The Geology of Indonesia/Banda Arc

The Banda Arc, a west facing horse-shoe shaped arc in eastern Indonesia, defines the locus of three converging and colliding major plates, the Indo-Australia Plate, the Pacific Plate, and the Eurasian Plate. Splinters of the Mesozoic southern Tethyan crust now form the base of the Banda Sea. On the surrounding islands, Asian affinity forearc nappes of the Banda Terrane, which may include ultramafic rocks and young mafic volcanics, structurally overlie Australian affinity passive margin units accreted to the upper plate. Banda Terrane metamorphic rocks yield Eocene cooling ages and record events happening in the Great Indonesian Arc (GIA) before opening of the Banda Sea and Miocene to present arc-continent collision. The GIA was an active continental arc along the southern margin of Sundaland from Jurassic to Oligocene (Carter et al., 1976; Haile et al., 1979; Earle, 1983; Harris, 2006). Australian affinity syn-collisional metamorphic rocks are exposed along the North Coast of Timor and are known as the Aileu Complex, which have protolith ages of Permian to Jurassic. Metamorphic cooling ages of these rocks range from 12-3 Ma.

Banda Arc tectonic map

From the inner (Banda Sea) side to the outer (Foreland Basin) side the following belts has been distinguished in the Outer Banda Arc geology by de Smet (1999, Figure 1): 1. A forearc belt, which is in general blocky and narrow, 2. A collision-related metamorphic belt, composed of low to high grade metamorphosed rocks, 3. A thrust and fold belt dominated by Permo-Triassic and Jurassic sediments of Australian continental margin origin. 4. A thrust and fold belt dominated by Late Mesozoic and Tertiary deep water sediments 5. A belt of uplifted Late Neogene basins.

11.1. TECTONICS edit

Banda Arc cross section

The origin of the arc has been a subject of continuing controversy and proposals then be grouped into three: 1. The arc was formed by just a simple 180° counterclockwise rotation of an originally straighter east-west trending arc, by which its northern part has rotated relative to the southern part (Katili, 1975; Carrey, 1976; Audley-Charles, 1972; Carter et al., 1976). 2. The arc has had achieved its present curvature at least from Late Cretaceous (Norvick, 1979). 3. The arc was formed by a pieces of northern margin of the Australian continental masses leading to the enclosure of the Banda Sea oceanic basin (Silver et al., 1985; Bowin et al., 1980; Lee & McCabe, 1986; Lapouille et al., 1985; Pigram & Panggabean, 1983, 1984; Hartono, 1990a).

Richardson & Blundell (1996) summarized the proposed structural models into three main groups, resulting mainly from near-surface observations on Timor:

1. The imbricate model (Fitch & Hamilton, 1974; Hamilton, 1979) is based largely on marine geological and geophysical data (e.g. von der Borch, 1979; Silver et al., 1983; Karig et al., 1987). In this model, Timor is interpreted as an accumulation of chaotic material imbricated against the hanging wall of a subduction trench, the Timor trough, and essentially forms a large accretionary prism.

2. The overthrust model is probably the oldest model in which Timor was interpreted in terms of Alpine-style thrust sheets (e.g. Wanner, 1913). This model was dominantly based on surface geology where overthrust sheets of the Timor allochthon are well exposed. Subsequent workers (Carter et al., 1976; Barber et al., 1977; Barber, 1979; Haile et al., 1979.; Brown & Earle, 1983; Audley-Charles, 1981, 1986a, b; Price & Audley-Charles, 1983, 1987; Harris, 2011; Audley-Charles & Harris, 1990) have made a clear distinction between allochthonous units of non-Australian origin and paraautochthonous units derived from the Australian continent.

3. The rebound model (Chamalaun & Grady, 1978) suggests that the Australian continental margin entered a subduction zone in the vicinity of the wetar Strait. Subsequently, the oceanic lithosphere detached from the continental part, resulting in the uplift of Timor by isostatic rebound on steep faults. This model is based on the assumption that convergence has stopped, which is not consistent with GPS measurements of at least 26 mm/a of convergence across the Timor Trough and fault plane solutions that show active convergence along the subduction interface and near the suture zone. There is no evidence of slab tear or subduction polarity reversal.

Closely linked with the origin of the Banda Arc is speculations around the age and mode of occurrence of the Banda Sea, which origin remains controversial i.e. whether it formed by back-arc spreading (Barber, 1981; Carter et al., 1976; Hamilton, 1988; Nishimura & Suparka, 1986) or whether it represents a trapped oceanic lithosphere of Indian Ocean affinity (Bowin et al., 1980; Lee & McCabe, 1986; Pigram & Panggabean, 1983; Silver et al. 1985).

As originally proposed by Earle (unpublished PhD, 1981), recent studies in Timor (Sopaheluwakan, 1990a, b, 1991; Sopaheluwakan et al., 1989; Helmers & Sopaheluwakan, 1989, Helmers et al., 1989; Stanley and Harris, 2009; Harris 2011) and Buru-Seram microplate (Linthout et al., 1989, 1991; Sopaheluwakan, 1993; Sopaheluwakan et al., 1992) have provided an alternative interpretation of the origin of the metamorphic basements in the Banda Arc, emphasizing the important role of lhezolitic ultramafic rocks, which was overlooked by previous models.

Plate tectonics theory which predict the presence of an accretionary wedge, composed of distal sediments scraped from the ocean floor at the inner side of the arc. At this place, however, de Smet (1999) found continental basement rocks and Early Mesozoic continental sediments. The situation is explained as the result of an inversion process of the Australian continental margin during the Neogene plate collision between Australia and the Banda subduction zone. De Smet (1999) concludes that the Outer Banda Arc is not an accretionary complex, but instead the compressed northern rim of the Australian continent. The load of Australian continental margin sediments that accumulated during the Mesozoic and Tertiary is pushed up on the back of crustal blocks of continental basement to form the present islands and mountains of the Outer Banda Arc. However, no evidence for involvement of Australian basement is found. All the ages of metamorphic rocks in Timor is pre-Jurassic (Banda Terrane) or Miocene-Pliocene).


11.2.1. TIMOR

The stratigraphy of Timor is very complicated due to structural complexity. In general the stratigraphy of Timor is divided into three sequences, named Gondwana, Passive Margin and Syn-Orogenic Sequences, ranging from Permian to Pleistocene. This chapter is a shortened version of Sawyer et al. (1992) BASEMENT ROCK

The affinity of metamorphic rocks that outcrop on Timor is not well understood. Schist, phyllites, amphibolites, and associated serpentinites of the Mutis-Lolotoi Complex represent Late Jurassic to Early Cretaceous arc-trench lithosphere (e.g. Earle, 1979; Harris, 2006). No pre-late Carboniferous rocks are found in Timor. The oldest metamorphic rocks are Eocene (Standley and Harris, 2009). Earle (1981) showed that heavy mineral assemblages identified in the metamorphic complexes of West Timor are different to those in the Permo-Triassic formations on Timor, and incompatible with derivation from continental basement. KEKNENO SEQUENCE

Kekneno Sequence ranges from the Early Permian to Middle Jurassic with an Upper Jurassic hiatus. Units include the Atahoc and Cribas Fm., Triassic Niof, Aitutu and Babulu Fm., and the Jurassic Wai Luli Fm. On strictly stratigraphic basis, the Permian age Maubisse Fm. can be linked to the Kekneno Sequence. This, however, contrasts with structural observations, which indicates a need for tectonic separation (Audley-Charles, 1968). Sawyer et al., tentatively introduce the term “Tethys Margin nappe” to reflect the disparity in occurrences of the Maubisse. MAUBISSE FORMATION

