The Geology of Indonesia/Papua

Papua is the eastern-most province of Indonesia and is located on the western half of the island of New Guinea. The island of New Guinea is the second largest island in the world and includes Puncak Jaya (4884 m), the highest peak in Southeast Asia and Australia. Traditionally, the outline of the island of New Guinea has been described as similar to a bird flying westward with open mouth (Fig. 1). As a result, the island has been geographically divided into the Bird’s Head, Neck, Body and Tail regions. The geologic of the Irian Jaya is very complex, involving interaction between two plates; the Australian and Pacific plates (Fig. 1 ). Most of the Cenozoic tectonic evolution of New Guinea is the result of oblique convergence between the Indo-Australian and Pacific plates (Hamilton, 1979; Dow et al., 1988). New Guinea and the mountainous Central Range, is commonly cited as the type locality of an active oceanic island arc - continent collision (Dewey and Bird, 1970). The Central Range is a 1300 km long, 150 km-wide belt with rugged topography and numerous peaks over 3000 m in elevation. Most of the range is composed of folded and faulted Mesozoic and Cenozoic strata that was deposited on the Australian passive continental margin. The purpose of this paper is to make a general summary of the geology and tectonics of Irian Jaya from various sources of published information.


In general, from the north to the south, the geology of Irian Jaya can be divided into three broad geologic provinces: Continental, Oceanic and Transitional provinces. Every geologic province has its own characteristic in stratigraphic, magmatic and tectonics history. The Continental province consists of sediments as apart of Australian craton. The Oceanic province consists of ophiolite rocks and island-arc volcanics complex as a part of Pacific plate. The transition province is a zone that consists of highly deformed and regional metamorphic rocks as a product of interaction between two plates. However, this relatively simple zonation doesn’t apply perfectly to the Bird’s Head and Bird’s Neck. Some authors believe that this region consists of widely diverse terrains where their origin is not at the present position. Therefore, their geologic history differs and should be separated from Bird’s Body (Pieters et al, 1983; Pigram and Davies, 1987).

The central portion of New Guinea (the Bird’s Body) can be divided into four lithotectonic provinces (Fig. 1.2): the New Guinea Foreland/foreland basin (Arafura Platform), the Central Range fold-and-thrust belt, a metamorphic (the Ruffaer Metamorphic Belt) and ophiolite belt, and a collided Melanesian arc island arc complex (the Meervlakte depression/north coast basin and the Mamberamo Thrust Belt).

The New Guinea Foreland (Arafura Platform) consists of the Arafura Sea and southern coastal plain of Irian Jaya which lie on Australia continental crust. The stratigraphy of the platform is composed mostly of unmetamorphosed marine and nonmarine Pliocene and Holocene siliciclastic sedimentary rocks which are underlain by the Cenozoic carbonate and Mesozoic siliciclastic strata deposited on the northern passive margin of Australia (Dow and Sukamto, 1984a, b). Foreland thrust and the Central Range Fold-and-Thrust Belt are defined as the New Guinea Mobile Belt (Dow et al., 1988). The Central Range is an orogenic belt that stretches 1300 km from Irian Jaya to the Papuan Peninsula. The 150 km-wide belt has rugged topography and numerous peaks over 3000 m in elevation. This wide zone is a south verging fold-and-thrust belt that largely involves Paleozoic to Tertiary rocks of the Australian continent. The amount of shortening across this belt in Irian Jaya is unknown, but palinspatic reconstruction of cross-section across the thrust belt in Papua New Guinea shows approximately 75 km north-south directed shortening (Hobson, 1986).

