Tectonostratigraphic evolution of Cenozoic marginal basin and continental margin successions in the Bone Mountains, Southwest Sulawesi, Indonesia

Tectonostratigraphic evolution of Cenozoic marginal basin and continental margin successions in the Bone Mountains, Southwest Sulawesi, Indonesia

Journal of Asian Earth Sciences 38 (2010) 233–254 Contents lists available at ScienceDirect Journal of Asian Earth Sciences journal homepage: www.el...

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Journal of Asian Earth Sciences 38 (2010) 233–254

Contents lists available at ScienceDirect

Journal of Asian Earth Sciences journal homepage: www.elsevier.com/locate/jseaes

Tectonostratigraphic evolution of Cenozoic marginal basin and continental margin successions in the Bone Mountains, Southwest Sulawesi, Indonesia Theo M. van Leeuwen a,*, Eko S. Susanto b, Sigit Maryanto b, Sapri Hadiwisastra c, Sudijono b, Muhardjo d, Prihardjo b a

Jl. H. Naim IIIB No. 8, Jakarta 12150, Indonesia Geological Development and Research Centre, Jl. P. Diponegoro No. 57, Bandung, Indonesia Lembaga Geoteknologi, LIPI, Jl. Cisitu, Sangkuriang No. 21, Bandung, Indonesia d PT Rio Tinto Exploration, 15th Floor Menara Anugrah, Kantor Taman E3.3., Jl. Mega Kuningan Lot 8.6–8.7, Jakarta 12950, Indonesia b c

a r t i c l e

i n f o

Article history: Received 19 August 2008 Received in revised form 15 November 2009 Accepted 28 November 2009

Keywords: Southwest Sulawesi Continental margin Marginal basin Stratigraphy Geochemistry Paleontological dating

a b s t r a c t The Bone Mountains, located in Southwest Sulawesi along the SE margin of Sundaland, are composed of Oligocene to possibly lower Miocene marginal basin successions (Bone Group) that are juxtaposed against continental margin assemblages of Eocene–Miocene age (Salokalupang Group). Three distinct units make up the latter: (i) Middle–Upper Eocene volcaniclastic sediments with volcanic and limestone intercalations in the upper part (Matajang Formation), reflecting a period of arc volcanism and carbonate development along the Sundaland margin; (ii) a well-bedded series of Oligocene calc-arenites (Karopa Formation), deposited in a passive margin environment following cessation of volcanic activity, and (iii) a series of Lower–Middle Miocene sedimentary rocks, in part turbiditic, which interfinger in the upper part with volcaniclastic and volcanic rocks of potassic affinity (Baco Formation), formed in an extensional regime without subduction. The Bone Group consists of MORB-like volcanics, showing weak to moderate subduction signatures (Kalamiseng Formation), and a series of interbedded hemipelagic mudstones and volcanics (Deko Formation). The Deko volcanics are in part subduction-related and in part formed from melting of a basaltic precursor in the overriding crust. We postulate that the Bone Group rocks formed in a transtensional marginal basin bordered by a transform passive margin to the west (Sundaland) and by a newly initiated westerly-dipping subduction zone on its eastern side. Around 14–13 Ma an extensional tectonic event began in SW Sulawesi, characterized by widespread block-faulting and the onset of potassic volcanism. It reached its peak about 1 Ma year later with the juxtaposition of the Bone Group against the Salokalupang Group along a major strike-slip fault (Walanae Fault Zone). The latter group was sliced up in variously-sized fragments, tilted and locally folded. Potassic volcanism continued up to the end of the Pliocene, and locally into the Quaternary. Ó 2010 Elsevier Ltd. All rights reserved.

1. Introduction The geology of the Bone Mountains in Southwest Sulawesi (Fig. 1A), located along the southeastern margin of Sundaland, is characterized by two contrasting domains. Along the western foothills of this mountain range a series of marine sedimentary and volcanic rocks is exposed in a narrow, 31 km long, faultbounded strip. These were deposited in a continental margin setting between the Middle Eocene and Middle Miocene. They can be divided into three distinct lithostratigraphic units, which we have named Matajang, Karopa and Baco Formations, and grouped together into the Salokalupang Group. The remaining part of the * Corresponding author. Tel./fax: +62 21 7250 326. E-mail address: [email protected] (T.M. van Leeuwen). 1367-9120/$ - see front matter Ó 2010 Elsevier Ltd. All rights reserved. doi:10.1016/j.jseaes.2009.11.005

Bone region is largely underlain by a sequence of volcanic rocks of marginal basin affinity, the Kalamiseng Formation of Sukamto (1982). Sandwiched between the Kalamiseng Formation and the Salokalupang Group is a sequence of mudstones and intercalated volcanics, referred to in this paper as Deko Formation, which we interpret to have also been formed in a marginal basin environment. We group the two formations of marginal basin affinity together into the Bone Group. The Bone Mountains consist of a chain of mountains, roughly aligned in a N–S direction, which are up to 850 m high. A second mountain chain, the Western Divide Mountains, occupies the western part of SW Sulawesi. The two mountain belts are separated by a graben-like structure that is known as the Walanae or Central Depression. This structure forms parts of a major N–NW trending ‘‘hinge-line”, named the Walanae Fault Zone (van Leeuwen, 1981).


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Previous investigations of the Bone Mountains are limited. t’Hoen and Ziegler (1917) were the first to make geological observations in this area as part of a reconnaissance mapping programme of SW Sulawesi during the early part of the 20th century. They reported the occurrence of nummulitic limestone in the western foothill zone and altered volcanics further to the east. More than 60 years later a second geological mapping survey was undertaken in SW Sulawesi, again on a reconnaissance scale, by Sukamto (1982) and Sukamto and Supriatna (1982) of the Geological Survey of Indonesia. They gave the name Salokalupang Formation to the various rock units found along the western margin of the Bone Mountains, including the Deko Formation, and assigned a Middle Eocene–Oligocene age to it, based on limited paleontological evidence. Yuwono (1987) carried out petrographic and geochemical analyses of a number of Kalamiseng volcanic samples collected along a traverse across the Bone Mountains. The results of his study, which also included limited K/Ar dating, are briefly discussed by Yuwono et al. (1988a). In this paper we present the results of investigations that we undertook intermittently between 1995 and 2002. It involved mapping and sampling along 12 river traverses that cut across the western margin of the Bone Mountains (Fig. 1B). Part of the work was done in the framework of a stratigraphic research program carried out by the Geological Research and Development Centre (GRDC) in South Sulawesi between 1995 and 1999. It consisted of measuring detailed stratigraphic sections along four traverses, viz S. Kalupang, S. Deko, S. Katumpang, and S. Baco (Note: S. stands for Salo, which means river in the local Bugis language). The results of this exercise have been presented in an earlier paper (Maryanto et al., 2004). In total we collected more than 900 outcrop and float samples for paleontological studies, geochemical analyses and petrographic examinations. The main objective of our investigations was to elucidate the stratigraphic succession of the Salokalupang Formation of Sukamto (1982), i.e. our Salokalupang Group and Deko Formation. We have been able to establish a broad stratigraphic scheme, despite considerable structural complexities resulting from the location of the rocks within the Walanae Fault Zone. Another aim was to determine the age and origin of the Kalamiseng Formation, and its tectonostratigraphic relationship to the other units. The new field and laboratory data from the Bone Mountains, integrated with published data from adjacent regions, provide the basis for a new model for the evolution of Southwest Sulawesi.

2. Geologic setting The Sulawesi area is tectonically one of the most complex regions of the world, as it is the site of the interaction of three major lithospheric plates – the Eurasian Plate to the west, the Pacific Plate to the east, and the Australian–Indian Plate to the south. Since the Cretaceous it has been subjected to multiple phases of rifting, subduction, magmatism, collision, strike-slip faulting and rotation, but the exact timing of events is still not well understood (Hall and Wilson, 2000). Sulawesi has been divided into several provinces (Sukamto, 1975; Hamilton, 1979) which are, from west to east, the West Sulawesi Plutono-Volcanic Arc, the Central Sulawesi Metamorphic Belt, the East Sulawesi Ophiolite, which is interspersed with smaller units of Mesozoic and Cenozoic sediments, and the microcontinental blocks of Banggai-Sula and Buton-Tukang Besi (Fig. 1C). The West Sulawesi Plutono-Volcanic Arc consists of two distinct segments, namely ‘‘Northern Sulawesi”, a Cenozoic tholeiitic to calc-alkaline island arc system, underlain in part by oceanic crust, and ‘‘Western Sulawesi” which developed in a continental margin setting along the southeastern edge of the Eurasian Plate, known as Sundaland, during the Late Mesozoic–Cenozoic,

and which is characterized by extensive Neogene potassic magmatism (Taylor and van Leeuwen (1980); van Leeuwen and Muhardjo, 2005). The oldest dated events related to accretion and/or suturing are Late Oligocene–Early Miocene (Hall and Wilson, 2000). The East Sulawesi Ophiolite, which in part may represent the basement of the pre-collisional fore-arc complex, was obducted onto imbricated Australian continental crust now seen in the Buton microcontinent and the continental metamorphic complexes of Southeast Sulawesi (Hall and Wilson, 2000). There may have been a significant strike-slip component to convergence before collision of the earliest fragments (Hall, 1996). Convergence continued throughout the Neogene until the present day in the development of several other major thrusts, each active in different part of the region and at different times since the Early Miocene (Hall and Wilson, 2000). SW Sulawesi (Fig. 1A) is separated structurally from the rest of Western Sulawesi by the Tempe Depression which is filled with Quaternary sediments (van Leeuwen, 1981). To the west and east it is bounded by deep marine basins, i.e. the Makassar Straits and Bone Bay. A Mesozoic basement is overlain with angular unconformity by the Paleocene–Eocene Alla and Langi Volcanics and the terrestrial to marginal marine Eocene Malawa Formation (van Leeuwen, 1981; Sukamto, 1982). The dominantly siliciclastic Malawa Formation passes transgressively upwards in shallow marine carbonate successions of the Tonasa Formation, the base of which range in age from Middle Eocene to Late Eocene (Sukamto, 1982; Wilson, 1995; Harahap et al., 1999). In some areas platform carbonates accumulated until the Middle Miocene, in others deeper marine marls and redeposited carbonates occur (Wilson and Bosence, 1996). The carbonate successions are overlain by thick sequences of Mio-Pliocene volcanic and volcaniclastic rocks of the Camba Formation and associated units, which are exclusively potassic in nature (Sukamto, 1982; Yuwono et al., 1988a; Polvé et al., 1997; Elburg and Foden, 1999; Elburg et al., 2002). In eastern SW Sulawesi, i.e. the Bone Mountains and adjacent areas, the oldest rocks exposed are Middle–Upper Eocene volcaniclastic sediments, volcanics and carbonates of the Matajang Formation, which are overlain by Oligocene carbonates of the Karopa Formation and Lower–Middle Miocene sediments, volcaniclastics and volcanics of the Baco Formation, forming together the Salokalupang Group (this issue). These rocks are juxtaposed against Oligocene–Lower Miocene volcanics and mudstones of the Kalamiseng Formation and Deko Formation (Bone Group, this issue). The formations are unconformably overlain by the Middle– Upper Miocene Camba Formation. Coral-rich shallow marine carbonates of the Middle Miocene–Pliocene Tacipi Formation outcrop along the northern margin of the Bone Mountains (Ascaria et al., 1997; Grainge and Davies, 1985; Mayall and Cox, 1988). The mountain range is surrounded by a thick sequence of uppermost Miocene–Pliocene tuffaceous rocks, volcaniclastics, marls and other marine beds, which towards the top becomes more terrestrial (Walanae Formation) (t’Hoen and Ziegler, 1917; Sukamto and Supriatna, 1982; Grainge and Davies, 1985; van den Bergh et al., 1995). The structure of SW Sulawesi is relatively simple. In some areas the Camba Formation overlies tilted or locally folded older formations with angular unconformity, whereas in others there is a conformable contact (van Leeuwen, 1981; Samodra and Purnamaningsih, 1993; Wilson, 2000). The younger strata are generally not significantly deformed, with the exception of the Sengkang Basin in the northeast, where they are folded in a series of northerly trending anticlines and synclines (Grainge and Davies, 1985), and the Bantimala area, which underwent localized thrusting and folding (Sukamto, 1985; Wilson, 1995).

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Fig. 1. (A) Summary of the geology of SW Sulawesi; modified after Sukamto (1982) and Sukamto and Supriatna (1982). (B) Traverse location map. (C) Principal tectonic provinces of Sulawesi.