The Maubisse Fm. consists primarily of Early through Late Permian limestone and extrusive igneous members that are the oldest known rocks in West Timor (de Roever, 1940; Audley-Charles, 1968). Comparison of brachiopod assemblages led to an interpretation that the Maubisse was formed as part of Gondwana (Bird and Cook, 1991). The most common Maubisse lithologies are red to purple color biocalcarenites, packstones and boundstones rich in debris of coral, crinoids, bryozoans, brachiopods, cephalopods, and fusilinids. The matrix is typically recrystallized micrite with sparry cement replacing most bioclasts. Subordinate Maubisse facies include a massive white to grey limestone, well bedded micrites, rare interbedded clastics and channel-fill deposits lithologically equivalent to the Permian age Cribas Fm. ATAHOC FORMATION

The Atahoc fm of East Timor was dated on the basis of ammonoids as Early Permian Sakmarian (Bird, 1987). In West Timor, the Atahoc Fm. is not widely exposed, occurring only along the remote northwest coastline and in the north of West Timor. Atahoc basal contacts are not documented in West Timor. Bird (1987) described an upper contact consisting of amygdaloidal basalt between the Atahoc and overlying Cribas Fm. Large areas of West Timor assigned to the Permian Kekneno Series by Rosidi et al., (1981) were found to be Triassic age, and are reassigned to the Niof and Babulu Fm. Atahoc Fm sandstones are a fine-grained, moderately sorted arkose, with large pyritized woody fragments, and lithic fragments of Mutis / Lolotoi-equivalent (?) phyllite and shale (Sawyer et al., 1992). CRIBAS FORMATION

A classification of the Early Permian Cribas Formation established in East Timor by Audley-Charles (1968). Bird (1987) described five major facies with laterally continuous, sharp bed boundaries consisting of varicolored sands, silts, black shales, and bioclastic limestones and a thickness greater than 400 meters. This formation are exposed only in portions of the Northern Range. Sandstones are classified from petrographic analysis as bimodal, fine to coarse-grained feldspathic litharenites. The provenance was proximal to basic igneous rocks. A shallow shelf environment of deposition is in agreement with Bird’s (1987) identification of Atomodesme communities that represent temperature to subtropical waters in depths of 20-50 m. Sedimentary structures indicate that turbidity currents were a common mode of sediment transport. NIOF FORMATION

The Early to Middle Triassic Niof Formation bedding contacts are typically sharp and exhibits numerous sedimentary structures. Intraformational growth faulting, large-scale slumps and other syn-sedimentary structures are common. Dominant lithologies include laminated, thin-bedded claystones, red, grey, black and brown shales with rip-up clasts, siltstones, sandstones that fit a greywacke field classification, carbonate mudstones, and brittle limestones. In the northern range, nort of Atambua, Sawyer et al. (1992) observed a series of small exposures tentatively assigned to the Niof Fm. These consist of subrounded to angular conglomerates with clasts of Maubisse limestones, trachyte with sanidine phenocryst, intermediate volcanics, andesite, and granite. Microlitic matrix texture is composed of siltstone and micaceous sandstones. The section is in faulted contact with a fine-grained, indurated volcanic breccia and to the north, serpentinized peridotite of the Maubisse Fm. Atapupu facies. AITUTU FORMATION

The dominant lithologies of the Aitutu Formation are white to occasionally pink limestone with interbedded carbonate mudstones that vary from pale to grey to black color. Sometimes rounded chert precipitated within limestones. In outcrop, units are very well bedde3d on the order of 45-60 cm, with sharp, planar contacts. Along bedding planes, macrofauna Halobia, Daonella, Monotis, various ammonites and other fossil fragments are common. The Aitutu outcrops in the Northern Range and contacts are diachronous and in part graditional with Niof and Babulu fm., in agreement with Cook (1986). The upper contact with the Lower Jurassic Wai Luli Fm. is synchronous, and recognizable from a pronounced lithologic charge to dominantly shale. The Aitutu environment of deposition was predominantly open marine, outer shelf to slope, and probably distal to both the Niof and Babutu Fm. BABULU FORMATION

Lithologies of the Babulu Fm. consist of interbedded shale and silt, sand stringers, and massive sandstones of the Lapunif member (Gianni, 1971; Cook, 1986). Contacts are distinct but undulatory. Uppermost Babulu lithologies may consist of brittle, cross-bedded limestone and light grey calcareous shales that are similar to parts of the aitutu Fm. In outcrop, weathered surfaces of the Lapunuf member always appear tan to buff colored with fresh surfaces a salt and pepper gray color. Bedding planes within massive sandstones are typically 60 cm to 3 meters in thickness. Sedimentary structures are very common and include current bedding, biogenetic features, and mud cracks. Bedding plane surface contain brachiopods and small ammonites, abundant oriented plant fragments, mud cracks, sole marking, and trace fossils of the Nereites facies (Cook, 1986). Sawyer et al.’s date largely agree with Cook’s (1986), and range from Early Middle Norian to latest Carnian. Transitional Babulu and Niof Formation shales, however, all yields, a Ladinian age. The Babulu Fm. was deposited as a massive regressive sand wedge at the end of a major Triassic sea level transgression. We interpret the paleoenvironment as fluctuating between proximal nearshore to shelf-slope break, with deposition from both deltaic progradation and turbidity current. WAI LULI FORMATION

Homogeneous dark grey colored claystone and shale with interbedded organic-rich limestones, calcilutites, and siltstones dominate the Jurassic age Wai Luli Formation. The siltstones are always a homogeneous dark grey color on unweathered. Claystones and limestones are either pastel blue, yellow or green on fresh surfaces are. Organic –rich limestones exhibit millimeter color banding and laminations, and branching trace fossils that along with filled burrows, produce a characteristic mottled grey appearance. Interbedded units are always less than 60 cm thick, with sharp contacts that commonly contain bedding-parallel assemblages of ammonites (Sawyer et al, 1992). Thicknesses of sections in East Timor were estimated by Audley Charles (1968) at between 800-1000 meters.

Sawyer et al. (1996) dated some samples, all of which range from the Early to Middle Jurassic Hettangian to Callovian. Depositional environment was shallow inner shelf to middle shelf. The overlying Early Cretaceous Nakfunu Fm follows a nondepositional hiatus that encompassed the upper Callovian through Tithonian. Upper Wai Luli contact has yet been described between the Wai Luli and younger Oe Baat Formation lithologies.