The Ruffaer Metamorphic Belt is a 50 km wide zone of highly-deformed generally low-temperature (<300°) metamorphic rocks which is bounded on the north by Irian Jaya Ophiolite Belt and on the south by deformed, but unmetamorphosed, passive margin strata (Dow et al., 1988; Nash et al., 1993; Warren, 1995; Weiland, 1999). The Irian Jaya Ophiolite Belt is separated from the Ruffaer Metamorphic Belt by a series of faults and is covered by alluvium of the Meervlakte depression. The suture separating the rocks from two different plates may be at the boundary between the Ruffaer Metamorphic Belt and the Irian Jaya Ophiolite Belt. The Derewo fault zone was mapped primarily with aerial photographs and satellite images as the boundary between the Ruffaer Metamorphic Belt and unmetamorphosed strata in the Irian fold belt by Dow et al. (1986). However, recent mapping shows that the transition from the Ruffaer Metamorphic Belt to unmetamorphosed strata is gradational from north to south and that from near 137° to 138°E longitude (Warren, 1995). Thus, the mapped DFZ does not correspond to the southern boundary of the metamorphic rock as shown on previous maps.

The most northern orogenic belt in Irian Jaya is a poorly exposed, complex zone involving oceanic rock from a collided Melanesian island arc built into the Pacific Plate. This belt consists of the Meervlakte (lake plain) depression and the Mamberamo Thrust-and-Fold Belt (MTFB). The Meervlakte is an intramontane basin and this basin has been actively subsiding from Middle Miocene to the present in which the rate of subsidence is greater than the rate of sedimentation (Dow et al., 1988). MTFB is a 200-km wide, NW-trending zone of convergent deformation, largely within the Melanesian arc terrane, that began in the Pliocene and is still active (Dow and Sukamto, 1984a, b; Dow et al., 1988).


The details of the Cenozoic tectonic evolution of New Guinea are the subject of some debate. The most commonly published scenario is the subduction polarity reversal (or arc reversal) model which entails movement of the Australian continental crust and mantle into a northward dipping subduction zone, followed by collision and initiation of southward subduction of the Pacific plate at the New Guinea Trench (Dewey and Bird, 1970; Hamilton, 1979; Milsom, 1985, Dow et al., 1988; Katili, 1991). A second model proposed to explain relationships in eastern New Guinea shows the island underlain by a doubly-dipping slab of oceanic lithosphere (“zippering” model), which is the westward continuation of the subducted Solomon plate (Ripper and McCue, 1983; Cooper and Taylor, 1987). A third model is similar to the first, but the subducted Australian plate is simply dipping vertical without a reversal in subduction direction (Johnson and Jaques, 1980). Both of these later models require significant oblique convergence across most of the island.

All of these authors agree that southern New Guinea is underlain by the passive northern margin of the Australian continent which was thickly blanketed from the early Mesozoic by siliciclastics grading into Cenozoic carbonate strata. Most of these authors argue for one major collisional event with an oceanic island arc. Based on the change from carbonate sedimentation to widespread clastic sedimentation derived from orogenic uplifts, the collision apparently began in the late Miocene (Visser and Hermes, 1966; Dow and Sukamto, 1984a; Dow et al., 1988). However, based on the metamorphic age of rocks in Papua New Guinea and the island arc terrane, some workers conclude that the collision began in the early Oligocene (Pigram et al., 1989; Davies, 1990). To account for these relationships, Dow et al. (1988) propose that New Guinea is the product of two distinct island arc-continent collisions: one during the Oligocene and a second during the Miocene (Melanesian Orogeny). Quarles van Ufford (1996) proposes that the island is the site of two orogenic events that are spatially and temporally distinct. The earlier (Eocene-Oligocene) event, termed the “Peninsular orogeny,” was restricted to the Bird's Tail region of easternmost New Guinea. The generation and erosion of a significant landmass is recorded in Oligocene and younger clastic sedimentation in the Aure Trough. The “Central Range orogeny,” on the other hand, was an event that began in the middle Miocene and generated the present width of the island and caused widespread clastic sedimentation. Importantly, Quarles van Ufford (1996) divided the Central Range orogeny into pre-collisional and collisional stages. The pre-collisional stage is related to the bulldozing and metamorphism of passive margin sediments in a northward-dipping subduction zone. The collisional stage only occurs when buoyant Australian lithosphere actually jams the subduction zone (Cloos, 1993) and crystalline continental basement becomes involved in the deformation. Collisional delamination is proposed as the tectonic process that occurred within the subducting Australian lithosphere between 7 to 3 Ma. Besides involvement of crystalline basement, this process causes late-stage igneous activity and a vertical mountain uplift of 1 to 2 km. This process culminated with the initiation of the E-W trending, left-lateral strike-slip that dominates the recent tectonics of western New Guinea.