The dominant structure in SW Sulawesi is the Walanae Fault Zone, which consists of two fault systems that bound the 15 km wide Walanae Depression on either side, i.e. the West and East Walanae fault systems (van Leeuwen, 1981). The latter, which is the most prominent of the two, runs along the western margin of the Bone Mountains with a distinct topographic expression and continues in a northerly direction (Sukamto, 1982) where it divides

the Sengkang Basin into two sub-basins (Grainge and Davies, 1985). In the latter area its present configuration is a high angle reverse fault downthrown to the east (Grainge and Davies, 1985). Further to the north the Walanae Fault Zone has been interpreted to be truncated by a series of NW–SE strike-slip faults found along the northern margin of the Tempe Depression (Berry and Grady, 1987) or to continue into Central-West Sulawesi, where it links


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with the Masupa fault (Guritno et al., 1996), also referred to as Sadang fault (Hamilton, 1979). To the south it continues into the Bone Basin. The Walanae Fault Zone is commonly interpreted as a major sinistral strike-slip fault (e.g. Sukamto, 1975, 1982; van Leeuwen, 1981; Grange and Davies, 1985; Berry and Grady, 1987; Guntoro, 1995), based primarily on regional considerations, such as the straightness and considerable strike length of the fault zone. Second order reverse faults seen on seismic data support Pliocene wrenching (Grange and Davies, 1985) and stream offsets in the Biru area along the West Walanae fault system indicate young lateral movement (van Leeuwen, 1981). However, seismic data (Grange and Davies, 1985), geological relationships and geomorphological features (van Leeuwen, 1981) also suggest significant vertical movement. Further to the north, motion along the Masupa fault initially involved left-lateral movements, which subsequently (Late Pliocene or Quaternary) changed to right-lateral (Guritno et al., 1996). Inception of the Walanae Fault Zone probably took place during the Middle Miocene (van Leeuwen, 1981; Grainge and Davies, 1985; this issue). Another significant NNW-trending strike-slip fault, which was active during the Pliocene, occurs further to the west in the Bantimala-Barru area (Berry and Grady, 1987).

3. Stratigraphy This section gives an overview of the lithological and biostratigraphic data that we collected in the field, supported by laboratory analyses. Field mapping and stratigraphic subdivisions were predominantly based on lithostratigraphic units. Fossils (benthonic and planktonic foraminifera, and nannofossils) provided the time control. Foraminiferal determinations were made from standard petrographic thin sections. 142 samples were prepared for paleontological analysis. Of these 51 samples were barren or contained non-age specific or exclusively reworked assemblages. Foraminifera and nannofossil assemblages recorded from the remaining samples varied in species diversity, abundance and preservation. Reworking of taxa is a significant feature of the Oligocene and Miocene strata, giving the rocks in a number of cases a seemingly older age. The Geological Time Scale used in this paper is that of Gradstein et al. (2004). The planktonic foraminifera and nannofossil assemblages have been interpreted using the zonation schemes of Blow (1969, 1979) as modified by Bolli et al. (1985) and Okada and Burky (1980) respectively. Large foraminifera have been assigned to the letter-stages of the Far East (Adams, 1970). A summary of the paleontological data is presented in Appendix A. Schematic cross sections of selected traverses are presented in Fig. 2. They show the locations of paleontological and geochemical samples mentioned in the text, Appendix A and Table 1. A generalized stratigraphic scheme for the Salokalupang and Bone Groups is shown in Fig. 3.

3.1. Salokalupang Group The Salokalupang Group (Matajang, Karopa and Baco Formations) spans the period late Middle Eocene–Middle Miocene with apparent gaps in the Early Oligocene and Late Oligocene–Early Miocene. The reason(s) for these gaps is/are not clear. They may be an artifact of sampling or may be due to a lack of diagnostic fossils. Alternatively the intervals may represent condensed successions, or (unlikely) a non-deposition/erosion event, or they form limited outcrops because of faulting. The latter is a distinct possibility in view of the strongly tectonized character of the sequence.

3.1.1. Matajang Formation The Matajang Formation, a thick sequence of sedimentary and subordinate volcanic rocks of Middle–Late Eocene age, forms the lower part of the Salokalupang Group. It is exposed in all river traverses except S. Kalupang and S. Baco. The formation can be subdivided into two units; Member A (predominantly volcaniclastic sediments and mudstone) and Member B (volcanics with limestone intercalations). Member B overlies the lower part of Member A and is interpreted to interfinger with the upper part. Member A, which is >450 m thick, is predominantly composed of volcaniclastic sandstones, siltstones and mudstones with subordinate intercalations of lava and limestone in the upper part. It is exposed in S. Deko, S. Deko Kanan, S. Linrung, S. Matajang, S. Katumpang, and S. Mate (Fig. 2). It is mostly in faulted contact with various other units and in the west unconformably overlain by the uppermost Miocene–Pliocene Walanae Formation, that consists of a thick series of sedimentary rocks with a high volcanic component (Sukamto, 1982; van Leeuwen, 1981). In some localities sandstones and mudstones are interbedded in a regular fashion; in others sandstones are dominant, containing interbedded mudstones that vary in volume from <5% to 40%; in yet others thick mudstone intervals (up to 40 m) are present with interbedded muddy breccias and minor intercalations of volcaniclastic sandstone. The sandstone–mudstone sequence contains rare intercalations of breccia, lava and limestone. Details of the main sedimentary rock types are presented in Appendix B. Lava intercalations were observed only in S. Matajang (olivine basalt) and S. Katumpang (reddish aphanitic lava, very similar to the hematite-rich lavas from Member B in S. Mate). Limestone intercalations occur as thin (<3 m) interbeds or somewhat thicker lens-like bodies (or in some cases transported blocks). They include micritic limestone, calcarenite with abundant broken foram tests, and bioclastic limestone containing large benthonic foraminifera and red algae, and locally miliolidae. Paleontological evidence (Appendix A) suggests that the lower (exposed) part of Member A was deposited during the P13–14 interval, constraining its age between 39.8 and 38 Ma, i.e. late Middle Eocene, and that the upper part was formed during the Late Eocene. Member B is exposed in S. Matajang and S. Mate, and has been traced several km to the south of S. Matajang. It consists of a series of interbedded volcanic breccias (dominant), lava and tuffs with intercalations of volcaniclastic sandstone, mudstone and limestone. Its thickness in S. Matajang is estimated to be at least 650 m. The breccias are poorly sorted and monomict or polymict; both clast and matrix supported varieties occur. Some breccias contain minor amounts of sedimentary carbonate clasts that show soft sediment deformation, suggesting mobilization from within the same sedimentary environment. The beds (1–10 m thick) are massive or well stratified, with coarse and fine grained layers alternating. Fragments are usually less than 5 cm in diameter. Some beds have coarser bases with clasts up to 60 cm in diameter. In S. Mate, the breccias show in part hyaloclastic features. Lava lithologies (breccia, tuff and sandstone fragments, and lava flows) include olivine basalt, clinopyroxene ± hornblende andesite, clinopyroxene ± hornblende quartz andesite, and biotite–clinopyroxene andesite. The breccias and lavas in S. Mate are commonly red brown. This colour is caused by pervasive hematite formed after ultra fine-grained magnetite. The hematite is interpreted to be a result of rapid oxidation through extrusion into a subaqueous environment. The pyroclastic rocks include crystal-lithic tuff, crystal-vitric tuff, and pumice tuff. The crystal component of these rocks consists of broken feldspar, clinopyroxene, biotite, and hornblende. Limestone intercalations within the volcanic sequence, which are particularly well developed in the S. Matajang area, vary in thickness from 5 m to 110 m. The rocks are bioclastic packstones consisting of debris derived from red algae (very common), green

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Fig. 2. Schematic geological cross sections of selected traverses. Scale 1:15,000.

algae (variable), bryozoa (variable), echinoderm, corals (rare), and molluscs, together with abundant benthic foraminifera, including miliolids, and rare globigerinids. The limestone intercalations in Member B commonly contain Pellatispira provalae, indicative of the upper part of zone Tb, with the exception of B6 in the bottom part of the sequence, which is characterized by the abundant presence of Nummulites javanus

(Appendix A). The latter species was long regarded to be a Middle Eocene guide fossil in SE Asia, but in SW Sulawesi it may also occur in the basal part of the Upper Eocene (van Leeuwen, 1981; Crotty and Engelhardt, 1993). At the top of the sequence sample B14 contains Pellatispira morphotypes close to P. absurda, which may be considered as an evolutionary ad-form of the Pellatispira lineage, indicating a biostratrigraphic position at the top end of Tb (A. Won-


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Fig. 2 (continued)

ders, written comm., 1996). This evidence indicates that the age of Member B spans the entire Upper Eocene. In addition, several larger ‘‘isolated” Upper Eocene limestone blocks are found along the fault contact between the Bone and Salokalupang Groups in S. Karopa, S. Deko and S. Deko Kanan. They are bounded to the west by either the Karopa Formation or the Matajang Formation. We have included these limestones tentatively in the Matajang Formation, but in view of their tectonic position it is also conceivable that they are allochthonous. The fossil assemblages in the lower part of Member A suggest that the sedimentary rocks were deposited in an open marine environment. The clastic material was largely derived from volcanic sources. Its immature nature indicates a first-cycle erosional origin. The absence of proximal volcanic deposits suggests that these sources were located some distance away. Cross-bedding channelization, graded bedding, parallel lamination, and soft sediment

deformation features shown by the sandstones suggest rapid deposition in a subaqueous environment, probably (in part) from gravity driven currents in a slope setting. The thicker mudstone intervals represent periods of quiescence that were occasionally interrupted by subaqueous debris flows. More or less continuous influx of clastic material hindered development of carbonate deposits, with the exception of Middle Eocene limestones exposed in the lower reaches of S. Deko. Around the boundary between the Middle and Late Eocene significant calc-alkaline volcanic-deposits (lava flows, pyroclastics, volcanic breccias) started to develop. Volcanic activity continued up to the end of the Late Eocene. The proximal nature of the volcanics in Member B suggests that the eruptive centres were nearby. Evidence for a subaqueous environment of deposition of the volcanic rocks includes the presence of carbonate clasts showing soft sediment deformation in some breccias, common limestone

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Fig. 3. Chronostratigraphic diagram for the Bone region, SW Sulawesi. Time scale drawn from data in Gradstein et al. (2004). Planktonic foraminiferal and nannofossil zones use the nomenclature of Blow (1969, 1974), as modified by Bolli et al. (1985) and Okada and Burky (1980) respectively.

intercalations, and quenching and pervasive hematite alteration of lava flows. During the Late Eocene deposition of clastic material continued (upper part Member A). It was similar in character to the lower part of Member A, the main difference being the appearance of subordinate lava flows and limestone beds and lenses, suggesting an interdigitating relationship with Member B.

The abundant presence of large benthonic foraminifera, coralline algae, echinoderms etc. in the limestones found in the upper part of Member A and in Member B indicate an inner to middle neritic environment of deposition, including fore-reef and carbonate shoal. The commonly fragmented nature of the fossils suggests prevalent medium to high energy conditions.