In contrast to the Permian and Triassic age units of the Kekneno Sequence, the Wai Luli has not been observed in the Northern Range. This may result from erosion or perhaps because Wai Luli overpressured shales were the site of a major decollment between the Kekneno Sequence and upper thrust sheets. The latter implies that the Wai Formation was tectonically incorporated into the Sonnebait melange (Harris et al., in prep). KOLBANO SEQUENCE

Kolbano Sequence lithologies range from Upper Jurassic Tithonian to Lower Pliocene stage. Formations include the Tithonian to Berriasian Oe Baat, Early Cretaceous Nakfunu, Cretaceous Menu, and Tertiary Ofu (Harris 2011). The succession is punctuated by four major hiatuses or condensed section that occur in the: 1) Middle Cretaceous Albian trough Turonian, 2) Early Paleocene, 3) Oligocene and extends locally into the Lower Miocene, and 4) post-Early Pliocene. Exposures of Kolbano Sequence are rare north of the Southern Range due to erosion, removal by thrusting, subduction beneath the Banda Terrane, or a distal facies change to the Palelo Group of the Banda Terrane that has since been tectonically shortened (Sawyer et al, 1992). Kolbano Sequence lithologies in the Southern Range form a hinterland (north) dipping leading imbricate fan composed of frontally accreted Australian margin material (e.g. Harris, 1991; 2011). The unit is bounded to the north by tectonic contacts with the Kekneno Sequence from which the decollment level rise progressively from the footwall thrust cutoff ramp of the Boti-Merah thrust to within the upper part of the Kolbano Sequence by the southern coast. OE BAAT FORMATION

The massive sandstone facies of Oe Baat Fm. has rare bedding surfaces, but when observed consist of alternating silts and sandstones. The base of the section consists of brown to black siltstone and shale with limonite-encrusted silt nodules. The shale yielded a Late Jurassic Tithonian age (Sawyer, 1992). The environment of deposition of the bedded, glauconitic facies was shallow shelf, whereas the massive sandstone facies is less marine. Charlton (1987) concluded that the abundance of glauconite and phosphatic minerals implied proximity to a shallow shelf in an area of upwelling. Massive member textural maturity, feldspar composition, and subangularity of quartz grains indicated proximity to an uplifted sedimentary and continental granitic or gneissic source provenance. Sawyer et al. (1992) and Charlton (1987) suggest 480 meters thickness of this formation. Two upper formation contact relationships were observed by Sawyer et al. (1992). The first is a disconformity between the Oe Baat massive facies and Eocene age Ofu Formation Boti member. The second consists of Oe Baat overlain by Zanclian age Viqueque Formation. Although a major hiatus exist between units, structural dips are concordant. This implies that at least part of the section was eroded previous to deposition of Tertiary lithologies, wit a post Zanclian folding event responsible for the present antiformal structure. NAKFUNU FORMATION

Nakfunu Fm. lithologies consist of radiolareites, claystones, calcilutites, interbedded shales, and less commonly, calcarenites, wackestones and packstones. One distinguishing feature of the Nakfunu is bedding that consistently varies in thickness from only 3 to 30 cm with sharp, flat, planar to undulatory contacts. Shale units can be interbedded or massive. Black colored ferro-manganese nodules are common in outcrop. Measured sections indicate an average formational thickness of 500 m. Sawyer et al (1992) biostratigraphic results indicate Early Cretaceous ages in essentially two age range populations; Berriasian to Aptian, and Hauterivian to Aptian, with a major nondepositional hiatus between the Albian and the Turonian. Age relationships and allogenic composition indicate that the base of the Nakfunu is equivalent to the bedded, glauconitic facies of the Oe Baat Formation. Lithology was nearly identical to the Oe Baat, consisting of very fine grained, moderately sorted sublitharenite sandstone. The Nakfunu Fm. was deposited during a gradual Early Cretacous eustatic sea level rise that culminated in the Albian. The depositional environment was probably a starved distal continental rise. Abyssal water depths near or below carbon compensation depth are presumed by he occurrence of manganese, palynoflora species, and a lack of terrestrial phytoclasts, which suggests a low clastic input. Foraminifera were rare or absent in samples, however radiolarian were abundant and nannofossils common. Several samples contained reworked Middle to Late Triassic palynomorphs. MENU FORMATION

The Cretaceous Menu formation is lithologically similar to the Tertiary age Ofu Fm., and includes some of the units assigned by Charlton (1987) to the Boralalo Fm. In contrast to the nearly always massive lithologies of the Ofu, the Menu Fm. exhibits sharp, planar bedding where individual unit thickness is usually range from 6 to 60 cm. Limestones may contain 1 to 2 cm horizons and nodules of red chert, and often exhibit intense internal cleavage. Bedding planes show branching trace fossil casts up to 70 cm in length and 5 cm in width. Lithologic similarities between the Menu and Ofu Fm strongly suggest a stratigraphic contact. In the Noil Menu type area, the Early Cretaceous Nakfunu is imbricated with the Menu, but the original contact is suspected to be stratigraphic. Menu lithologies were deposited as distal calciturbidites in a deep marine environment similar to the Ofu Fm. (Sawyer et al, 1992) OFU FORMATION

The dominant lithology of the Ofu Formation is a hard, white to pink massive limestone that exhibits conchoidal to subconchoidal fracture, and is lustrous or porcelainous on fresh surfaces. In outcrop, units may contain very fine millimeter laminations and intense pressure solution cleavage, giving rise to calcite veining within stylolites, joints and fractures. The large size of abundant, inherited reworked Cretaceous and Paleogene forams suggests an origin from rapid downslope transportation by turbidity current processes. Like the Nakfunu and Menu, the Ofu Fm. was deposited in an essentially clastic starved, deep marine setting. Clastic preservation, however, indicates accumulation above the carbon compensation depth (CCD). This suggests either a relative water depth shift from lower rise in the Early Cretaceous to an upper rise or slope setting, or a favorable change in world ocean temperature or chemistry. VIQUEQUE SEQUENCE

The Viqueque Sequence consists largely of Plio-Pleistocene sediments of synorogenic “molasse-type” origin. The sequence includes the Viqueque Fm. and various melange units, although a genetic tie is not necessarily implied. Wells drilled in the Suai Basinof East Timor encountered up to 3,000 meters of Viqueque Sequence in contact over a folded Mutis / Lolotoi thrust sheet. VIQUEQUE FORMATION

The Viqueque Formation is an overall coarsening-upwards succession from chalks and calcilutites to sandstones capped by quaternary gravels and reef limestones. This formation occurs within the Central Basin, west and south of the Kolbano imbricated units, and perhaps into the Northern Range. It shows strong lithologic variations that reflect rapid proximal uplift and variations in the original depositional topography. Units are nearly always bedded, on the order of 10 cm or more. Contact at the base of the Viqueque Fm. are highly variable. Sawyer et al. (1992) observed angular contact over Triassic Aitutu, structurally conformable over eroded or nondepositional Oe Baat Fm., high-angle faulted with the Kolbano Sequence, drapped over klippen of the Banda and Tethys nappes, and in contact with Sonnebait and Bobonaro melange. Batu Putih member lithologies are primarily massive white calcilutites or chalks and light grey marls with common plant debris. Units are soft to firm, and bedding indistinct. Tuffaceous horizons are rare outside the type locality, although accessory vitric glass shards are common. Coarse bioclasts and clastic allogens occur where the unit interfingers with the Noele member. MELANGES

Throughout timor, several units can be described as melange or are easily confused with melange. Harris et al. (1998) distinguishes between a diapiric and reworked sedimentary deposit called the Bobonaro Scaly Clay, and a significantly more deformed melange of probable tectonic origin called the Sonnebait Melange. Bobonaro Scaly Clay occur where basal Viqueque Fm. is in contact with grey shales and common entrained cobble to boulder-size blocks, in the Oeleu diapir and within active diapirs of Pulau Semau, Oecussi and Halilukiuk. The extruded shales contain blocks and fauna as old as Middle Triassic and as young as Pleistocene. In contrast, the Sonnebait Melange appears to be a product of tectonic deformation. Shales are usually recrystalliezed, and associated or entrained blocks exhibit sheared contacts. In general, the size of tectonized blocks and ratio of blocks to melange matrix decreases across the island from north to south. BANDA TERRANE

The Banda Terrane is regard as a dismembered, high level nappe consisting of forearc basin and volcanic arc lithologies. The upwards shallowing sequence from ocean floor to continental shelf to reefal (Barber, 1978) begins with Mutis=Lolotoi equivalent metamorphic no older than Later Jurassic (e.g. Earle, 1981; Standley and Harris, 2009). The Palelo Series or Group unconformably overlies this oceanic lithosphere basement, and comprises a thick sequence of forearc clastics and volcanics that were deposited on the Asian plate prior to its collision with the continental margin of Australia (Earle, 1979, 1981, 1983).