Fig. 14.2. Papua Stratigraphy, Sapiie, 2000

The details and complete stratigraphic nomenclatures of Irian Jaya are best described in the geological report and map from Dow et al (1988) published by GRDC Bandung. In this paper the stratigraphic of the region will be compiled, generalized and simplified from various publications. The stratigraphy of the Australian Craton mainly known from exposed rocks in the deformed northern margin of Australian Craton resulted from regional uplift during the Central Range Orogeny and drill wells when exploring for oil in the Bird’s Head platform. The simplified stratigraphic nomenclatures of the rock from Australian Craton are summarized in Fig.. This diagram shows regional stratigraphic co-relation from the Bird’s Head in the NNW to the Bird’s Body in the ESE.


The distribution of Paleozoic rocks in Irian Jaya is poor due to the lack of exposures. Therefore it is difficult to generate regional correlation of this strata throughout the region. In addition, a lot of older strata have been regionally metamorphosed. There are several locations that expose Paleozoic strata in Irian Jaya. The largest block of the rocks exposed In the mountain of the northeastern Bird’s Head is known as Kemum High. On the other hand, the best out crop of unmetamorphosed section of Paleozoic strata is exposed along the Gunung Bijih Mining Access (GBMA) at the southwestern Central Range. The southwestern Central Range forms a simple north-dipping homocline, approximately consisting of 30-km-wide, 18-km-thick sedimentary sequence. The GBMA contains road cuts, which expose the most stratigraphically continuous section in all of Irian Jaya (Sapiie et al, 1999). In the Bird’s Head region, the oldest strata known are the thick Kemoem Formation which consist mostly of slate, phyllitic and minor quartzite. In the Bird’s Head this formation is intruded by Carboniferous biotite granite (Melaiurna Granite). Therefore, the metamorphism is interpreted as occurring in Devonian to early Carbonaceous (Pigram et al, 1982a). However, in the northeast Kemum high, Triassic granitic rocks intrude the lower Paleozoic (i.e. Anggi Granite, Kwatisore, Netoni Intrusive Complex) (Dow et al, 1988). Kemoem Formation is overlain unconformably by the Aifam Group.

Aifam Group is used for describing a group of mappable rocks consisting of shallow-water shelf sediments in the lower part and grading upward to fluviodeltaic environment. The Aifam Group is best known from the northern margin of the Bird’s Head Platform and in this region the group has been subdivided into three formations: Aimau Formation, Aifat Mudstone and Ainim Formation (Dow et al, 1988). The Aifam Groups occurs widely in the Bird’s Head region where it appears to be unmetamorphosed. This group is more strongly deformed and metamorphosed in the Bird’s Neck area. In the Bintuni area, the Tipuma Formation is unconformably overlying the Aifam Group (Biantoro and Luthfi, 1999).

In the Central Range (Bird’s Body), Awigatoh Formation is the oldest rock known in Irian Jaya. This rock unit was named Awigatoh Formation by Bar et al (1961) and Visser and Hermes (1962), and later named as Nerewip Formation on Timika Sheet Map by Parris (1994). This rock is exposed in the Awigatoh Mountain close to the border, the core of Mapenduma and Digul Range anticlines (Paris, 1994; Granath and Argakosoemah, 1989). In the Mapenduma anticline, the formation is exposed along Baidu and Nerewip Rivers west of GBMA (Quarles van Ufford, 1996). The formation consists of metabasalt, metavolcanic with minor limestone, shale and siltstone. Based on limited field observation, this formation appears to be overlain disconformably by Kariem Formation.