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3.1.2. Karopa Formation The Karopa Formation is present in S. Deko, the catchment area of S. Karopa, S. Matajang and S. Linrung. It is a least 450 m thick. The unit is made up of a well-bedded sequence of limestones alternating with commonly calcareous mudstones and siltstones, and subordinate calcareous sandstones. Bed thicknesses range mostly between 2 cm and 65 cm. In the Deko and Karopa area the formation is in contact with the Deko Formation to the east and the Baco Formation to the west. The contacts are interpreted to be faults. In S. Matajang and S. Linrung the formation is exposed in narrow fault slices sandwiched between rocks belonging to the Matajang Formation. The limestones are light to medium grey and comprise predominantly fine to medium grained calc-arenites. They contain abundant fragments of coral, coralline algae and benthic forams. The mudstones and siltstones are grey to brownish, and are commonly rich in planktonic forams. Small benthic species are usually present in subordinate amounts. They include Textularia sp., Buliminia sp., Bolivina sp., Textularia sp., and Pelecypoda sp. Planktonic foraminifera and nannofossil evidence (Appendix A) suggests that the formation was deposited during intervals P20–22 and CP18–19 respectively, giving a maximum age range of 30.5–23 Ma (Middle–Upper Oligocene). Sample B1096 contains a rich Globigerina assemblage, indicative of zones P20–21 (30.5–27 Ma). The presence of reworked Middle–Upper Eocene planktonic foraminifera (P14–P15) and the co-occurrence of benthic and planktonic foraminifera and their fragmental nature are interpreted to indicate an open, high energy carbonate shelf environment with an abundance of coarse-grained detritus, which is partly older material. 3.1.3. Baco Formation Baco Formation is the name given to a sequence of upper Lower–lower Middle Miocene sedimentary rocks that contain a significant volcanic component. It can be divided into two members; Member C is composed of sandstones, mudstones and mud-rich breccias, whereas Member D consists predominantly of volcaniclastic sandstones and breccias with subordinate lava. The two members have nowhere been observed in contact with each other. Paleontological evidence suggests that they are in part contemporaneous. The lower (exposed) part of Member C consists predominantly of mudstone, commonly calcareous, with intercalations of siltstone and sandstone. Some of the mudstone beds contain, in most cases sparse, fragments of rounded to angular limestone, up to 4 m in diameter, and locally also mudstone and volcanic clasts. The member is exposed in fault blocks bounded by older formations in S. Linrung and S. Matajang, and also in S. Katumpang Kanan, where it is unconformably covered by the Walanae Formation (Fig. 2). Its estimated thickness is >500 m. The upper part of Member C is exposed in the lower course of S. Deko where it is divided into two sections by faults. A detailed stratigraphic section measured by Maryanto et al. (2004) suggests that the southern sequence is about 360 m thick, and the one occurring to the north about 140 m. Four distinct sedimentary facies have been recognized: (i) claystone facies; (ii) thin bedded sandstone facies; (iii) graded sandstone facies; and (iv) bouldery mudstone facies. Their main characteristics are summarized in Appendix C. The bouldery mudstone facies is well developed in the northern sequence and the upper part of the southern sequence. Nannofossil assemblages observed in samples from the lower part of Member C suggest zones CN3–4 (Table 1), i.e. a maximum age range of 16.8–13.65 Ma (late Early–early Middle Miocene). Nannofossil and planktonic foraminifera assemblages in samples from the upper part of Member C equate to zones CN5a and N11–12 respectively (Appendix A), giving a maximum age range of 14.1–11.61 Ma (Middle Miocene).

Initial formation of Member C took place in an open marine, probably relatively deep environment under predominantly quiet conditions. Mudstone deposition was occasionally punctuated by debris flows. Conditions became subsequently less stable. We interpret the thin bedded sandstone and graded sandstone facies to represent products of low and high concentration turbidite currents (Walter, 1978), respectively. The bouldery mudstone facies probably originated from subaqueous debris flows. Paleocurrent analysis by Maryanto et al. (2004), using individual clast imbrications and current ripple lamination structures in thick-bedded turbidites and bouldery mudstone respectively, suggest slumping of slope sediments towards the southeast and subsequent transport in a northwesterly direction, presumably parallel to the long axis of the basin. Planktonic foraminifera and nannofossil assemblages indicate an open marine environment. The presence of Braarudosphaera and Micronlolithus species together with carbonaceous grains in some mudstone samples suggest reworking of shallow material. Member D unit is exposed in S. Kalupang and S. Baco (Fig. 2). It consists of a series of alternating volcaniclastic sandstones, pebbly sandstones and breccias/conglomerates with intercalations of mudstone, siltstone and lava. In the east it is in faulted contact with the Kalamiseng Formation, and in the west it is unconformably covered by the Camba Formation (S. Baco) or in faulted contact with the Tacipi Limestone (S. Kalupang). The estimated thickness of the formation is about 450 m in S. Baco, and at least 350 m in S. Kalupang. The sequence exposed in S. Kalupang consists predominantly of volcaniclastic sandstones with interbedded siltstones/mudstones, volcanic breccia and minor conglomerate. There is an increase in average bed thickness upwards the sequence accompanied by a decrease in the amount of fine-grained interbeds. Interbedded breccia and conglomerate occur as thin beds and lenses in the lower part of the sequence, increasing in thickness (up to 30 m) and frequency upwards. Lithological and sedimentary features are summarized in Appendix D. The sequence exposed in S. Baco differs in several respects from that in S. Kalupang. In the lower part a 97 m thick series of massive grey calcareous claystones is exposed, containing several intercalations of coarse to fine grained sandstones, 10 m to 100 cm thick, which show parallel lamination and graded bedding structures. Furthermore a series of lava flows are present at both the bottom and top ends of the exposed sequence. Lavas (flows; fragments) in the S. Kalupang and S. Baco sequences include clinopyroxene ± biotite andesite, olivine basalt (one variety is notable for the lack of feldspar), and biotite–clinopyroxene–hornblende ± olivine trachyte. They are porphyritic and show vitric, hyalophytic or trachytic textured groundmasses. Geochemical analyses (see below) indicate that the volcanic rocks are potassic in nature. Variably sized blocks of reworked Upper Eocene nummilitic limestone are scattered throughout both sequences. Planktonic foraminifera observed in four samples collected from S. Kalupang suggest a Middle Miocene age (zones N10–N12; 14.01–11.8 Ma; Appendix A). Seven claystone samples from S. Baco were analyzed for microfossils. Two of these contained a nannofossil assemblage that is considered to reflect an age equating to zones CN4–CN5a (Table 1), i.e. 14.5–11.8 Ma (Middle Miocene). The other five samples contained abundant but poorly preserved (mostly crystalized) planktonic foraminifera consisting of reworked Upper Eocene–Lower Oligocene and Lower Miocene assemblages. In several volcaniclastic sandstone samples collected from the middle course of S. Baco various palynomorphs were observed including Casuarina. This is an Australian taxon, which dispersed into the Sundaland region at the end of the Oligocene or Early Miocene (R.J. Morley, written comm. 2001).

T.M. van Leeuwen et al. / Journal of Asian Earth Sciences 38 (2010) 233–254

The volcaniclastic sequence exposed in S. Kalupang is likely to have been deposited in a shallow marine environment as suggested by the common presence of Orbitolina species and rare occurrence of wave base structures. The presence of unstable minerals such as olivine, hornblende and pyroxene indicates a first-cycle volcanogenic provenance of the sediments. This feature combined with the overall coarsening upwards of the clastic sequence suggests that the volcaniclastics were deposited as progradational aprons. A marine environment is also indicated for the lower part of the Baco sequence based on the abundant presence of nannofossils in the mudstone intervals. The common presence of palynomorphs higher up in the sequence may indicate nearshore or terrestrial conditions of deposition.

3.2. Bone Group 3.2.1. Kalamiseng Formation The Kalamiseng Formation (Sukamto, 1982) occupies large parts of the Bone Mountains. It is exposed over an area of about 2300 sq km. It is partly surrounded by the unconformably overlying Middle–Upper Miocene Camba Formation. Its western margin is in faulted contact with the Deko Formation, Salo Kalupang Group and Tacipi Limestone. Sukamto (1982) estimated the unit to be at least 4250 m thick. The lower part of the formation consists of pillow lavas that alternate with volcanic breccias, whereas the latter rock type dominates in the upper part and has intercalations of massive lavas (Yuwono, 1987). The rocks are of basaltic (dominant) to andesitic composition. Minor rhyolite is also present. The lavabreccia sequence contains subordinate interbeds of volcaniclastic sandstone, red calcareous mudstone and limestone, and is intruded by dykes and stocks of similar composition as the host rocks. The volcanic rocks and their intrusive equivalents (Yuwono, 1987; this issue) are composed essentially of plagioclase (andesine–labradorite) and pyroxene. The pyroxene assemblage is dominated by clinopyroxene (augite and minor endiopside) and includes small amounts of orthopyroxene (hypersthene). Olivine was observed in a few samples only. Opaque minerals (2–4 vol.%) constitute titaniferous magnetite, and lesser ilmenite, chromite and chromiferous spinel. The groundmass of the volcanic rocks is made up of an admixture of microlitic/microcrystalline plagioclase, pyroxene and opaques with variable amounts (0–20 vol.%) of glass. The igneous rocks have undergone weak to strong alteration. Secondary minerals include chlorite, clay, epidote, silica, actinolite, amphibole, albite, carbonate, and zeolite, indicating a greenschist facies assemblage. The age of the Kalamiseng Formation is poorly constrained. Sukamto (1982) suggested a possible Early Miocene age. Yuwono (1987) reported the results of K/Ar dating of four samples, which yielded ages of 33.3 ± 1.67 Ma (Eocene–Oligocene boundary) and 21.72 ± 1.09 Ma, 17.5 ± 0.88 Ma and 18.7 ± 0.94 (late Early Miocene). Yuwono et al. (1988a) considered the latter two ages to reflect the time of tectonic emplacement of the Kalamiseng rocks in view of their altered state. The 33.3 and 21.7 Ma ages were obtained from samples with a low loss on ignition (Yuwono, 1987; Table IV.7), suggesting they were not significantly affected by alteration, and hence may represent formation ages. Two samples that we collected from red mudstone intercalations yielded Oligocene nannofossil assemblages with a maximum age range of 32– 27 Ma (Appendix A). The common presence of pillow lavas indicates that the Kalamiseng Formation was deposited in a subaquaous environment, probably deep marine, as suggested by the nature of the red mudstone intercalations.


3.2.2. Deko Formation The Deko Formation is composed of a sequence of mudstones with intercalations of volcanic rocks and volcaniclastic sandstone. It is exposed in S. Karopa, S. Deko, S. Linrung, S. Matajang and S. Katumpang. To the east it is invariably in contact with the Kalamiseng Formation, whereas to the west it is flanked by various units (Fig. 2). In S. Matajang it is also present as narrow fault slices within a thick package of Eocene rocks. The Deko Formation has its widest exposure in the S. Karopa and S. Deko, where it is estimated to be at least 375 thick. The mudstones are massive to shaley, predominantly calcareous, and vary in colour from reddish and brownish red (dominant) to brownish grey, greenish grey and medium grey. Some beds show parallel lamination. Several samples consist predominantly of planktonic foraminifera and are interpreted to represent Globigerina ooze. The sandstone interbeds are generally less than 5 cm thick, but locally attain thicknesses of up to 2 m. The sandstones are fine to coarse grained and moderately well to poorly sorted and mostly clast supported. Graded bedding is a fairly common feature. The framework clast assemblages comprise relict clinopyroxene and plagioclase, lava fragments, magnetite and minor biotite. Magnetite commonly defines a primary bedding structure, including cross bedding. Lava fragments include olivine basalt and basaltic andesite, and biotite–clinopyroxene andesite/ trachyte. The volcanic intercalations vary in thickness from 5 m to 45 m. They consist predominantly of porphyritic fine-grained, dark coloured lavas, including pyroxene trachyte, pyroxene dacite and a vitric andesite that is notable for an absence of mafic minerals. Other rock types include pyroxene vitric tuff and volcanic breccia. The mudstones contain rare interbeds of graded calcarenite (<5 cm thick) and pebbly mudstone (<25 cm). Pebbles in the latter comprise mudstone, which is of similar composition and age as the host mudstones, and various altered volcanic rocks. Blocks of bioclastic limestone, containing Upper Eocene benthic foraminifera assemblages, and volcaniclastic sandstone occur scattered throughout the mudstone sequence. These blocks are usually less than 5 meters in diameter. In addition small to medium-sized blocks of diorite are embedded in the mudstones. The mudstones of S. Deko contain rich nannofossil and planktonic foraminifera assemblages, indicating an Oligocene age, with a maximum age range of 32–27 Ma (Appendix A). Reworked Eocene–Lower Oligocene nannofossil species are common. The nature of the sedimentary rocks suggests a relatively deep marine environment of deposition in which quiet conditions prevailed. Turbidites and debris flows occasionally interrupted pelitic sedimentation. 4. Geochemistry Volcanic and intrusive samples (45) from the Salokalupang and Bone Groups were analyzed for major oxides and a wide range of trace elements by fusion-inductively coupled plasma and fusioninductively coupled plasma-mass spectrometry methods respectively, using a lithium metaborate fusion rather than acid digestion to place the sample into solution. The analyses were carried out by Actlabs (Canada). For further details the reader is referred to Elburg et al. (2003). The samples have been divided into four main suites, which are named after the formation in which they occur; Matajang volcanics (Mv, 8 samples), Kalamiseng volcanics (Kv, 14), Deko volcanics (Dv, 14) and Baco volcanics (Bv, 5). Included in the Kv suite are several samples of fragments and intercalations from the Deko Formation that have geochemical characteristics similar to the Kalamiseng volcanics.