11.3. TANIMBAR edit


The oldest rocks so far recognised on Tanimbar are of Triassic age found only in decimetre-sized blocks ejected from mud volcanoes. The Triassic rocks comprise yellow, brownish and grey coarse to fine grained sandstones which commonly show sedimentary structures including cross- bedding, ripples, sole and tool marks typical of turbidite sedimentation. Other sandstones show herringbone cross-bedding, possibly indicative of tidal deposition environments. The rocks are immature subarkosic micaceous sandstones, with grains subangular to subrounded and with variable degrees of sorting. They are petrographically reminiscent of Triassic sandstones in Timor (the Babulu Formation of Cook et al. 1989, Bird et al. 1989, Bird and Cook 1991). Palynomorphs indicate Middle – Late Triassic (Anisian – Rhaetian) and Early Jurassic ages for the sandstones (P. T. Corelab Indonesia, written communication 1987). Occasional coal fragments found in the mud volcanoes also yield Early Jurassic ages, and these may have been deposited contemporaneously with the shallower water, possibly inter-tidal sandstones. Alternatively they may represent drifted logs within the turbiditic sequence. The likely environment of deposition for the Late Triassic – Early Jurassic sandstones was within a fluviodeltaic system, shallowing from fairly deepwater during the Triassic to marginal marine in the Early Jurassic (Charlton et al, 1991).


The matrix of the mud volcanic ejecta is primarily Jurassic shale. These are medium-dark grey in colour, partly pyritic, containing abundant ferromanganiferous nodules. Barite mineralisation also occurs commonly in association with the ironstones. The mud volcanic ejecta also locally includes numerous ammonites and belemnites which in general occur loose, but are sometimes enclosed within shale blocks, demonstrating their primary relationship with the shales. Another specimen was identified as late Toarcian based on its ribbing pattern. Thus the entire Lower Jurassic is represented by this faunal collection. Other poorly preserved ammonites and a nautiloid (Cenoceras) are also typical of the Lower Jurassic. Palynological determinations (P. T. Corelab Indonesia, written communication 1987) indicate a Pliensbachian – Callovian age range for shale samples. Part of the jawbone of a Jurassic ichthyosaur was also collected. The environment of deposition for the Jurassic shales is interpreted as a low-energy restricted marine environment, at least partly under anoxic conditions (as indicated by the abundance of pyrite). The palynological determinations indicate a shallow marine, inner shelf environment (P. T. Corelab Indonesia, written communication 1987). The shales were deposited contemporaneously with the upper part of the Triassic – Jurassic sandstone sequence, and the two lithofacies may have interdigitated (or even been parts of the same inter- bedded sequence). The Jurassic shales are again strongly reminiscent of time equivalents in Timor (the Wai Luli Formation), and are also similar to the Buya Formation of the Sula islands (Garrard et al. 1988) and the Yefbie Shale of Misool (Pigram et al. 1982). UNGAR FORMATION

This is a new stratigraphic division proposed here to cover part of what was recorded as the Molu Complex by Sukardi and Sutrisno (1981, 1990). The formation comprises a thick sequence of coarse to medium, massive to well bedded sandstones virtually uninterrupted by finer sediments. It is extensively exposed on Ungar, Vulmali and Natraal islands (Fig. 2), and similar litho- logies are seen commonly as clasts in mud volcanoes in the Yamdena Straits region. Two distinct members of the Ungar Formation have so far been recognised, First, orange or yellow weathering massive to poorly bedded, very coarse (up to 3 – 4 mm) mature quartz sandstones, composed of well rounded and moderately to well sorted quartz grains, with a minor clay matrix sometimes present. These sandstones occur on Vulmali island, and in the ejecta of some mud volcanoes. Second there are greenish or buff fine – medium arkosic sandstones composed of angular to rounded grains, with a glauconite and clay matrix. Parallel, wavy and low-angle cross-bedding on a mm – dm scale are common, but these are the only sedimentary structures so far observed. The glauconitic sandstones have been observed on Ungar and Natraal islands, and in the ejecta of some mud volcanoes. Both sandstone lithologies are variably cemented with calcite, and porosity in the sandstones varies from poor to good depending on the degree of cementation. The environment of deposition for the Ungar For- mation remains uncertain. The formation is at least in part a marine sediment as it contains glauconite, but the large, well-rounded quartz grains may have been derived from an aeolian environment. Sedimentary structures are scarce, apart from parallel- and minor cross-bedding in the glauconitic sandstones. The well-rounded quartz grains may have been reworked from an aeolian environment into a marine one, possibly suggesting a fluviodeltaic system. Neither macro- nor microfossils have so far been found in this sequence, and so the Ungar Formation is not accurately dated. The Ungar Formation is the source for fine-grained sandstones in the Early Miocene Tangustabun Formation (see below), and so pre-dates the Neogene. As sedimentary sequences have already been described above covering much of the Triassic and Early – Middle Jurassic, the Ungar Formation was probably deposited within the age range Late Jurassic – Palaeogene. The Late Cretaceous and Palaeogene of Tanimbar were probably developed in deepwater carbonate facies as will be discussed in the next section. The Ungar Formation is therefore most likely Late Jurassic – Early Cretaceous in age. The dating of the formation will be discussed in more detail in a later section. The thickness of the Ungar Formation is presently unknown, but the extensive outcrops on Ungar Island suggest that the formation is several hundred meters thick. FORMATION

The Tangustabun Formation consists of reddish-brown and grey clay alternating with dirty white vitric tuff, reddish- brown to grey limestone and calcareous sandstone, with quartz sandstones present in the upper part of the sequence. Planktonic foraminifera in this formation indicating a Middle Palaeocene (P4) age for part of the sequence, and concluded a Palaeogene age overall for the formation (Charlton et al, 1991). The geological map shows the Tangustabun Formation cropping out in a number of inliers in the centre of Yamdena, surrounded by the younger Batimafudi Formation. During the 1986 field season the Tangustabun Formation was sampled in the Keustenan river section (De Smet et al. 1990a). In this section the formation consists of greyish quartz sandstone with intercalations of red clay. The sandstone is fine grained and very well sorted, virtually unconsolidated, and lacks obvious sedimentary structures. Bedding is only indicated by occasional clay interbeds which occur irregularly at intervals of metres to tens of metres. Contacts between sandstone and claystone horizons are sharp and flat, with no evidence of bioturbation or scouring. The sandstones contain a reworked fauna of Late Cretaceous (Cenomanian – Maastrichian) and Palaeogene planktonic foraminifera. The youngest faunal element in the samples was dated as planktonic foraminiferal zone N8 (latest Early Miocene), which indicates a maximum age for part of the sequence. The Tangustabun Formation is thus late Early Miocene in age rather than Palaeogene as has been previously suggested. Although no basal stratigraphic contacts have yet been recognized, it is likely that the Tangustabun Formation rests unconformably on the Ungar Formation, which was the primary source of siliciclastic grains in the younger formation. No provenance source has yet been found in Tanimbar for the Late Cretaceous – Palaeogene microfossils which occur reworked in abundance within the Tangustabun For- mation. It is suggested that rocks of this age, which were originally developed in deepwater carbonate facies, were eroded away in the latest Palaeogene or Early Miocene, immediately before deposition of the Tangustabun Formation. This will be discussed in more detail in a later section. The Tangustabun Formation has a thickness of at least 300m in the Keustenan river section (De Smet et al. 1990a), and Sukardi and Sutrisno (1990) estimated a total thickness of 600 m for the formation. De Smet et al. (1990a) were unsure as to the environment of deposition, but overall favoured a relatively deep outer shelf setting.