Kariem Formation along GBMA has been given different names, such as Kemoem Formation based on lithologic correlation with the northeastern Bird’s Head Region (Martodjojo et al., 1975), or a new name Otomona Formation on the Timika Map Sheet (Parris, 1994). Originally, Kariem Formation was the description for a group of sedimentary rock exposed in the Kariem river at Eastern Irian Jaya (Bar et al., 1961; Visser and Hermes, 1962). Lithologically this formation consists of fine-grained quartzose turbidite. In the northeastern Bird’s Head Region, this sediment were metamorphosed, intruded by granite, eroded before late Carboniferous, and overlain unconformably by the Aifam Group (Dow et al, 1988). The age of Kariem Formation is interpreted as Precambrian or Early Paleozoic. This interpretation is based on stratigraphic position that is below the Silurian and Devonian Modio Formation and from the result of reset age of zircon fission-track (ZFT) showing age of 650±63 Ma (Quarles van Ufford, 1996). At the GBMA, the relationship between Kariem Formation with the overlying Tuaba Formation is inferred to be disconformable (Quarles van Ufford, 1996).

Tuaba Formation was named by Pieters et al. (1983) for describing unit exposed in the Tuaba river. Tuaba Formation is composed of thick bedded of coarse- to medium-grained quartz sandstone with interbedded conglomerate and shale. The age of Tuaba Formation is constrained as Precambrian or early Paleozoic. The formation is stratigraphically below the Silurian to Devonian Modio Formation. However, along the GBMA this formation is in the fault contact (Hannekam Fault) with Modio Formation. Therefore the nature of the contact is unknown (Sapiie et al, 1999).

Modio Formation is previously named Modio Dolomite (Pigram and Panggabean, 1983; Dow et al., 1988). Quarles van Ufford (1996) renamed this rock unit from Modio Dolomite to Modio Formation to incorporate the siliciclastic member in the upper part. This formation is divided into two members. The lower A Member is dominated by carbonate specifically well-bedded stromatolitic dolostone. On the other hand, fine-grained clastic rocks consisting of bioturbated mudstone and siltstone dominate the upper B Member, fine-grained planar cross-bedded to horizontally laminated sandstone (Quarles van Ufford, 1996). Modio formation is interpreted as a transgressive sequence deposited from tidal to marine self. The age of Modio Formation is constrained as Silurian to Devonian based on Late Devonian (Frasnian) coral found and identified from the limestone in the Modio B member (Oliver et al., 1995). The upper contact with Aiduna is not well exposed and is interpreted to be disconformable (Quarles van Ufford, 1996).

Aiduna Formation was first named by Lehner et al. (1955) in the western part of the Waghete sheet as a part of the lower member of the Aifam Formation (Parris, 1994). In GBMA Martodjojo et al. (1975) placed this formation within Aifam Group as Aifam C Member of Visser and Hermes (1962). Pigram and Panggabean (1983) used Aiduna Formation in Waghete Sheet area because of difficulty of subdividing the Aifam Group. Parris (1994) on the Timika Sheet area preferred the use of Aiduna Formation replacing Aifam C Member, since he had already subdivided the lower Aifam into The Tuaba and Modio Formation. Aiduna Formation is characterized by well-bedded coal bearing silisiclastic rocks. This formation is interpreted to have been deposited in fluvial to deltaic environment (Visser and Hermes, 1962; Dow et al., 1988). However, the presence of Brachiopods indicates that some of the Aiduna Formation was deposited in marine environment or perhaps lagoonal area (Martodjojo et al., 1975; Parris, 1994; Quarles van Ufford, 1996). The age of Aiduna Formation is constrained by Brachiopods fossils as Permian (Martodjojo et al., 1975) and by plant flora as Late Permian (Quarles van Ufford, 1996). The contact with overlying Tipuma Formation is conformable.