T.M. van Leeuwen et al. / Journal of Asian Earth Sciences 38 (2010) 233–254

Table 1 Whole rock analysis of representative volcanic samples. B 1029 Mv

B 1072 Mv

B 905 Bv

B 988 Bv

B 435 Kv1

B433 Kv1

B 417 Kv2

B926 KV2

B 419 Dv1

B1106 Dv1

B 318 Dv2

B325 Dv2

SiO2 Al2O3 Fe2 O3 MnO MgO CaO Na2O K2O TiO2 P2O5 LOI Total

57.98 17.59 6.59 0.11 2.44 1.61 7.37 1.68 1.10 0.23 3.01 99.7

67.93 14.88 4.48 0.07 0.78 3.48 3.92 1.31 0.51 0.2 2.41 99.9

47.81 13.31 10.81 0.19 8.78 10.83 2.02 2.03 0.66 0.26 2.89 99.6

47.47 14.79 9.99 0.29 4.69 9.01 2.42 4.5 0.74 0.62 5.4 99.9

50.8 14.36 13.16 0.18 4.33 7.33 4.64 0.65 2.31 0.01 2.51 100.3

54.79 15.82 9.64 0.14 3.97 4.03 6.19 0.67 1.61 0.02 3.05 99.9

57.6 12.83 10.32 0.16 3.25 4.94 3.85 1.38 1.57 0.28 2.59 98.8

46.93 14 8.14 0.07 5.28 10.82 3.03 0.18 0.73 0.16 10.37 99.72

61.72 15.39 3.35 0.25 2.98 2.35 5.52 0.52 1.18 0.35 2.74 98.7

59.07 16.55 6.18 0.29 4.22 2.54 6.18 0.12 1.12 0.38 3.2 99.85

48.24 15.72 10.42 0.26 4.11 6.1 6.02 2.31 0.79 0.57 5.78 100.3

46.54 12.83 11 0.29 7.18 8.88 4.91 0.83 0.73 0.73 6.05 99.96

Sc V Cr Co Ni Cu Zn Ga Ge Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U

23 122 28 27 20 15 104 18 2 46 284 25 134 3 1.7 388 11.5 29.2 3.7 16.5 4 1.32 4.6 0.8 4.5 0.9 2.6 0.4 2.6 0.4 3.4 0.2 6 2.2 0.6

8 73 284 3 20 37 60 18 1 23 333 18 174 6 76.6 230 26.5 57.7 6.22 22.5 4.6 1.44 4.2 0.6 3.4 0.6 1.8 0.27 1.6 0.24 4 0.5 9 5 1

33 312 355 43 63 161 98 16 2 28 1.020 14 37 1 0.7 559 5.7 13.5 1.76 8.5 2.2 0.82 2.5 0.4 2.5 0.5 1.5 0.22 1.4 0.22 1.2 0.1 12 1.4 0.6

28 298 54 27 20 106 118 17 2 100 2.160 19 73 2 304 1490 15.5 34.6 4.08 17.9 4.2 1.4 4.3 0.7 3.6 0.7 2 0.3 1.9 0.3 2.1 0.1 36 5.4 2.4

36 351 52 35 25 21 88 19 2 19 202 22 145 3 0.9 31 1.6 4.2 0.73 4.1 1.8 1.17 2.7 0.6 3.8 0.9 2.7 0.45 3.1 0.47 3.7 0.21 5 0.3 0.5

27 303 24 23 10 10 67 16 1 13 800 17 141 3 0.5 86 2.7 7.2 1.14 5.8 2 0.88 2.5 0.5 2.9 0.6 1.9 0.29 1.9 0.29 3.4 0.2 6 4 0.2

26 338 121 22 20 15 109 17 2 26 310 42 94 3 0.6 46 8 18.3 2.79 13.5 4 1.7 5.6 1 6.2 1.3 4.1 0.63 3.8 0.58 2.4 0.1 5 0.5 0.3

36 215 312 28 47 27 79 15 1 5 490 17 59 1 0.5 64 5.5 14.7 1.86 8.6 2.4 0.89 2.8 0.5 3 0.6 1.8 0.27 1.7 0.26 1.7 0.1 5 1.2 0.2

18 42 40 5 15 13 91 23 1 8 224 79 364 5 0.5 140 17 47 7.04 32 9.7 2.99 12 2.1 13 2.8 8.4 1.23 7.6 1.19 8.6 0.4 5 1.7 0.5

16 45 58 10 20 10 133 28 1 3 266 82 358 7 0.5 59 17.2 55 7.12 34.1 9.9 3.33 12.8 2.4 14.4 2.9 8.9 1.37 8.4 1.27 8.8 0.4 5 2 0.7

41 340 64 25.1 67 111 60 17 1 34.1 273.6 21 74 2 5.8 586 16.6 34.4 4.06 20.1 4.9 1.68 3.9 0.7 3.7 0.7 2.1 0.29 2 0.3 2 0.1 26 6.3 1.9

39 236 152 38 41 149 61 15 1 35.2 578.4 21 84.7 2 6.6 1.200 18.3 52.8 6.16 30.1 6.7 2.42 5.7 0.8 3.8 0.7 2.1 0.25 1.5 0.23 2.2 0.11 54 9 3

Abbreviations: Mv = Matajang volcanics, Bv = Baco volcanics, kV = Kalamiseng volcanics, Dv = Deko volcanics.

Analyses of representative samples from each suite are given in Table 1. All samples have been plotted on (1) the total alkali versus silica (TAS) diagram of Cox et al. (1979) to classify and to divide them into alkaline and subalkaline series using the boundary of Kuno (1966) (Fig. 4a); (2) the K2O versus SiO2 (wt.%) diagram (Peccerillo and Taylor, 1976) to subdivide the samples into low-K (tholeiitic), medium-K (calc-alkaline) and high-K types (Fig. 4b); (3) the Th/Yb versus Ta/Yb diagram of Pearce (Fig. 4c); and 4) the discrimination diagram of Wood (1980) (Fig. 4d). In addition the Kv samples were plotted on the Ti–V diagram compiled by Shervais (1982) (Fig. 4e) and the Cr–Y diagram of Pearce (1982) (Fig. 4f) with the objective to determine whether they are MORB or IAT-like. We also plotted selected samples from each suite as extended spidergrams (Fig. 5). For comparison we have added an example of typical arc basaltic andesite from Ijen (Handley et al., 2007) in Java, and a subduction zone basalt with a low recycled flux (weak subduction signature) from Lau Basin (Pearce et al., 1994). The Salokalupang Group samples (Mv and Bv) display features that are typical of arc volcanics. On the extended spidergram they are characterized by Nb–Ta throughs, LILE-, LREE- and Sr-enrich-

ments, and moderate HREE-depletion, similar to the pattern shown by the Ijen-arc sample (Fig. 5). The samples plot mostly in the continental margin arc field (Fig. 4 c and d). Zr/Y values are greater than 3, suggesting a continental crust enrichment (Pearce and Norry, 1979). The Mv samples have a composition range from basaltic to rhyolitic, and belong to the subalkaline series, with the exception of two trachyandesite samples (Fig. 4a). They have low to medium-K contents (Fig. 4b). In contrast, the Bv samples are high-K to shoshonitic and alkalic, and have a restricted SiO2 range (Fig. 4a and b). The Bone Group samples (Kv and Dv) differ from the Salokalupang Group in lacking continental margin arc signatures, but instead are MORB- to IAT-like. The Kv samples are predominantly basaltic in composition and low-K (Fig. 4a and b). In the Wood diagram they plot in the tholeiitic arc and N-MORB fields (Fig. 4d). Similarly in the Cr–Y diagram of Pearce (1982) they overlap both fields (Fig. 4f). In the Ti–V discrimination diagram of Shervais (1982) the samples plot in the MORB and back-arc basin basalt (BABB) fields (Fig. 4e). In the Th/Yb versus Ta/Yb diagram they straddle the boundary between the N-MORB and tholeiitic arc fields (Fig. 4c).


T.M. van Leeuwen et al. / Journal of Asian Earth Sciences 38 (2010) 233–254





Alkalic 5


(Shoshonite series)



Trachyte P-T Benmorite



Phonolite 12 9

Rhyolite Mugearite B+T Trachyandesite Hawaiite Dacite 6 Nephelin

High-K (High-K (calc-alkaline) series)




(Calc-alkaline series)

B-A Andesite 3



0 35






(Low-K (tholeiite) series)

0 40











Legend Bv Hf/3 SH

Dv 1 Dv 2



Kv 1





Mv D1





Kv 2

r ic h s o ed ur m a ce n tl


De p


le t s o ed ur m a ce n t le

B 0.1

.01 .01

D2 C

A = Active continental margin B = Oceanic island arc 0.1





Ta/Yb 650 600 550 500 450 400 350 300 250 200 150 100 50 0


10 ARC < 20 > OFB

1400 1000 MORB

Continental Flood basalt

Arc tholeiite

50 MORB & BAB Ocean-island & alkali basalt


Cr (ppm)




Calc-alkali basalts

10 0










Y (ppm)

Fig. 4. (a) Chemical classification and nomenclature of igneous rocks using the total alkalis versus silica (TAS) diagram of Cox et al. (1979). The curved solid line subdivides the alkalic from subalkalic rocks. (b) K2O versus silica diagram. The diagram shows the subdivisions of Le Maitre et al. (1989) (broken lines with nomenclature in italics) and of Richwood (1989) (nomenclature in parentheses). The shaded bands are the fields in which fall the boundary lines of various authors as summarized by Richwood (1989). (c) Th/Yb versus Ta/Y plot of Pearce (1983). Dashed lines separate the boundaries of the tholeiitic (TH), calc-alkaline (CA) and shoshonitic (SH) fields. (d) Hf/3–Th–Ta tectonomagmatic discrimination diagram from basalts and more differentiated rocks (after Wood, 1980). Island-arc tholeiites plot in field D1 and calc-alkaline continental margin volcanics in field D2. The boundary between the two fields is transitional. (e) Ti–V discrimination diagram from basaltic rocks (after Shervais, 1982). MORB = midoceanic ridge basalt; BAB = back-arc basin basalt. (f) Cr–Y diagram for basalts (after Pearce, 1982), showing the fields for MORB, volcanic-arc basalt (VAB) and within-plate basalts (WPB).


T.M. van Leeuwen et al. / Journal of Asian Earth Sciences 38 (2010) 233–254

The Kv samples can be divided into two groups, Kv1 and Kv2. The first group have trace element patterns that are relatively flat for the less incompatible elements (Fig. 5). They all show some degree of curvature down towards Nb and Ta, and then are more enriched in the most incompatible elements. There are positive spikes for K and Sr. These patterns may be interpreted in two ways: (i) the basalts originated remote from any subduction zone, i.e. they are true MORB, and were subsequently altered resulting in an increase of the most mobile elements; and (ii) they formed in a subduction zone with a low recycled flux. The first scenario would require a very depleted mantle source and/or a very high degree of melting in order to obtain the very low concentrations of high field strength elements (HFSE). This is in itself not necessarily a barrier to a ‘‘subduction-free” origin. However, the Th and U concentrations are significantly higher than those normally observed in MORB volcanics. The second interpretation is supported by similarities in trace element patterns with the Lau Basin basalt (i.e. a low recycled flux basalt), especially in the moderately and less compatible elements. Macpherson (written comm., 2007) notes that ‘‘while it would be reasonable to assume that some arcs may have a low recycled flux at any time in their history, the earliest stages of subduction might be a time when the recycled flux is just becoming established”. The Kv2 samples are less depleted than the first group and appear transitional between the low (Lau-type) and typical subduction (Ijen-type) cases. The mobile elements, in particular, are significantly more enriched. In the ‘‘altered MORB” model this would require very significant variations in the degree of partial melting and/or the prior depletion of the mantle from which the crust was extracted, and would also require that the crust extracted from the least depleted mantle had the greatest inventory of mobile elements added to during alteration. The ‘‘subduction-related” model would require that the mantle structure was heterogeneous in terms of its depletion/melting, and then that fluid flux was variable. This would appear quite reasonable in terms of our understanding of active magmatic systems in subduction zones (Macpherson, written com., 2007). The Deko volcanics can also be subdivided into two groups, Dv1 and Dv2, based on their distinct geochemical characteristics, e.g. the Dv1 samples have higher Na2O (5.5–7.3% versus 3–6%), Nb (5–7 ppm versus 1–2 ppm), Hf (6.3–9.7 ppm versus 2.0–2.4 ppm), and Zr (255–416 ppm versus 70–101 ppm) contents. The Dv1 samples are of basaltic andesite to dacite composition (54–69 wt.% SiO2) (Fig. 4b). Their geochemistry is rather unusual. They have very high and constant concentrations of REE, which differentiate them from typical arc lavas such as Ijen. They do have some ‘‘spikeyness” to their patterns, but in a somewhat inconsistent manner (Fig. 5). Some HFSE (Nb, Ta and Ti) are relatively depleted while Zr and Hf are enriched. There is also some enrichment of mobile elements but this is not particularly pronounced. The major elements are reasonably consistent with differentiation of a subduction zone melt at moderate to low pressure, except that they have relatively high MgO with respect to SiO2. Furthermore the high, relatively constant REEs over a range of SiO2 are difficult to reconcile with a subduction model. Two possible scenarios to explain the unusual geochemistry are: (i) the rocks are the product of considerable contamination of arc magma by arc crust, or (ii) they are derived purely from crustal melts. In the latter model various degrees of melting of a basaltic precursor would explain the major element variations and the flat but elevated trace element concentrations. The negative HFSE spikes might relate to an odd residual phase (rutile), while the moderate enrichment of the mobile elements could reflect that the source material of the melts was either formed in a subduction zone or moderately altered.