Several samples collected from mud volcanoes in the Yamdena Straits region are shallow marine sediments of Oligocene or Early Miocene age. These include quartz sandstones composed of well sorted and well rounded grains, and containing large benthic foraminifera. The sand grains are similar to those in sandstones of the Ungar Formation, and are almost certainly derived from erosion of that unit. One of these samples has been interpreted as a near-reef deposit based on its benthic foraminiferal assemblage (A. Racey, pers. commun. 1987). Another sample contains Oligo-Miocene shallow water benthic foraminifera mixed with deeper water planktonic foraminifera, and is probably turbiditic, but again indicates that shallow water environments existed in what is now western Tanimbar during Oligo-Miocene times. A second rock type collected from a mud volcano in the Yamdena Straits region that may be approximately contemporaneous is an immature deltaic coal, dated as possibly Miocene in age (P. T. Corelab Indonesia, written communication 1987). BATIMAFUDI FORMATION

The Batimafudi Formation was defined by Sukardi and Sutrisno (1981, 1990) as a Miocene sedimentary sequence composed predominantly of light grey or brown calcarenites interbedded with white or grey marls. These authors indicated an Early – Middle Miocene age range. On their geological map, the formation is shown outcropping widely in eastern Yamdena. Sukardi and Sutrisno (1981, 1990) also recognised a separate Marl Member of the formation, which they indicated as outcropping in central Yamdena and the Western Islands; that is, to the west of the Batimafudi Formation proper. De Smet et al. (1990a) described the Batimafudi Formation from a number of localities in eastern Yamdena. In that area the formation is composed of interbedded marls and calcarenites, with the calcarenites mainly of turbiditic and mass flow origin. The calcarenites are composed predominantly of planktonic foraminifera, but locally contain well-sorted benthic faunas derived from a carbonate shelf environment. Autochthonous faunas from the intervening marls indicate N16 – N17 (Late Miocene) ages, whilst allochthonous faunas range from Palaeogene to Late Miocene. De Smet et al. (1990a) interpreted a palaeo-bathymetry of 1000 – 2000m for the Batimafudi Formation. Somewhat restricted, probably near-land conditions are indicated by the small size of planktonic foraminifera and the common occurrence of biserial forms in samples collected from the upper Keustenan river section (L. J. van Marle, Free University Amsterdam, personal communication to M.E.M.S. 1986). The formation has an estimated thickness of 700 – 1000 m. The relationship between the Batimafudi Formation and the underlying Tangustabun Formation is not clear (De Smet et al. 1990a). On the one hand a transitional contact is suggested by the occurrence of reddish clays similar to those in the Tangustabun Formation within the lower parts of the Batimafudi Formation; on the other, no rocks were found of Middle Miocene age. At present this problem remains unresolved. The westward transition from predominant cal- carenites in the east to predominant marls in the west as suggested by Sukardi and Sutrisno (1981, 1990) was not supported by our fieldwork. Thick marly intervals occur locally near the east coast of Yamdena (e.g. in the Batimafudi river), and conversely thick calcarenites occur in the Western Islands (e.g. in SE Laibobar). It seems more likely that the transition is a vertical one, from predominantly marly sequences in the lower part of the formation to more calcarenitic at the top. The east – west transition suggested by Sukardi and Sutrisno (1981, 1990) more likely reflects a deeper level of erosion in the west compared with the east. The relative distribution of marl and calcarenite in the Batimafudi Formation does not therefore reflect a westward transition to more distal facies as might be expected on the outer edge of the Australian continental margin. More- over, six well-constrained palaeocurrent indicators were recorded from the Batimafudi Formation with a mean flow towards 120’ azimuth (range of recordings 065 – 160 ). This predominant flow to the SE is the opposite of what would be expected if Tanimbar occupied a position near the edge of the Australian continental block at this time. BATILEMBUTI FORMATION

The Batilembuti Formation was described by Sukardi and Sutrisno (1981, 1990) as comprising dirty white to light grey marl rich in planktonic and benthic fossils. A Plio-Pleistocene age was assigned to the formation, and the environment of deposition interpreted as open marine. At the type section on Cape Batilembuti near Saumlaki, the formation passes upward conformably into Quaternary reef described as the Saumlaki Formation. A possibly lower part of the formation was described by De Smet et al. (1990a) from þ near the village of Batuputih on the west coast of Yamdena. The Batuputih section consists of white, poorly consolidated foram-rich marls, poorly bedded on a scale of 1 – 3 m. Bioturbation is locally observed, and some levels are rich in shell fragments. The planktonic microfauna indicates an Early Pleistocene (N22) age, whilst the benthic fauna indicates water depths of 100 – 300 m. This is in marked contrast with the underlying Batimafudi Formation, which accumulated in water depths of 1000 – 2000m. Also unlike the Batimafudi Formation, the Batilembuti Formation shows no evidence of gravitational transport or faunal reworking. The most striking difference, however, between the Batimafudi and Batilembuti formations is the different degrees of deformation in the two sequences. The Bati- mafudi Formation is rather intensely deformed by folding and thrusting, whilst the Batilembuti Formation is only tilted and locally cut by normal faults. Although the contact between the two formations has not yet been observed, the different degrees of deformation and the age gap between the two (the Pliocene was not identified in the palaeontological studies of De Smet et al. 1990a) suggest an unconformable relationship. It appears that the main phase of deformation on Tanimbar occurred during the Pliocene, during which time the Batimafudi Formation and older rocks were transferred from the outer continental margin of Australia into the arc – continent collision complex, and were simultaneously raised from water depths of 1000 – 2000 m during the Miocene to depths of 100 – 300 m during the Early Pleistocene. The Batilembuti Formation is essentially a post-orogenic deposit, and may infill hollows in the paIaeotopography of the collision complex. It may have a stratigraphic thickness of up to several hundred metres over such topographic lows. SAUMLAKI FORMATION

Quaternary reefs, described by Sukardi and Sutrisno (1981, 1990) as the Saumlaki Formation, are widely distributed around the coast of the Tanimbar islands, and are locally found inland as raised reef terraces. The highest reef terraces recorded by Sukardi and Sutrisno (1981) are on the islands of Molu (200m), Wuliaru (188 m), Teneman (152 m) and Selu (148 m). Quaternary reefs are recorded at 123 m elevation in northeastern Yamdena, 127 m near Saumlaki in southern Yamdena, and 104 m in western Selaru. Elsewhere Quaternary reefs do not exceed 50 m above sea-level. The abundance of raised Quaternary reefs indicates that uplift is continuing in Tanimbar. However, the uplift has not been as dramatic in Tanimbar as in Timor, where Quaternary reefs have been locally uplifted to more than 1200m elevation (Rosidi et al. 1981).


At several localities along the Keustenan river and its tributaries, a marked unconformity was observed between deformed Miocene rocks of the Batimafudi Formation and an undeformed grey clay rich in mol- luscs. The clay is assumed to be Pleistocene in age based on its young appearance and the unrecrystallised nature of the gastropod shells. This claystone is thus the time equivalent of the Batilembuti and/or Saumlaki formations. The unconformity surface is planar and locally truncates folding in the Batimafudi formation. The unconformity surface was repeatedly seen in the banks of the Keustenan river over a distance of several kilometres, during which the river dropped several tens of metres elevation. This suggests that the river cut rapidly through the soft Pleistocene clays until it reached the hard underlying Batimafudi Formation below, after which downcutting of the river slowed considerably, and the river tended to follow the dip of the unconformity surface. Yamdena is topographically asymmetrical, with a sharp escarpment near the east coast, and a long, gentle backslope down to the west coast. It is likely that the backslope is controlled by the Pleistocene basal unconformity which was initially a peneplain surface, and that this surface has been subsequently tilted gently to the NW.