The Tipuma Formation is widespread in Irian Jaya, extending from the northwest Bird’s Head to the east near the border. Visser and Hermes (1962) were the first who formally gave the name Tipuma Formation for the rock unit derived from Kembelangan No. 1 well in the Bird’s neck area. The Tipuma Formation is characterized by a distinctive red color with minor light green mottling. Tipuma Formation was deposited in fluvial environment during the period of continental rifting (Pigram and Panggabean, 1983). Field observation indicates that the thickness of the formation changes rapidly along the strike (Quarles van Ufford, 1994). This evidence is interpreted to be representing a horst and graben depositional topography resulting from active extension. The age of the Tipuma Formation is solely constrained by its stratigraphic position, that is, Triassic to Early Jurassic. Pigram and Panggabean (1983) on Waghete Map Sheet suggested that the contact between the Tipuma Formation and the overlying Kembelangan Group is unconformable (post-breakup unconformity). KEMBELANGAN GROUP

The Kembelangan Group is recognized from the Bird’s Head to the Arafura Platform and is a regionally extensive unit deposited on the northern passive margin of the Australian continent during Mesozoic time (Visser and Hermes, 1962; Dow et al., 1988). Pigram and Panggabean (1983) divided the Kembelangan Group into four formations, the Kopai Formation, the Woniwogi Sandstone, the Piniya Mudstone and the Ekmai Sandstone. In the Bird’s Head region, the Kembelangan Group cannot be subdivided into four formations. The upper part of this group is known as the Jass Formation (Dow et al, 1988). The Jass Formation consists of quartz sandstone and calcareous mudstone. The Kembelangan Group consists of interlayer carbonaceous siltstone and mudstone in the lower section, and fine-grained glauconitic quartz sandstone and minor shale in the upper section. This group was deposited as a passive margin sequence conformably overlying the Triassic rift sequences of the Tipuma Formation (Dow et al., 1988; Parris, 1994). The contact with overlying Waripi Formation of New Guinea Limestone Group appears to be conformable. NEW GUINEA LIMESTONE GROUP (NGLG)

During the Cenozoic time, approximately at the Cretaceous and Cenozoic boundary, the island of New Guinea is characterized by carbonate deposition known as the New Guinea Limestone Group (NGLG). The NGLG overlies the Kembelangan Group as originally defined by Visser and Hermes (1962). In central Irian Jaya, The New Guinea Limestone Group is generally divided into four formations.

The basal unit is the Paleocene to Eocene Waripi Formation, which is composed of fossiliferous dolostone, quartz sandstone and minor limestone. The Waripi Formation was deposited in a shallow marine, high energy environment. This formation has gradational contact with the Yawee Limestone (undifferentiated NGLG in ) and Late Cretaceous Ekmai Sandstone (Pieters et al, 1983).

The Eocene Faumai Formation conformably overlies the Waripi Formation. This formation is composed of thick-bedded (up to 15 m) to massive foraminifera-rich limestone, marly limestone, dolostone and a few quartz-rich sandstone layers up to 5 m thick. The Faumai Formation was deposited in shallow marine, medium energy environment. The Early Oligocene Sirga Formation conformably overlies the Faumai Formation. This formation is composed of a foraminifera-bearing, coarse- to medium-grained quartz sandstone and siltstone that is locally pebble-rich. The Sirga Formation was deposited in a fluvial to shallow marine environment after period of non-deposition. This formation is the only silisiclastic formation deposited in the Irian Jaya region between the Eocene and Middle Miocene. Pigram and Panggabean (1983) name this formation as Adi Member. The Sirga Formation deposition is the result of the transgression that followed the Oligocene sea-level fall as well as Oligocene orogenic activity in eastern New Guinea (Quarles van Ufford, 1994). The Imskin Formation is a pelagic limestone which consists of well-bedded carbonate mudstone, marl, chalk, chert and abundant pelagic foraminifera (Visser and Hermes, 1962; Koesoemadinata, 1978; Pieters et al, 1983). This formation represents deep-marine environment and grades upward into shallow-water carbonate. This formation ranges in age from Paleocene to middle Miocene (Pieters et al, 1983).