Salokalupang Group



B 905 B v (M id M io cene) B 978 B v (M id M io cene) B 432 M v (Eo cene) Lau Ijen

100 10 1 0.1 0.01


Cs Ba U Ta La Pb Sr Nd Hf Eu Gd Dy Ho Tm Lu Rb Th Nb K2O Ce Pr P2O Zr S Ti Tb Y Er Yb 10000

Bone Group


1000 100

B 409 Kv1(Oligo -M io cene) B 416 Kv2 (Oligo -M io cene) B 1105 Dv1(Oligo -M io cene) B 1017 Dv2 (Oligo -M io cene) Lau Ijen

10 1 0.1 0.01


Cs Ba U Ta La Pb Sr Nd Hf Eu Gd Dy Ho Tm Lu Rb Th Nb K2O Ce Pr P2O Zr S Ti Tb Y Er Yb Fig. 5. MORB normalized trace element patterns for selected samples. Normalizing values and order from Sun and McDonough (1989).

The Dv2 samples (7) are basalts and basaltic andesites, mostly low-K, that plot in the tholeiitic island arc field on the Wood diagram (Fig. 4d). Their geochemistry is quite similar to the Ijen arc sample (Fig. 5). They show similar enrichment and depletion patterns to this (and other) arc lavas. They probably resulted from melting of mantle wedge material that had been fluxed by fluid derived from the subduction of oceanic lithosphere and sediments. The more incompatible elements are slightly more enriched than at Ijen, while the less incompatible elements are more depleted. This ‘‘rotation” of the pattern relative to typical arc lavas could result from slightly different recycled fluxes. Alternatively, elevated incompatible elements could reflect differentiation at greater depth than typical arc lavas, which could occur at any stage during the history of an arc, but most likely during the earliest stages when it is difficult for melts from the young mantle wedge to reach the shallow (crustal) lithosphere (Macpherson, 2008). In conclusion, the Salokalupang and Bone Group volcanics show evidence for a range of subduction inputs. These vary from very small (Kv1) to moderate (Kv2) to typical arc (Mv, Bv, Dv2). The exception is the Dv1 rocks, which do no appear to record the history of mantle processes that took place beneath the area where DV2 magmatism occurred, but instead reflects heating and melting of a basaltic precursor in the crust. The Mv and Bv rocks are typical

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continental margin arc volcanics. They are geochemically very similar to the Eocene calc-alkaline and Middle Miocene potassic volcanics respectively in the Western Divide Mts (Letellier et al., 1990; Elburg et al., 2002). The Kalamiseng volcanics are likely to have been formed in a back-arc basin or incipient arc setting. The latter setting is also indicated for the Dv2 rocks. 5. Structure The geology of the western margin of the Bone Mountains shows a prominent NNW–N structural trend. It reflects the trend of the East Walanae fault system, which is about 2–2.5 km wide, and faults cutting the Kalamiseng Formation. It is also the dominant strike direction of the sedimentary strata of the Salokalupang Group. The most eastern fault of the East Walanae system forms the boundary between the Kalupang and Bone Groups. In several localities the contact is exposed as a narrow zone of highly disturbed sedimentary rocks, showing brecciation and shearing. The faults that constitute the East Walanae system cut the Kalupang Group into a series of tectonic slices. Faults bounding these slices are rarely exposed, but their existence can be deduced in a number of cases from the juxtaposition of rocks of different ages, abrupt lithological changes and/or sudden changes in strike shown by sedimentary strata. In rare cases where they have been observed in the field they appear steeply dipping to vertical. On the cross sections (Fig. 2) interpreted faults are shown with a similar attitude. Difficulties in correlating between sections suggest that the fault slices consist of variably sized blocks/lenses rather than of slices that have more or less constant widths over a considerable strike length. Evidence for faults in the Kalamiseng Formation includes fault breccias, slicken-slides, aerial photo lineaments, and geomorphological indications such as cliff faces and sudden changes in river courses. The latter features suggest young movements. In the upper Matajang spacing between faults is of the order of 300–500 m. As mentioned above, the East Walanae fault system is believed to be a sinistral strike-slip fault along which significant vertical movements have also taken place. In the study area it has not been possible to establish direction and amounts of fault movements. To the east the system separates two units of different origin, to the west the Walanae Formation largely obscures older formations, and within the fault system itself strata trend predominantly parallel to the fault zone, thus hampering correlation across faults. The uppermost Miocene–Pliocene Walanae Formation commonly shows variable strikes and dips close to the fault system, suggesting Pliocene and/or younger movements. The strata of the Salokalupang Group dip predominantly E/ENE. Dips are moderate to steep. Evidence of overturning of beds has been locally observed. In places dips and strikes deviate significantly from the general trend, which is partly related to small to moderate scale folding and fault movements. The folds trend roughly NNW and are arranged en echelon. As suggested by Sukamto (1982) they may have been formed as the result of strike-slip horizontal movements along the East Walanae fault system. A large SE plunging anticlinal structure has been interpreted to be present in S. Kalupang (Sukamto, 1982; Maryanto et al., 2004). Another possible interpretation is that the observed change in dip direction is caused by faulting. The latter interpretation is supported by lithological differences in the westerly and easterly dipping sections: i.e. breccias are more prominent in the latter. It should be noted that in S. Baco beds dip consistently to the west. 6. Discussion The Salokalupang and Bone Groups represent two contrasting rock sequences, which are partly contemporaneous.


The Salokalupang Group spans the period late Middle Eocene– Middle Miocene with apparent gaps in the Early Oligocene and (Late Oligocene) – Early Miocene. Lithologically it can be correlated with deposits of the same age group found in the Western Divide Mountains and we therefore interpret it to form part of the SE Sundaland continental margin domain. The Middle Eocene section (lower part Member A, Matajang Formation) is characterized by thick deposits of volcaniclastic material, that was probably derived from a volcanic arc located along the eastern margin of the present Western Divide Mountains (see below). During the Late Eocene the volcanic arc appears to have shifted to the east as evidenced by the presence of proximal volcanics in the upper part of the Matajang Formation (Member B) and apparent lack of uppermost Eocene volcanics further to the west (van Leeuwen, 1981). It is generally accepted that in the Western Divide Mountains region the calc-alkaline volcanic activity had ceased by the end of the Eocene (e.g. t’Hoen and Ziegler, 1917; van Leeuwen, 1981; Yuwono et al., 1988a), but some authors have suggested that it continued into the Oligocene in the Bone Mountains region (e.g. Wilson, 2000). We did not find any evidence in support of the latter assumption. The Oligocene Karopo Formation does not contain contemporaneous volcanic material and, as discussed below, the volcanics present in the Kalamiseng and Deko Formations were not formed in a continental margin environment. The absence of Oligocene calc-alkaline volcanic rocks in the stratigraphic record of SW Sulawesi has been variously interpreted to indicate a change of active continental margin to strike-slip (transform) margin (Polvé et al. 1997; van Leeuwen and Muhardjo, 2005), steepening of the subduction zone (Guntoro, 1999), tectonic erosion of the volcanic arc (Polvé et al., 1997), or a change from west to east dipping subduction polarity (Parkinson, 1991). All these scenarios assume that the Eocene volcanics formed above an active subduction zone in an arc setting, a view that is widely accepted (e.g. van Leeuwen, 1981; Yuwono et al., 1988a; Soeria-Atmadja et al., 1998), and supported by the geochemical characteristics of the Eocene volcanics in SW Sulawesi (Elburg et al., 2002; this issue). The oldest recorded Neogene rocks of the Salokalupang Group are represented by the lower part of the Baco Formation (Member C) that consists predominantly of mudstone and was formed some time between 16.8 and 13.65 Ma, i.e. late Early Miocene–early Middle Miocene. The common presence of turbidites and debris flows in the upper part of Member C indicates that conditions of sedimentation became less stable during the early Middle Miocene, reflecting the initiation of a tectonic event. This event roughly coincided with the onset of a period of potassic magmatism throughout Western Sulawesi around 13–14 Ma ago (Polvé et al., 1997; Elburg et al., 2002, 2003), which in the Bone Mountains region resulted in thick deposits of volcaniclastic and minor lavas and pyroclastics (Member D, Baco Formation). The tectonic event culminated with the deformation of the Salokalupang Group assemblages, which were tilted, locally folded and sliced up into blocks of varying sizes along a major northerly trending fault zone (East Walanae fault system) prior to the deposition of the unconformably overlying Camba Formation. The basal part of this formation is older than 11.8 Ma (upper limit zone N12) and the upper part of the Baco Formation is younger than 14.1 Ma. This constrains the timing of the deformation to around 13–12 Ma. Turning to the Bone Group, its age is less well constrained than that of the Salokalupang Group due to limited available paleontological and radiometric age data. These suggest that the rocks were deposited at least in part during the Oligocene, and possibly also during the Early Miocene. The Kalamiseng Formation was originally thought to represent the remnants of a calc-alkaline arc of either Eocene (van Leeuwen, 1981) or Early Miocene age (Sukamto,


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1982). Nishimura et al. (1980), who undertook a gravity survey along the northern edge of the Bone Mountains, interpreted the results to indicate that the area might be underlain by oceanic crust. Further support for this hypothesis was lent by the work of Yuwono (1987), which included whole rock analyses and analysis of clinopyroxene phenocrysts hosted by basalt, showing a N-MORB-like, non-orogenic signature. This suggests that the formation forms the upper part of an ophiolite suite (Yuwono et al., 1988a; Letellier et al., 1990). As mentioned above, we believe that the Deko Formation is closely related to the Kalamiseng Formation in time and space, for the following reasons: (i) the Deko Formation is everywhere bounded on its eastern side by the Kalamiseng Formation; (ii) both units are of similar age; (iii) red mudstone is the dominant sedimentary rock type in both units; and (iv) the Deko Formation contains fragments and possibly intercalations of igneous rocks, which geochemically and petrographically are very similar to the Kalamiseng rocks. It is now generally recognized that the majority of ophiolites formed above subduction zones, especially in fore-arc environments (e.g. Pearce, 2003; Pubellier et al., 2004). This class of ophiolites is known as ‘‘supra-subduction zone” (SSZ) ophiolites (Pearce et al., 1984). A similar setting is suggested for the Bone Group rocks based on the following evidence: (i) the volcanic rocks show weak to strong subduction signatures; (ii) hemapelagic mudstones are the dominant sedimentary lithology; this facies is commonly associated with marginal basin ophiolites, whereas chert and reefoidal limestone are more characteristic features of midoceanic ridge and oceanic plateau environments respectively (Xenophontos and Osozawa, 2004); and (iii) a relatively short travel time is indicated for the Bone Group to move to their present position from it place of formation (maximum 15 million years), which is more typical of subduction related ophiolites than of seamount/oceanic plateau or oceanic ridge rocks (Xenophontos and Osozawa, 2004). The presence of plutonic fragments with a Kalamiseng volcanics geochemical signature in the Deko Formation suggests that plutonic rocks of the ophiolite sequence were exposed at the seafloor, supplying debris from submarine fault scarps, a feature that is common in fore-arc environments (Hall, 1990). Boninitic volcanics, a key indicator of a fore-arc setting, have not been identified among the Bone Group volcanics. Assuming they are indeed absent, this may indicate that temperatures or water contents were too low, or pressures too high, to create boninitic magmatism (R. Hall, written comm., 2009; Bloomer et al., 1995). The Bone Group was juxtaposed against the Salokalupang Group before the deposition of the Camba Formation, as this unit overlaps both sequences. Most likely it took place during the same tectonic event that faulted and tilted the Salokalupang Group rocks, i.e. around 13–12 Ma. This is about 5–6 Ma later than proposed by Yuwono et al. (1988a). The two Early Miocene K/Ar dates interpreted by them to reflect the emplacement age of the ophiolite may instead be the result of partially resetting of the radiogenic clock, or possibly represent formation ages. About 150 km to the north of the Bone Mountains another exposed oceanic fragment is exposed juxtaposed against Sundaland basement, and has been referred to as the Lamasi Volcanics (Djuri and Sudjatmiko, 1974) or Lamasi Complex (Bergman et al., 1996; Barber and Simandjuntak, 1996) (Fig. 1C). This unit consists predominantly of basaltic-andesitic lavas and dykes, with minor rhyolite and rare intercalations of red mudstone, chert and volcaniclastics, intruded by gabbro (e.g. Simandjuntak et al., 1991; Priadi et al., 1994; Bergman et al., 1996). Barber and Simandjuntak (1996) report also the occurrence of serpentinized peridotite, layered gabbro, troctolite and sheeted dyke complexes, suggesting that a complete ophiolite suite may be present.