Non-stratigraphic mapping units (metamorphics and ”tectonites”) and some other problematic rock types were grouped together by Sukardi and Sutrisno (1981, 1990) as the Molu Complex. As has already been described, one part of the Molu Complex, the Ungar Formation, is here regarded as a normal mappable stratigraphic sequence. Preliminary investigations of other parts of the Molu Complex (and parts of what was mapped as the Batilembuti Formation) suggest that these areas can be interpreted either as normal strati- graphic sequences, or as part of a unit here described as the Bubuan Mud Complex. We suggest that the name Molu Complex has little geological value, and should be abandoned. Our fieldwork in 1986 and 1987 was insufficient to delineate in detail the normal stratigraphic parts of the ”Molu Complex” other than the Ungar Formation, and these units will not be described further here. Only that part of the ”Molu Complex” (and the Batilembuti Formation) here re-assigned to the Bubuan Mud Complex will be described in the present paper. BUBUAN MUD COMPLEX

The Bubuan Mud Complex is a melange unit, consist- ing typically of decimetre sized blocks of various litholo- gies set in a non-scaly clay matrix. Upon erosion, the clay matrix is often differentially removed, leaving a lag deposit of mixed boulders. The Bubuan Mud Complex outcrops in areas shown on the geological map of Sukardi and Sutrisno (1990) as both the Molu Complex and the Batilembuti Formation (Fig. 3). In general, the designation of the unit as Molu Complex corresponds to areas where the clay matrix has been removed, leaving only the bouldery lag. These areas occur mainly along river courses where the rivers have removed much of the clay matrix. Where the clay and mud matrix remains predominant, rocks that are here assigned to the Bubuan Mud Complex were mapped by Sukardi and Sutrisno (1990) as part of the Batilembuti Formation. This includes extensive areas of western Yamdena (Fig. 3). The name proposed for this melange is taken from the island of Bubuan oA’ the west coast of Yamdena (Fig. 2), on which representative examples of the Bubuan Mud Complex are exposed. As the map of Sukardi and Sutrisno (1990) indicates, Bubuan is also the site of active mud volcanism. The significance of mud volcanism as a rock-forming process in eastern Indonesia was discussed by several pre-war Dutch geologists (e.g. Brouwer 1922, Heim 1942, van Bemmelen 1949). More recently Williams and Amiruddin (1983), Williams et al. (1984) and Barber et al. (1986) have revived interest in mud volcanism in this region by interpreting it as a major element in the formation of melange. In Timor, where levels of erosion are considerably deeper than in Tanimbar, the process of shale diapirism which feeds the mud volcanic activity has been used by Barber et al. (1986) to explain the origin of the Bobonaro Scaly Clay melange. In Tanimbar, the Bubuan Mud Complex is here interpreted as the superficial coalescence of the ejected material from numerous mud volcanoes. The Bubuan Mud Complex might also be comparable with the Salas Block Clay of Seram (Audley-Charles et al. 1979).

The matrix of the Bubuan Mud Complex is fairly uniform throughout Tanimbar. It consists of a medium to dark grey clay, thixotropic and sticky when wet, but generally completely desiccated at outcrop during the dry season. As described above, the clays locally contain an abundant ammonite fauna that indicates the clays are primarily Early Jurassic in age. A variety of rock types are found as boulders in the mud complex, and relative proportions vary considerably from locality to locality. The most abundant com- ponent is ferro-manganiferous, sometimes septarian, nodules that are thought to have originated within the Jurassic claystones. Barite and calcite mineralisation are also cogenetic with the ironstones. The second commonest component is sandstone: both Triassic sand- stone and Late Jurassic – Early Cretaceous(?) sandstone of the Ungar Formation. Triassic sandstones are par- ticularly abundant from the Yamdena Straits region, but are absent from the Bubuan Mud Complex in the south of the island near Saumlaki. Other subordinate rock types include pastel pink unfossiliferous calcilutites (very similar to the latest Cretaceous – earliest Palaeogene Borolalo Limestone of Timor), Oligo-Miocene shelf sediments, serpentinite and high grade metabasic amphibolite.

11.4. KAI BESAR & KAI KECIL edit

Geological map of Kai Besar and Kai Kecil Islands

The stratigraphy of Kai Besar (Big Kai) has been described by Achdan and Turkandi (1982) with additional data from Charlton et al. (1991) which are used in this chapter. The oldest stratigraphic unit consists of flat-bedded poorly fossiliferous calcilutites and marls interbedded on a decimeter scale, named Elat Formation. Achdan and Turkandi (1982) reported Upper Eocene ages from planktonic foraminifera in the marls, with reworked Middle-Upper Eocene benthic foraminifera in the calcilutites. They estimated the formation to be about 500 m thick. In 1987, Charlton et al. logged approximately 450 m of this unit from continuous coastal exposure north and south of Mun, and they suspect that the total exposed thickness may be some what greater, perhaps 600-800 m. The Elat fm. is interpreted as pelagic or hemipelagic carbonates deposited in a distal continental slope setting, possibly slightly shallowing upward. The Elat fm. is overlain (disconformably?) by yellowish or reddish-brown shallow water limestone of Tamangil Formation. The limestone is a characteristic calcirudite containing numerous Lepidocyclina benthic foraminifera up to 6 cm diameter. Achdan and Turkandi (1982) reported a Middle-Upper Oligocene age for this unit, and identified a stratigraphic thickness of up to 50 m. In our opinion the unit is sufficiently distinct to warrant separate formational status. The overlying Weduar fm. consists of reef limestone, calcilutite, calcarenite and marl deposited in neritic shelf environments. It is believed to be entirely Miocene in age. According to Achdan and Turkandi (1982), the formation is approximately 500 m thick. The youngest stratigraphic unit recognised in Kai Besar by Achdan and Turkandi (1982) is the Weryahan Formation, which consists of shallow water limestone and marl of Pliocene age. On their map, Achdan and Turkandi (1982) indicated stratigraphic contacts between the Weryahan Formation and both the Elat Formation (Eocene) and Weduar Formation (Miocene). This suggests that the Weryahan formation has an unconformable relationship with these older sequences. Charlton et al (1991) were unable to confirm these relationships after visiting the locality, as the area indicated consisted of unbroken outcrops of the Elat Formation. At the second indicated locality of the Weryahan Fm. immediately north of Weduar village, a simple gradation was observed in a series of outcrops from the Weryahan Formation down into the Weduar Formation (Charlton et al., 1991). Bedding dips in the Weryahan and Weduar formation are similar, indicating that the boundary is at most oa disconformity.

There are limited publication related to the geology of Kai Kecil (Small Kai). Pertamina-BEICIP map (1982) shows that Kai Kecil island complex is dominated by Quarternary reefs. Pliocene Weryahan Fm occur in isolated places as olistostrome, which probably generated by mud volcanoes. The Kai Kecil is separated from Kai Besar by a significant thrust fault dipping to the west.