The Oligocene to Middle Miocene Kais Formation conformably overlies the Sirga Formation. This formation is composed primarily of foraminiferal limestone with interbedded marl, carbonaceous siltstone and coal. The Kais Formation was deposited on a medium- to low-energy carbonate shelf. Biostratigraphic analysis indicates the youngest strata to be ~ 15 Ma (Quarles van Ufford, 1996). In the Bird’s Head, the Kais Formation represents a reef complex comprising platform and patch reef facies. This formation is laterally equivalent to the Klamogun Limestone of Salawati basin. In addition, in the Salawati and Bintuni basins the Kais Formation partly interfingers and is conformably overlain by the Klasafet Formation (Dow et al, 1988).


The late Cenozoic sedimentation in the Australian continental basement is characterized by the kilometer-thick siliciclastic sequences overlying middle Miocene carbonate strata (Visser and Hermes, 1962; Dow et al, 1988). In the Irian Jaya region, three major formations are known and all of them are similar in terms of age and lithology (Pieters et al, 1983). These are the Klasaman, Steenkool and Buru Formations. They occur respectively in the Salawati and Bintuni Basins and in the southern part of the Central Range (Akimeugah and Iwur Basins). Locally, they are overlain by younger clastic sediments (i.e. Upa and Sele Conglomerates). In northern Irian Jaya, siliciclastic rocks occurred in the North Coast Basin (Meervlakte) in the early middle Miocene (Visser and Hermes, 1962; Dow et al, 1988). This rocks unit known as the Makats Formation covers the oceanic basement.


The detail stratigraphy of the Pacific plate (Oceanic province) is presented by Pieters et al (1983) and Dow et al (1988). In general, the Pacific rocks consist of mantle derived rocks, island-arc volcanis and shallow-marine sediments. The mantle derived rocks are exposed extensively along the Irian Jaya Ophiolite Belt (IOB), the Cyclop Mountain, Waigeo Island, north of Gauttier Mountain and as sliver blocks along the Sorong and Yapen Fault Zone (Dow et al, 1988). The IOB is approximately 400 km long and 50 km wide of east-west belt made of ultramafic, basic plutonic and high-grade basic metamorphic (Dow et al, 1988). The age of the IOB is unknown, but it is interpreted as Mesozoic based on metamorphic block.

The Auwewa Volcanic Group is the volcanic rocks of the Pacific plate (Dow et al, 1998). Originally, Visser and Hermes named it Auwewa Formation. However, because all the volcanic rocks have the same ages and is very similar in composition. They are all put together within one group. The rocks in the group are mostly the product of island-arc volcanism which are remarkably uniform (Dow et al, 1988). They are mostly characterized by basic composition. Throughout Irian Jaya they range in age from Paleogene to Early Miocene (Visser and Hermes, 1962).

The sediments in Pacific plate are characterized by shallow-marine carbonate with interbedded terrigenous sediments derived from island-arc and less commonly from the mantle rocks. This unit, named as Holandia Formation (Visser and Hermes, 1962) and Dow et al (1988), was raised in status of unit to a group. The group is extensively distributed at Waigeo, Biak, and Yapen islands and on the flank of the Cyclop Mountains. The age of this group ranges from Early Miocene to Pliocene.


The convergence between Australian and Pacific plates generated rocks of within the zone of deformation. This group of rocks is classified as transitional zone, which consists mostly of metamorphic rocks. The metamorphic rocks formed continuous belt (>1000km) from Irian Jaya to Papua New Guinea.