Published K/Ar and Ar/Ar dates for 17 volcanic and intrusive samples (Bergman et al., 1996; Polvé et al., 1997) and two unpublished Rb/Sr dates of around 40 Ma obtained from gabbro samples (A.J. Barber, pers. comm., 2006) fall into three groups (excluding excessively old ages): Late Jurassic–Early Cretaceous (162– 120 Ma), Late Eocene–Middle Oligocene (40–28.5 Ma), and Early Miocene (24–15 Ma). Bergman et al. (1996) interpreted their data to indicate Cretaceous–Paleocene ophiolite crystallization followed by Eocene–Oligocene obduction and Miocene exhumation. Polvé et al. (1997) proposed two magmatic events, i.e. Late Jurassic and Late Eocene–Oligocene, and considered the youngest ages that they obtained (around 15 Ma) to reflect tectonic emplacement or deformation ages. The Lamasi rocks display geochemical signatures that are similarly to the Kalamiseng volcanics in that they range between MORB and IAT (Priadi et al., 1994; Priadi et al. (1997); Polvé et al., 1997; Bergman et al., 1996). Several authors (Yuwono et al., 1988a; Bergman et al., 1996; Polvé et al., 1997; Kadarusman et al., 2004) have suggested that the Lamasi and Bone segments constitute fragments of the East Sulawesi Ophiolite (ESO), which has been variously interpreted to represent trapped Indian Ocean crust (Mubroto et al., 1994; Bergman et al., 1996), Celebes Basin Crust (Monnier et al., 1995), a seamount (Barber and Simandjuntak, 1996) or oceanic plateau crust subsequently overprinted by magmatism in different environments (Kadarusman et al., 2004). None of these scenarios take into account that the Sulawesi ophiolite is probably composite (Hall, 2002) indicated by two or more distinctive age groups obtained from the ESO (Simandjuntak et al., 1991; Monnier et al., 1994) and Lamasi ophiolite, suggesting different tectonic settings of formation for different parts. It would seem reasonable to assume that the Bone segment and the Cenozoic portion of the Lamasi segment share common formation and tectonic histories in view of similarities in geochemical characteristics and age, and their proximity to each other. They may possibly form a single terrane. Gravity modeling of the East Sengkang Basin located about halfway between the two segments suggests the presence of a high density basement (Panjaitan, 1998). We propose that the Bone and Tertiary Lamasi ophiolites formed during the Oligocene and Early Miocene in a transtensional basin, bordered to the west by a transform passive continental margin, i.e. Sundaland, where, subduction-generated calc-alkaline volcanism had ceased by the end of the Eocene, and to the east by a newly initiated, westerly-dipping subduction zone. The Mesozoic portion of the Lamasi Complex may represent the remnant of oceanic crust subducted beneath Sundaland prior to the Oligocene. Many authors have proposed that ophiolites are predominantly emplaced onto a continental margin during a collision event (e.g. Coleman, 1971; Dewey, 1976; Ben-Avraham et al., 1982; Milsom, 2003). This is the tectonic setting favoured by Bergman et al. (1996) for the Lamasi Complex. However, there is little sign that Western Sulawesi was affected by collision events taking place further to the east during the Early and Middle Miocene (Hall and Wilson, 2000; Hall, 2002; Calvert and Hall, 2003, 2007). Typical collision-related features such as a regional unconformity, widespread thrusting and folding, and syn-post orogenic sediments are absent in the Cenozoic record of SW Sulawesi (e.g. van Leeuwen, 1981; Wilson, 1995). Further north such features date from the Early Pliocene (Calvert and Hall, 2003; Hall, 2002; van Leeuwen and Muhardjo, 2005). Karig et al. (1986) and Hall (1990) have proposed an alternative mechanism whereby strike-slip faults juxtapose ophiolitic fragments (and other terranes) against continental assemblages along high-angle faults. They observe that this mode of terrane emplacement may be more common than generally believed, particularly in orogenic belts developed in response to oblique convergence. Examples of this type of emplacement can be found in the Philippines (Karig et al., 1986),

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Halmahera (Hall, 1990), and Sumatra (Page et al., 1979). A similar scenario may be applicable to the Western Sulawesi ophiolite fragments. The Bone fragment is juxtaposed against the Sundaland


margin along the East Walanae Fault, which, as discussed above, is a high-angle fault rather than a thrust fault dipping at a low angle, and has been interpreted to represent a sinistral strike-slip

Fig. 6. Schematic sections illustrating the Cenozoic geologic evolution of SW Sulawesi. (a) Mid-Eocene: regional extension; grabens filled with siliclastic sediments; formation of Makassar Straits; development of volcanic arc in the central part of SW Sulawesi. Late Eocene deepening of Makassar Straits; development of limestone platform; arc shifts to the east. (b) Oligocene: Cessation of subduction and volcanism; continuation of carbonate deposition; formation of marginal basin to the east of SW Sulawesi; initiation of new subduction zone at eastern margin of basin triggering volcanic activity; deposition of hemipelagic sediments; Early Miocene: continuation of carbonate deposition along passive margin; continuing volcanic activity in marginal basin. (c) Mid-Miocene: major tectonic event; block faulting of carbonate platform and older units; initiation of potassic volcanism; formation of Walanae Fault Zone; juxtaposition of marginal basin sequences against continental margin successions. (d) Late Miocene: continuation of potassic volcanism; formation of Bone Basin and Walanae Depression; Pliocene: tectonic event having main impact to the north; in SW Sulawesi effects included movements along main lines of weakness and widespread uplift.


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fault with a significant vertical component, although horizontal movement is difficult to prove in the field. The nature of the contact between the Lamasi ophiolite and the Sundaland basement is generally believed to be largely a thrust contact (e.g. Coffield et al., 1993; Bergman et al., 1996; Barber and Simandjuntak, 1996), an interpretation that is consistent with its outcrop configuration. The thrust structures could have subsequently developed during the Early Pliocene tectonic event with initial high-angle shear zones evolving into low-angle shears in a manner described by Karig et al. (1986) from the Philippines. Alternatively, they may have formed during the initial emplacement of the ophiolite fragment in a restraining bend in a strike-slip fault.

7. Geological evolution of SW Sulawesi In this section we present our interpretation of the geological evolution of SW Sulawesi based on published data and the results of our study. The Cenozoic part of the geologic evolution is illustrated in Fig. 6. The history of SW Sulawesi starts in the Late Jurassic–Early Cretaceous when northeasterly directed subduction of Meso-Tethys oceanic lithosphere took place beneath the Sundaland margin resulting in the development of an Early Cretaceous continental arc in South-Central Kalimantan (e.g. Hamilton, 1979; Parkinson et al., 1998; Hall, 1996, 2002). Based on the observation that eastern Kalimantan and Western Sulawesi share Cretaceous and Early Paleogene features, Hamilton (1979) suggested that the two regions were positioned closer together during this time interval. Similarities in basement structure and paleomagnetic data (Haile, 1978; Cloke et al. (1999); Guntoro, 1999; Fuller et al. 1999) support this hypothesis. Around 120–115 Ma, several continental fragments derived from Gondwana arrived at the subduction zone (Metcalf, 1996; Parkinson et al., 1998). In SW Sulawesi, one of these fragments was subducted to a depth of 65–85 km based on P–T estimates of eclogite at Bantimala (Miyazaki et al., 1996). Jurassic shallow marine sedimentary rocks, which are incorporated as tectonic slices in the Bantimala Complex (Sukamto and Westermann, 1992), represent remnants of the supracrustal sedimentary cover of this fragment (Wakita et al., 1996). Termination of underflow by collision and buoyancy of the microcontinent facilitated rapid uplift of the subducted material during the early Late Cretaceous (Parkinson et al., 1998), which was eroded before and during the deposition of upper Albian to lower Cenomanian chert and siliceous shale (Wakita et al., 1996). Turbiditic deposits (Balangbaru and Marada Formations) were laid down during the Campanian–Maastrichtian (80–65 Ma), as suggested by planktonic foraminifera (van Leeuwen, 1981; Sukamto, 1985) and nannofossil evidence (Hasibuan, 1996). These formations, and the lithologically similar Latimojong Formations exposed in the central and northern parts of Western Sulawesi, contain abundant subangular quartz, fresh plagioclase and volcanic fragments, suggesting a magmatic arc provenance. Volcanic intercalations are a minor constituent. Most authors (e.g. van Leeuwen, 1981; Sukamto and Simandjuntak, 1983; Parkinson et al., 1998) interpret the Upper Cretaceous sequences to have been deposited in a deep marine fore-arc basin, situated to the west of a northwest-dipping subduction zone, with (part of) the related magmatic arc located in SE Kalimantan, where Upper Cretaceous volcaniclastics are intruded by dykes and stocks with K/Ar ages from 87 ± 4 Ma to 72 ± 4 Ma (Yuwono et al., 1988b). During the Paleocene and Early Eocene Kalimantan and Western Sulawesi were largely emergent (Wilson and Moss, 1999). The only rocks identified from this period in SW Sulawesi are calc-alkaline volcanics at Bantimala (Alla/Bua Volcanics) and possibly at Biru (lower part Langi Volcanics), thought by some authors