11.5. SERAM (After Kemp & Mogg, 1992) edit

Geological map of Seram Island

The stratigraphy of Seram is strongly controlled by the Island structural development. As a result, an integrated analysis and understanding of the two must be undertaken. Seram Island itself sits within a complex region of crustal plate interactions; the Australian, Eurasian and Pacific-Philippine Plates have all had a major effect on the development of the Miocene to Recent sedimentation (Seram Series) and have overprinted and obscured stratigraphic relationships within the pre-Late Miocene section (Australian Series). The stratigraphic and structural model set out below is an update of that first presented by Kemp and Mogg (1992)


The Australian Series (Kemp and Mogg, 1992) consists of Permian through Late Miocene units that form the bulk of the sequence deposited at Seram. These units, deposited on and along the northern margin of the Australian Continental Plate, constitute an integral part of the Gondwana (Lower Triassic and older) and Westralian (Mid-Triassic to Latest Miocene) Superbasins as defined by Bradshaw et. al. (1988, 1994). During the Late Cabonisferous through Early Permian, major intra-cratonic extension and rifting occurred throughout the region of the north and northwest, Australian margin. This event resulted in a series of north to northeast trending rift basins bounded by normal faults that define the margins of individual basin elements. To the south of Proto-Seram, these formed the margins of the Vulcan - Malita - Calder Graben system. At Proto-Seram in the north, the extensional basin system formed the depocentres into which the sediments of the Kobipoto-Taunusa and Tehoru Complexes were deposited. Subsequent to deposition, a major metamorphic event occurred in the region of Proto-Seram resulting in high to low grade metamorphism of these Permian and older units. This early stage metamorphism is confirmed by the observed inclusion of metamorphic clasts of these unit within younger sediments (ie. Kanikeh Formation). This initial heating phase may be related to a localised extremes in crustal thinning and an associated heat pulse although this is still speculafive. DeSmet and Barber (1992) and Linthout et. al. (1991) suggest an additional phase of late stage Neogene metamorphism resulting from ophiolite obduction along the present-day south side of the Island (palaeo-west). This late second stage event is consistent with the main thrusting episode of the structural model presented.

From the Late Permian through the Early Jurassic, the Westralian Superbasin underwent a period of relative tectonic quiescence and regional thermal subsidence. This period of intracratonic sag, resulted in depocentre axes coincident with the previous rift basins, filled with a series of fluvio-deltaics in the south / southeast to marine units in the area of proto Seram. The Kanikeh Formation represents this period of deposition on Seram and consists of a series of fine to coarse grained, at times conglomeratic, litharenites and feldspathic litharenites that show distinct graded bedding and are interbedded with siltstones and mudstones. Lithic fragments consists of volcanic, igneous, metamorphic and sedimentary fragments. Metamorphic material derived from the Kobipoto and Tehoru Complexes is seen as intra-clasts. Calcarenites, calcilutites, shaly limestones, limestones and calcareous sandstones are found as intercalations, interbeds and discrete units within the Kanikeh Formation. The clastic Kenikeh units on Seram are mainly turbidite / gravity flow deposits with well defined Bouma sequences identified in outcrop in central and eastern Seram. Rich carbonaceous beds of detrital coals are seen throughout. Shallower water units have also been identified which may be associated with storm sand deposition. The shales and mudstones within the clastic Kanikeh Formation are believed to act as important decollement layers during the Late Miocene and younger thrusting episode. Some of the coarser and conglomeratic carbonate units also have gravity / mass flow origins.

Palynological data has shown that the Kanikeh Formation is Middle to Late Triassic in age, ranging from Ladinian to Norian (P.T. Geoservices, 1991) with the palynological assemblages being identical to those seen along the Northwest Shelf of Australia (Dirk Hos, pers comm., Price, 1976).

From the late Triassic through the Early Jurassic, the Westralian Superbasin underwent a renewed period of compression, fault reactivation and uplift. In the region of Proto- Seram, a developing high was isolated from the main clastic depocentres adjacent to the mainland by an intermediary deep basinal region. This resulted in isolation of the area from clastic input and the deposition of deepwater carbonates of the Saman-Saman Formation. These deeper water facies grade laterally and upward into the shallower facies Manusela Formation. Weber (1926), however, felt that the Saman – Saman was apparently conformable with, and possibly gradational to, the Kanikeh Formation. This may be based on lithostratigraphic correlations to similar carbonate facies within the Kanikeh section although this is speculative. The Saman -Saman Limestone consists of marls with interbedded calcilutites and nodular chert inter-beds (Tjokrosapoetro & Budhitrisna, 1982; O’Sullivan et. al. 1985). The Saman- Saman Limestone is interpreted to have been deposited in a moderate to deep water, outer- shelf to bathyal setting. Outcrops of the Saman- Saman Limestone are mapped in the central highlands region of Seram. Weber (1926) described similar limestones with strong bituminous staining in the central mountains.

The Saman-Saman Limestone is overlain by and is, in part, laterally equivalent to, the shallow water limestones of the Manusela Formation. This unit consists of skeletal oolitic grainstones and is found in outcrop in the Nief Gorge and in the exploration wells East Nief – 1 and Oseil - 1. Outcrops of the Manusela Formation have also been mapped in central Seram and in the Watubela Islands (Weber, 1926, Tjokrosapoetro & Budhitrisna, 1982). The Manusela Formation has extensive sand- size oolite grains dominating the groundmass with lesser bioclastic skeletal material sitting in a fine grain, partly fecal, matrix. Dolomitisation has occurred in some outcrop samples and is widespread in the section penetrated in East Nief - 1 and Oseil - 1.

Dating foi the Manusela Formation has proved difficult due to a lack of biostratigraphic material typical of such shallow, high energy, clean formations. However, outcrop data and samples from East Nief - 1 and Oseil - 1 have yielded ages ranging from Early Jurassic (Pliensbachian or older) Callovian - Lowermost Oxfordian or Bathonian. The Manusela Formation was probably deposited on a regional outer rise high. This rise was further uplifted as a result of the Late Triassic to Early Jurassic compression and the shallowing upward Saman-Saman / Manusela sequence deposited. Similar carbonate prone highs are recorded along the North West Shelf of Australia during the Triassic and Early Jurassic, such as the Ashmore and Exmouth Platforms (Exon et al, 1991; Barber, 1982). From the Late Callovian through Early Oxfordian, the Westralian Superbasin underwent extension and half graben reactivation. This was quickly followed by continental break-up, rapid subsidence, marine transgression (O’Brien, 1993) and development of the widespread Callovian unconformity (Mory, 1988; Bradshaw et al; 1988, Struckmeyer et al; 1991, among others - Table 4). The Manusela Formation was rapidly flooded and carbonate deposition halted. This marine transgression resulted in the deposition of the overlying Kola Shale at Proto-Seram. The unit consists of grey and red-brown claystone and shale that was deposited in a shelfal (possibly neritic) to outer shelf environment. The Kola Shale in East Nief - 1 has been dated as Berriasian - Kimmeridgian to as old as Middle Oxfordian (Lowermost Cretaceous to Upper Jurassic) and Late Jurassic to Early Cretaceous at Oseil - 1. A possible minor disconformity within the Kola Shale has also been identified from palynological evidence between the Mid-Tithonian and Upper-Kimmeridgian. This coincides with the intra-Kimmeridgian unconformity seen in the Vulcan Graben (Patillo & Nichols, 1990). The Kola Shale may act as an important decollement surface during thrusting.

The shelf setting for the Kola Shale also conforms to the paleogeographic model of Struckmeyer et. al. (1990) for this time section. The Kola can be correlative with the Maril Shale of the Papuan Basin, the outer shelf deposits of the Kopai (Lengguru region) and the Lelinta Shale on Misool (Pigram et. al., 1982). Time equivalent units in the Vulcan Graben in the Timor Sea region are important source units; however, the Kola Shale represents a more distal, open marine deposit and is not a potential source at Seram, a conclusion supported by geochemical studies (Corelab, 1988 & 1994).