In Irian Jaya low T regionally metamorphosed rocks are exposed along the Weyland Range and the northern flank of the Central Range. Historically, several different name have been proposed for these metamorphic belt such as Derewo Metamorphics (Pieters et al., 1983), Derewo Metamorphic Belt (Nash et al., 1993) and Ruffaer Metamorphic Belt (Dow et al., 1988, Warren, 1995). This belt generally consist of low T ( 300-350C and 5-8 kb) metapelites derived from the Mesozoic passive-margin of the Australian continent. The fact that the metamorphic rocks from Weyland Range recorded higher T (straulite-biotite-garnet; Dow et al., 1988) than in the Central is interpreted as result of the intrusion of the Utawa batholith (Warren, 1995). Isotopic ages from metapelitic rocks in Irian Jaya and Papua New Guinea record a late Oligocene to Earliest Miocene regional metamorphic event (Weiland, 1999). The contact separating this belt and the Irian Fold Belt is gradational (Warren, 1995).


Only little structural geologic evidence has been found prior to Late Miocene tectonic event (Melanesian orogeny by Dow et al., 1988) in Irian Jaya. The evidence of the oldest structures are recorded in the Paleozoic sections. However, the exposure of this group are limited, therefore the knowledge of the Paleozoic tectonic is also very poor. Most of the structural features in the island today are product of Late Miocene arc-continent collision. Later tectonic event (<4 Ma) is reactivated some of the older structures becoming dominated left-lateral strike-slip faults (Sapiie et al., 1999). In general, the structural pattern in Irian Jaya can be divided in three major structural domains: Bird’s Head, Neck and Body. In the Bird’s body, W to NW trending structures is dominant throughout the Central Range. This W-NW belt known as New Guinea Mobile Belt, a 300-km wide zone of continuos belt from Papua Nugini to Irian Jaya (Dow et al., 1988). The New Guinea Mobile Belt is terminated by EW trending continental strike-slip fault, the Tarera-Aiduna Fault Zone (TAFZ), at Bird’s neck. The structures in the Bird’s neck are dominated by N-NW trending fold known as Lenguru Fold Belt (LFB). This fold belt is terminated at the Kemum high in the Bird’s head region. In this region, the majority structures are dominated by EW trending fault system.

Sapiie et al. (1999) proposed that the piece of lithosphere north of New Guinea is moving as yet another distinct kinematic entity, the Caroline plate. This microplate is moving nearly, but not exactly, with the Pacific plate. In western New Guinea, the dominant effect of this interaction is to cause left-lateral transform motion. This motion starts in the Bismarck Sea, comes on land forming the 250 km long Bewani-Torricelli fault zone, makes a 200 km long right step with convergent deformation along the Mamberamo thrust and fold belt, returns to left-lateral offset along the Yapen fault zone. Further west, a 300 km long left step forms the divergent Waipona Trough which links to the Tarera-Aiduna fault zone which in turn extends westward to the Banda Sea. The Tarera-Aiduna fault zone separating the west directed underthrusting at the Seram Trough from the northward subduction at the Timor/Aru Trench. The restraining bend forming the Mamberamo thrust and fold belt is well defined from the regional seismicity. As earthquakes occur down to depths approaching 150 km, at least 200 km of convergence has occurred in this area and hence a similar magnitude of strike-slip is indicated along the Bewani-Torricelli and Yapen fault zones. Near Biak Island, convergence probably only recently began as earthquakes are shallow and there is no evidence of significant recent deformation in the nearby New Guinea Trench (Milson et al. 1992). In western New Guinea, significant west-directed convergence has been accommodated at the Seram Trough, which has seismicity extending down to depths of 100 km.

The major releasing bend in the system is the little studied Waipona Trough, a major depocenter since the Pliocene (Dow et al., 1988). Perhaps 50 to 100 km of lithospheric extension could have been centered on this zone. Based upon seismicity, active divergence is concentrated along the east coast of the Bird’s Head, cutting across the Lengguru fold belt. A southern extension of this divergent motion is the Aru Trough, which marks the rifting of the edge of the Australian plate. Total extension in either of these areas is small, perhaps 10 km. Because they do not appear to be long-lived phenomena, it is possible that underthrusting in the Biak area and extension along the back of the Bird's Head are very recent developments in the area.