to be related to subduction that had shifted again southeastward at the beginning of the Tertiary (Yuwono et al., 1988a; van Leeuwen, 1981). This interpretation is compatible with the widely held view that subduction at the Sundaland margin was continuous through the Late Mesozoic into the Cenozoic (e.g. Yuwono et al., 1988a; Soeria-Atmadja et al., 1998; Sribudiyaning et al., 2003) with an Andean magmatic arc stretching from Sumatra to Western Sulawesi. A different scenerio is proposed by Hall (2009), who in view of the near-absence of igneous rocks of Paleocene–Early Eocene age along the Sundaland margin suggests that at the beginning of the Cenozoic there was no northward subduction, with a passive margin being present all round southeastern Sundaland, except for minor subduction beneath, West Sulawesi and Sumba. Only from about 45 Ma did subduction (and volcanic activity) resume as a result of Australia moving rapidly northwards. There was widespread extension and formation of basins around the margins of Sundaland by Middle Eocene times (Van de Weerd and Armin, 1992; Wilson and Moss, 1999). In SW Sulawesi, terrestrial to marginal marine siliciclastics of the Malawa Formation were deposited (t’Hoen and Ziegler, 1917; Sukamto, 1982), possibly in a series of fault-bounded sub-basins. The siliciclastic material was probably derived from the Sundaland continental basement (Hamilton, 1979; Fraser and Ichram, 2000). The clastic supply was from time to time interrupted, allowing peat deposits to be formed, which are preserved in the stratigraphic record as thin coal beds (t’Hoen and Ziegler, 1917). The coal-bearing strata contain abundant pollen with Sundaland affinities (Morley, 1998). The age of the bottom part of the formation is poorly constrained, but is assumed by some authors to be Early Eocene (e.g. Harahap et al., 1999; Wilson and Moss, 1999). Pollen, nannofossil and foraminifera assemblages described from several localities higher up in the sequence (Crotty and Engelhardt, 1993; GRDC, 2001; Sukamto, 1985; Morley, 1998) indicate a Middle Eocene age. During the Middle Eocene a volcanic arc developed along the eastern margin of the present Western Divide Mountains. The Malawa basin(s) received occasional pyroclastic showers, as evidenced by the presence of tuffaceous material mixed with the siliciclastics (t’Hoen and Ziegler, 1917; Samodra and Purnamaningsih, 1993). On the other side of the arc large amounts of volcaniclastic material were deposited in a shallow marine basin (lower part Member A, Matajang Formation). During this time the Makassar Straits was the site of extension, block-faulting and subsidence, as can be seen on seismic lines across the straits (Guntoro, 1999; Nur Aini et al., 2005; Puspita et al., 2005). This caused the severing of the land connection between Kalimantan and Sulawesi (Wilson and Moss, 1999), which in turn ruptured the supply of siliciclastic material to SW Sulawesi (Fraser and Ichram, 2000). Carbonate development (Tonasa Formation), that had began locally during the Middle Eocene (Harahap et al., 1999), became widespread in SW Sulawesi during the Late Eocene forming a 100 km by 40–50 km platform area (Wilson and Bosence, 1996; Wilson et al., 2000). By this time the volcanic arc had shifted about 15 km to the east, producing lavas and pyroclastics (Member B, the Matajang Formation), which were deposited in a shallow marine basin and partly reworked (upper part Member A). Volcanogenic input mostly hindered carbonate development in this area, but during short periods of relative volcanic quiescence and/or a decrease in sediment supply, foraminiferal shoals formed. The near-absence of volcanic material in the platform carbonates in the Western Divide Mountains may be explained by assuming that the main development of the Tonasa Carbonate Platform occurred on faulted highs separated from the volcanic arc by a basinal area that acted as a trap for volcaniclastic debris (Wilson, 2000). The occurrence of scattered outcrops of Middle–Upper Eocene calc-alkaline volcanics throughout Western Sulawesi (van Leeu-

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wen and Muhardjo, 2005), the presence of Upper Eocene volcanic material at the bottom of the Kampung Baru I well in the East Sengkang basin (Yulianto, 2004), a 60 m intersection of Middle Eocene basalt in a well drilled offshore the south coast (Sudarmono, 2000), and an exposure of andesitic breccia of probable Eocene age on Tanahjampe island, about 150 km to the south of Western Sulawesi (Guntoro, 1999), suggest that volcanism took place over a strike distance of more than 1200 km. The relatively restricted dimensions of most of the exposures may indicate that the volcanic activity was not very intense and took place intermittently. A possible explanation for the apparent low intensity of the Eocene volcanic activity is that at the time subduction beneath Western Sulawesi was oblique and had a significant strike-slip component (e.g. Hall, 1996). By the end of the Eocene volcanic activity had ceased all together and SW Sulawesi became a passive margin, possibly as the results of increasingly oblique plate convergence, changing the active margin into a strike-slip transform margin (Polvé et al., 1997; van Leeuwen and Muhardjo, 2005) as Western Sulawesi rotated in a counterclockwise direction. Results of reconnaissance paleomagnetic investigations suggest a counterclockwise rotation of SW Sulawesi of 35–50° sometime during the period of 63–13 Ma (Haile, 1978; Sasajima et al., 1980). Shear portioning along wrench faults resulted in the formation of a marginal basin bordered to the west of Sundaland, while a new westerly-dipping subduction zone was created along its eastern side. Blueschists exposed further to the east in SE Sulawesi, comprising graphite–mica schists and MORB-like metabasites, which were subducted around 33 Ma ago and exhumed around 20 Ma (Helmers et al., 1989; Wijbrans et al., 1994) possibly represent the remnants of this subduction zone.Combined extension and subduction gave rise to the formation of volcanic rocks ranging from basalt and andesite with very weak to moderate subduction signatures (Kalamiseng volcanics) to more typical subduction-related arc volcanics (Deko 2 volcanics). The establishment of a new arc system facilitated transfer of heat (as well as mass) into the overriding plate, encouraging melting of this crust (Deko 1 volcanics). Further to the west along the continental margin extensional faulting resulted in segmentation of the platform causing localized drowning and subaerial exposure, with carbonates being redeposited in a region of deep water sedimentation together with marls in the northwestern part of SW Sulawesi (Wilson and Bosence, 1996; Wilson, 2000; Wilson et al., 2000). During the Oligocene the Bone Mountains area was occupied by a shallow continental shelf where benthic and planktonic foraminifera were fragmented and mixed with material derived from the carbonate platform (Karopa Formation). The Lower Miocene and bottom part of the Middle Miocene in SW Sulawesi are characterized by a variety of carbonate lithologies, indicating a range of shallow to deeper water environments (Wilson et al., 2000). In the deeper water areas e.g. Barru, Sengkang Basin and Bone Mountains (Member C, Baco Formation), marls and calcareous mudstones are interbedded with beds and lenses of coarse texturally immature material (Wilson et al., 2000; Yulianto, 2004; this issue). Around 13–14 Ma tectonic conditions became less stable. Widespread block-faulting took place in the Western Divide Mountains. Differential vertical movements in the order of 300–400 m are indicated at Biru where a basal Middle Miocene unit contains limestone clasts of increasingly older age (from Early Miocene to Late Eocene) going up the sequence (van Leeuwen, unpubl. data). Eroded material was deposited both close to fault-blocks (e.g. Biru) and transported over a considerable distance as turbidity currents and debris flows (upper part Member C, Baco Formation). Around the same time a phase of potassic magmatic activity began, which may have been triggered by deep-seated faults tapping into mantle metasomatized by (an) earlier subduction event(s) (van Leeuwen,


1981; Letellier et al., 1990). The potassic magmatism commenced throughout Western Sulawesi around the same time (Bergman et al., 1996; Polvé et al., 1997; Elburg et al., 2002, 2003), suggesting that during the Middle Miocene the entire region was affected by an extensional tectonic regime (Macpherson and Hall, 1999; van Leeuwen and Muhardjo, 2005). In some areas in SW Sulawesi carbonate accumulation continued more or less undisturbed, in others volcanic material was mixed with carbonates (van Leeuwen, 1981; Samodra and Purnamaningsih, 1993; Wilson, 2000), or was the dominant rock type (Member D, Baco Formation). The tectonic event culminated around 13–12 Ma with the juxtaposition of the Bone Group segment against the continental margin of Sundaland composed of the Salokalupang Group, along the Walanae Fault Zone. The Salokalupang rocks were tilted, sliced into tectonic lenses of variable size, and locally folded. In the Western Divide Mountains region, some fault blocks were tilted and/or folded and eroded, whereas others were subject to only vertical movements, resulting in both conformable and unconformable upper contacts (van Leeuwen, 1981; Wilson and Bosence, 1996). Similar tectonic features have been observed further to the north (Harahap et al. 1999) and in the Sengkang region (Ascaria et al., 1997). The Bone Basin, which is clearly bounded by major northerly trending marginal faults (Yulianto, 2004), may have started to form around this time as the result of extension (Hamilton, 1979; Polvé et al., 1997). The main cause of the Middle Miocene tectonic event was very rapid rollback of a subduction hinge further to the east in the Banda region, which induced massive extension of the region above the subducted slab, including the western part of Sulawesi (Hall, 2009). Soon after the emplacement of the ophiolite segment(s) potassic volcanic activity increased significantly. Volcanics and their erosion products (Camba Formation) and related units were deposited predominantly in a shallow marine environment (Sukamto, 1982; Sukamto and Supriatna, 1982), but there is also evidence for local subaerial exposure (van Leeuwen, 1981). The large influx of volcaniclastic debris inhibited carbonate production with the exception of the area around the northern edge of the Bone Mountains, where reef knolls formed during the late Middle–Late Miocene (Tacipi Formation, Grainge and Davies, 1985; Ascaria et al., 1997). The Walanae Depression may have started to develop during the Late Miocene. It was filled with sediments, pyroclastics, and volcaniclastics (Walanae Formation), initially in a marine environment, but becoming marginally marine towards the end of the Pliocene (t’Hoen and Ziegler, 1917). Prior to this the area probably experienced uplift, as suggested by the fact that along the western margin of the depression the Walanae Formation directly overlies Cretaceous basement (van Leeuwen, 1981; Wilson, 1995). During the Pliocene another major tectonic event took place in Western Sulawesi, which is widely attributed to the collision of the Banggai-Sula continental fragment with the east arm of Sulawesi, although some authors have suggested that it started somewhat earlier (e.g. Hamilton, 1979; Garrard et al., 1988; Polvé et al., 1997). R. Hall (pers comm., 2009) has pointed out that the Banggai-Sula fragment, in view of its small size, is unlikely to have been the main cause of this major orogeny, which in the central and northern parts of West Sulawesi produced fold and thrust belts, voluminous granitoid magmatism and rapid uplift (e.g. Bergman et al., 1996; van Leeuwen and Muhardjo, 2005). He suggests that as SE Asia is a globally unusual region of weak lithosphere and crust, lower crust may be flowing away from sites of sedimention (e.g. the major Cenozoic basins flanking Borneo) and towards and beneath elevated regions (e.g. west Sulawesi and NW Borneo). The impact of this tectonic event in SW Sulawesi has been largely limited to folding of the Neogene sediments of the Sengkang Basin into a series of north-trending anticlines (van Bemmelen, 1949) and a general uplift of the SW Sulawesi region (van Leeuwen,


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1981). Wrench and normal faulting occurred along the Walanae Fault Zone (van Leeuwen, 1981; Grainge and Davies, 1985; Berry and Grady, 1987). At Bantimala, the Middle Cretaceous melange complex was thrust over Eocene–Miocene formations (Sukamto, 1985). Folding and faulting in adjacent Tonasa rocks probably took place during the same time (Wilson, 1995). This area was also subject to significant strike-slip faulting dominantly trending NNW (Berry and Grady, 1987). However, the bulk of the Neogene strata in SW Sulawesi are flat lying and show little sign of significant deformation (t’Hoen and Ziegler, 1917; Sukamto, 1982). The difference in styles of Pliocene deformation between SW Sulawesi and the region to the north is similarly reflected in the structural setting of the adjacent Gulf of Bone; in the northern part folded Middle Miocene–lowermost Pliocene rocks are unconformably overlain by Pliocene orogenic sediments, whereas further to the south the young Cenozoic succession is largely undisturbed (Sudarmono, 2000).

8. Conclusions Our field and laboratory data provide new insights into the evolution of the Bone Mountains and Southwest Sulawesi as a whole. The Salokalupang Group spans a wider period than previously thought: Middle Eocene–Miocene, with apparent gaps in the Early Oligocene and Early Miocene. It is made up of three distinct units: Matajang Formation (Middle–Upper Eocene), Karopa Formation (Oligocene) and Baco Formation (Early–Middle Miocene). Observed/inferred contacts between these units are invariably tectonic in nature, reflecting the fact that the rocks are exposed within a major strike-slip fault zone (East Walanae fault system), which runs parallel with the western margin of the Bone Mountains. The units can be correlated with similar successions in the Western Divide Mountains, located further to the west, indicating that the Salokalupang Group forms part of the Sundaland continental margin. They record a history of Eocene calc-alkaline arc volcanism with deposition of siliciclastic sediments and platform carbonate development behind the arc, followed by cessation of subduction and continuing formation of carbonates and associated sediments during the Oligocene–Early Miocene, and finally in the early Middle Miocene widespread block-faulting, the onset of potassic volcanism, and sedimentation characterized by turbidity currents and debris flows related to the initial stage of a transtensional tectonic event. The Bone Group differs significantly from the Salokalupang Group in terms of lithologies and environment of formation. The bulk of the rocks consists of Oligocene–Lower Miocene MORB-like basaltic-andesitic volcanics, showing weak to moderate subduction signatures (Kalamiseng Formation). The unit is associated with a series of interbedded hemipelagic mudstones and intercalated volcanics, which in part have typical subduction-related arc volcanic features (Deko Formation). The majority of the Bone Group volcanics show geochemical characteristics that suggest that they may have formed during the early stages of subduction. This combined with the observation that the continental margin appeared to have changed from an active to a passive transform margin by the end of the Eocene, leads to the hypothesis that during the Oligocene a transtensional marginal basin formed adjacent to the passive margin, with west-directed subduction being initiated along its eastern border. Melting of a basaltic precursor in the overriding crust may explain the rather unusual geochemical composition of some of the Deko volcanics. The Salokalupang and Baco Groups are at present juxtaposed along the East Walanae fault system. Geological relationships and paleontological dating constrain the timing of the juxtaposition and the deformation of the latter into a series of variably sized tec-