Immediately following continental break-up in the Late Jurassic through Early Cretaceous, a brief period of transpression caused strike-slip reactivation along pre-existing normal faults, rotation and uplift that resulted in the development of the Valanginian Unconformity (Table 4). This event marks the top of the Kola Shale at Seram. The Westralian Superbasin then entered into a period typical of passive margins with marginal sag basin development from the Early Cretaceous (Falvey and Mutter, 1981; Patillo and Nicholls, 1990, Struckmeyer, 1990; and O’Brien at. al., 1993). Widespead regional transgression of the continental magin began at this time. The paleogeographic location of Seram on the outer margin of the Australian continent moved rapidly from neritic (Kola Shale) to a clastic starved, outer-shelf, shelf slope, bathyal environment. The Nief Beds were deposited over the unconformity surface. The Nief Beds consist of a condensed sequence of mudstones, calcilutites, marls, cherts, cherty limestones, sandy shales and lesser coraline and reefal limestones.

Seram remained, for the most part, in this distal setting throughout the Cretaceous, Palaeogene and into the Miocene. Although compressive events and transgressive-regressive phases have been clearly identified in the adjoining, more proximal areas along the northwestern Australian margin (Patillo and Nichols, 1990 and O’Brien et. al., 1993), these are less evident on Seram, partly due to the distal location and to complex tectonic overprinting obscuring detailed relationships. There is, however, some evidence for at least two shallowing phases in the geologic record during this marginal sag phase.

The earliest of these is represented by fractured Early Palaeocene coralline limestone’s (Kemp and Mogg, 1992) and may represent a period of shallow water deposition during regional uplift. Widespread uplift in the Late Cretaceous and Palaeocene has been identified elsewhere along the northern Australian margin associated with the opening of the Coral Sea to the east (Struckmeyer et. al., 1990; Patillo and Nichols, 1990; Etheridge et. al., 1991). A second period of shallow water coralline limestone deposition has been identified from outcrop samples in Central Seram. This episode, dated as late Miocene, probably represents the progressive uplift associated with the next major period of tectonism on the island; that of the Australia-Eurasian-Pacific plate collision that began as early as the Oligocene (Etheridge et. al, 1991). Evidence of clastics influx into the sequence is seen in Bolifar Utara - 1 where late Miocene clastics have been identified. These are probably the result of erosion and re- working of the Mesozoic to Late Miocene sequence that was being uplifted in response to early compression and associated thrusting.


The Late Miocene marks a critical phase in the geology and tectonic evolution of Seram. It was at this time that the collision between the northward moving Australian, eastward moving Eurasian and westward moving Pacific- Philippine plates had its major influence and accelarated thrusting and uplift of the section occurred at Seram. As the area of Proto-Seram moved northward as a part of the Australian Continental plate, it eventually entered a mobile belt that is presently bounded to the north by the Sorong Fault System and to the south by the Tarera-Aiduna Fault System (Figure 4). This mobile belt is a complex left-lateral strike-slip zone caused by the oblique convergence of the Philippine-Pacific plate against the northern margin of the Australian Plate. Uplift and erosion of the thrust belt since the Early Pliocene has provided the source material for the Seram Basin sediments of the Salas Complex, Wahai and Fufa Formations. Associated with the initial stages of thrusting and rapid orogenic uplift, a gravity slide / slump unit, the Salas Complex, was deposited in outer shelf to bathyal water depths and sits unconformably on the sediments of the Australian Series. The Salas consists of clays and mudstones and contains clasts, boulders and erratic block of the pre-thrusting sequence. The Salas represents a rapidly deposited gravity slide / slump unit that developed over the uplifting and eroding thrusted sediments as formation of the thrust belt progressed (Figure 8). The mechanism for developing these types of deposits are described by Jones (1987). Over pressuring within the Salas has resulted due to this rapid deposition and is seen in Bolifar Utara l. In both East Nief – 1 and Bolifar Utara - 1, the Salas unconformably overlies the Nief Beds. Direct age dating of the Salas is difficult due to the reworking of older biostratigraphic source material into the Salas. Dating of the underlying Nief and overlying Wahai, although unconformable, restrict its depositional age to the Early Pliocene although slightly older section may exist . As erosion of the uplifted thrust belt progressed, a reduction in the frontal slope resulted in a change from gravity slide and slump dominated sedimentation (the Salas Complex) to normal clastic deep water outer shelf to bathyal sedimentation. As further uplift of the thrust front progressed, a series of narrow thrust foreland basins developed parallel to the strike direction of the thrust front. These formed elongate ”perched” young basins that overly the older Mesozoic sequence. Of those identified, the best preserved are the Bula and Wahai basins which extend along the northern coast of Seram Island from east to west respectively. It was into these developing basins that the Wahai Formation (Figure 8) was first deposited. The Wahai has been dated as Early Pliocene to Early Pleistocene (Zillman and Paten, 1975) and consists of mudstones, siltstone and deep water limestones deposited in dominantly bathyal locations. During the Early Pleistocene, continued uplift of the Island resulted in progressively shallower water depths in the thrust foreland basins and deposition of the neritic mudstones, claystones, sands, silts, conglomerates and limestones of the Fufa Formation occurred (Figure 8). Both the Fufa and Wahai Formations have also been described by Zillman and Patten (1975), Tjokrosapoetro et. al. (1988) and others for north east Seram and by DeSmet et. al. (1989) in the south-west.

11.6. BURU edit

In Buru Island, based on regional correlation, the pre-Triassic interval is grouped as the Wahlua basement complex, comprises low-grade metamorphic rock. After transgression in Triassic, Dalan Formation was deposited, containing reef-slope carbonates and grading into a neritic to outer shelf sequence of flysch-type sediments. Folding and uplifting took place after the sedimentation of the Dalan Formation, and Jurassic is represented by a hiatus in the geological record and volcanic activity might have taken place during this time. The Ghegan Formation, is the oldest member of the Buru Group, consists of dolomitized limestone with minor fine clastic sediments, very rich in carbonaceous material. The carbonates represent reef-slope wackestones and minor packestones deposited in a neritic to outer shelf environment. The Kuma Formation, overlies the carbonates pass gradually, or interfingers with Ghegan Formation. Thinly bedded pelagic calcilutites, with chert layer alternations are the most common lithology of Kuma Formation. Andesitic tuff and lava flow is encountered in the uppermost part of the Kuma Formation,. This formation is dated as Middle Cretaceous to Eocene age. The Kuma Formation is unconformably ovelain by the Waekan Formation, which is dominated by coarse to fine sandstones and carbonates lithologies. Interbedded with these clastic sediments are volcaniclastic and lava flows with andesitic composition. The age of this formation is mostly Oligocene, but might range from Eocene to Oligocene. The Paleogene and older sediments are overlain with a distinct unconformity by Early to Middle Miocene Hotong Formation, consists of sandstones and conglomeratic sandstone, with minor amounts of shale, marl, clay and limestone. The Pliocene Leko Formation mainly consists of conglomerates and conglomeratic sandstone, which in the uppermost part interfinger with limestone. The Leko formation is overlain by Quaternary reefs, terrain deposits and alluvium.

External Links edit

Darman, H. & Reemst, P., 2012, Seismic Expression of Geological Features in Seram Sea: Seram Trough, Missol-Onin Ridge and Sedimentary Basins, Berita Sedimentologi #23