tonic slices to around 13–12 Ma, probably about 1 Ma after the extensional and potassic volcanism event commenced in SW Sulawesi. The causes of the major tectonic event that took place in Western Sulawesi during the Pliocene are still poorly understood. It had its greatest impact in the central and northern parts of the region. Stress was transmitted southward through the basement where it caused localized deformation along major zones of crustal weakness, i.e. the Middle Cretaceous Bantimala suture and Walanae Fault Zone. The two ophiolite segments exposed along the eastern margin of the south arm of Sulawesi, i.e. Lamasi Complex and Bone Group segment, may share, in part, a similar history of formation and emplacement. More detailed work is required in order to elucidate more precisely their age(s) and environment(s) of formation, and subsequent tectonic history. This may shed important new light on the Cenozoic development of the eastern part of Sulawesi, which is still much less known than that of the western part. Acknowledgments We thank Tony Barber for helpful comments on earlier versions of the paper. We also thank Tuti Mariani for preparing the manuscript, and Heru Pratomo for creating the figures. Special thanks are due to Colin Macpherson, who helped us greatly with the interpretation of the geochemical data, and to Robert Hall, who provided invaluable input throughout the writing of the paper and reviewed the final draft. Appendix A. Appendix 1 Taxonomic notes A.1. Matajang Formation, Member A SLD17, 19, 20, 23, 24, S. Deko; mudstone intercalations; abundant planktonic forams, including Acarinina bulbrooki, Globigerina medizzai, Globigerinatheka mexicana, Globorotalia cerroazulensis pomeroli, Globorotaloides carcoselleensis, Morozovella spinulosa, and Truncorotaloides topilensis; zones P13–P14 (late Middle Eocene). SL 29, S. Deko; limestone; abundant planktonic forams including Globigerina corpulenta, Gl venezuelana, Gl tripartita, Globorotalia cerroazulensis cocoensis, Gbt nana, Globigerinatheka mexicana, and Hantkenina alabamensis; top P14–bottom P15 (late Middle Eocene–early Late Eocene). B 1047, S. Matajang; claystone; abundant nannofossils including Cyclicargolithus floridanus, Cribrocentrum reticulatum, Discoaster barbadiensis, D. deflandrei, D. saipanensis, Markalius inversus, Reticulofenestra dictyoda, R. umbilicus and Helicosphaera seminulum; CP14 (Middle Eocene–early Late Eocene). B 920, S. Deko; mudstone; abundant planktonic forams, including Globigerina senni, Gl. eocaena, Gl. inaequispira, Gl. corpulenta, Morozovella spinulosa and Globoratalia cerroazulensis; P14 (late Middle Eocene). SKG 25, S. Katumpang Kiri; mudstone; very rich in planktonic forams including Globigerina ampliapertura, G. prasaepis, Globorotalia increbescens; P17–P20 (Late Eocene–Early Oligocene). B 931A, S. Deko; mudstone; Globigerina eocaena, Gl. ampliapertura; P17 (Late Eocene). B 955, 1022, 1026, S. Mate; B1055, S. Matajang; B 969, 1084, S. Katumpang Kiri; B 1005, B 1010, 1035, 1078, S. Linrung; B 315, 407, S. Deko; B 925 S. Deko Kanan; B 905, S. Kalupang; exotic block; B 979, S. Baco; exoctic block; all bioclastic limestone samples with the exception of B 1005 (lime mud) containing large benthonic forams including Asterocyclina penuria, A. matazensis, A. insuricamerata, Discocyclina omphala, Nummulites bagelensis, N. pengaronensis, N. semiglobula, Heterostegina saipanensis, H. aequatorial, Spiroclypeus

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vermicularis, Operculina complanata and in some samples Pellatispira provalae; Tb (Late Eocene). B 999, S. Linrung; sandy limestone with Reticulofenestra hampdenensis and abundant Cyclicargolithus floridanus suggesting a Late Eocene age. SLD 115, 116, 121, S. Deko; limestone; Globigerina officinalis, G. gortanii, G. yeguanensis, G. tripartita, Globigerinatheka mexicana, Gbk semiinvoluta, Globorotalia cerroazulensis cerroazulensis, Gb. Increbescens, Pseudohastigerina micra, Morozovella lehneri, M. spinulosa; P14–15.

SLD 4, 4A, 33, 34, 41a, 58, 118B, 120D; S. Deko; mudstones; abundant nannofossils showing poor to moderate preservation, including Discoaster variabilis, D. brouweri, D. deflandrei, D. challengeri, D. exilis, D. formosus, Spenolithus compactus and Reticulofenestra pseudoumbilica; based on the Discoaster chronostratigraphy established by Lambert and Laporte-Galaa (2005) for the Kutai Basin in East Kalimantan this assemblage suggests zone CN5a (early Middle Miocene).

A.2. Matajang Formation, Member B

SLK 05, 16, 17, 31, S. Kalumpang; mudstone; common Orbitulina species, including O. universa and O. suturali, indicating unit is not older than basal Middle Miocene (N9); presence of Globoratalia praemenardii in SLK 31 narrows age down to N10–N12 (Middle Miocene). B 979, 984, S. Baco; claystone; abundant nannofossils including Coccolithus miopelagicus, Cyclicargolithus floridanus, Discoaster adamanteus, D. formosus, D. exilis, D. deflandrei, D. variables, Dictyococcites antarticus, D. productus, Reticulofenestra pseudoumbilica, Sphenolithus heteromorphus; CN4–CN5a (Middle Miocene).

B 6, S. Matajang; limestone; Nummulites javanus; top Ta3-bottom Tb. B 8, 9, 13, 14, S. Matajang; limestone; abundant Pellatispira species including P. provalae, Discocyclina spp., Nummulites spp., Biarritzina sp., Operculna venosa; Tb (upper part). A.3. Karopa Formation B 949, S. Deko; mudstone; planktonic forams include Globigerina praebulloides, Gl. venezuelana, Gl. ciperoensis ciperoensis (common); P20–22 (Oligocene). B 1096, S. Deko; globigerina ooze; preservation generally poor; assemblage include Globigerina triparti, Gl ciperoensis, Gl. gortanii, Gl. ampliapertura, G. sellii, G. oauchitaensis, suggesting zones P20– 21 (Middle Oligocene). B 311, 313, 947, S. Deko; B 1048, S. Matajang; mudstone; calcareous nannofossils including Dictyococcites bisectus, Cyclicargolithus abisectus, C. floridanus; CP18–19 (Middle–Late Oligocene). B1048, S. Matajang; claystone; Discoaster deflandrei, D. tanii ornatus, Cyclicargolithus abisectus, C. floridanus; CP18 (Early Oligocene). B 1076, S. Linrung; grey mudstone; nannofossils showing poor preservation; assemblage includes Coccolithus pelagicus, Cyclicargolithus, floridanus, Dictyococcites productus, Sphenolithus moriformis. This assemblage suggests an Oligocene age. B 1080, S. Linrung; grey mudstone, nannofossils include Cyclicargolithus abisectus, C. floridanus and Dictyococcites scrippae, suggesting zones CP 18–19 (Middle–Late Oligocene). A.4. Baco Formation, Member C B 1003, 1012, 1021, S. Linrung; B1050, 1053, S. Matajang; B 1127, S. Katumpang Kanan; calcareous mudstone; common nannofossils, including Discoaster adamanteus, D. deflandrei, D. aulakos, D. formosus, D. exilis, Helicosphaera intermedia, H. ampliaperta, Sphenolithus belemnos and S. heteromorphus; CN3–4 (late Early Miocene). B 1050 A, S. Matajang; mudstone; abundant, but poorly preserved nannofossils, including Sphenolithus belemnos which is indicate of zones C2–3. SLD 1, 6, S. Deko; mudstone; abundant planktonic forams including Globorotalia praemenardii, Gt. siakensis, Globigeniroides subquadratus, G. immaturus, Hastigerina siphonifera, and Orbulina universa, suggesting zones N10–12 (early Middle Miocene). SLD 32, S. Deko; mudstone; planktonic foram assemblage includes Globoratolia peripheronda, Gt fohsi, Gt. siakensis, Globigerinoides subquadratus, Gg. succulifir, Globaquadrina altisphira, Orbulina universa, Sphaeroidinellopsis seminulina, suggesting zones N10– lower N11 (early Middle Miocene). SLD 36, 37, 38, 39, 43, 50, 51; S. Deko, mudstones; planktonic forams include Globorotalia praemenardii, Gb. lobata, Gb. siakensi, Gb. menardii, Globoquadrina druyi Globigerinoides subquadratus, Gs. trilobus, Gs. immaturus, Orbulina universa, Sphaeordinellopsis seminula and Hastigerina siphonifera; N11–12 (early Middle Miocene).

A.5. Baco Formation, Member D

A.6. Kalamiseng Formation B 1061, upper Matajang; red mudstone; abundant nannofossils, including Cyclicargolithus abisectus, C. floridanus, Dictyococcites bisectus and Sphenolithus pseudoradians, suggesting zones CP18– 19a (Oligocene). B 928, S. Deko Kanan; red mudstone, rich in nannofossils, including Discoaster deflandrei, D. adamanteus, Cyolicargolithus floridanus, Helicosphaera bramlettei, H. obliqua, H. euphratis, Dictyococcites bisecta, and Sphenolithus ciperoensis, indicating zone CP 19 (Late Oligocene). A.7. Deko Formation B 319, 323, 408, 413 A, B, C, D, 426, S. Deko; mudstone containing abundant nannofossils, including common Dictyococcites bisectus, Discoaster deflandrei, Cyclicargolithus abisectus and C. floridanus; rarer are Dictyococcites scrippsae, Sphenolithus predistensus, S. pseudoradians, Helicosphaea obliqua, H. compacta and Markalius inversus;suggesting zones CP18–CP19b (Oligocene). SL 101A, 104, 105, 109B, B1096, S. Deko; mudstones containing abundant planktonic foraminifera including Catapsydrax dissimilis, C. unicavus, Globergerina ampliapertura, Gl. tripartita, Gl. ciperoensis, Gl. gortanii, Gl. sellii, Gl. venezuelana, and Gl. oauchitaensis; suggesting zones P19–P21 (Oligocene). B 1054, S. Matajang; mudstone (globigerina ooze); Globigerina tripartita, G. yeguaensis, G. ampliapertura, G. ciperoensis, G. gortanii, G. venezuelana; P20 (Early Oligocene). Appendix B. Appendix 2 Rock types Member A, Matajang Formation

(i) Volcaniclastic sandstones: grey, greenish or brown; fine to very coarse grained (dominantly fine to medium), locally pebbly; clast to matrix supported; texturally immature (dominantly angular clasts; poorly sorted) and poorly to well cemented; clast assemblages consist predominantly of igneous material (in descending order of abundance: lava of intermediate to basic composition, plagioclase, magnetite, quartz, K-feldspar, granite), varying amounts of carbonate material, minor quartz sandstone and chert; locally coal


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fragments; rocks are commonly calcareous as a result of deposition of calcite in open space or recrystallization of calcareous mud matrix. Bed thicknesses vary from 1 cm to 7.5 m; bed contacts usually sharp; soft sediment deformation, parallel lamination and graded bedding relatively common; large scale trough cross stratification locally developed. (ii) Mudstones: massive to shaley; mostly greyish, but reddish brown and black varieties also occur; partly calcareous; locally silty to sandy. (iii) Muddy breccias: poorly sorted and matrix supported; matrix constitutes 50–90% of the rock volume and is of similar composition and colour as interbedded mudstones; clast angular to subrounded; show a complete size spectrum ranging from granules to boulders up to 7 m in diameter; bed fabric random clast lithologies include mudstone (dominant) volcaniclastic sandstone, limestone, breccia, conglomerate and various chloritized volcanics; beds typically <5 m thick, but locally up to 30 m. (iv) Breccias/conglomerates: poorly sorted; matrix or clasts supported; angular to subrounded fragments (<2 cm to 3 m in diameter) consisting of two or more of the following rocks types; lava, volcaniclastic sandstone, volcanic breccia, calcareous mudstone, and bioclastic limestone.

Appendix C. Appendix 3 Sedimentary facies Member C, Baco Formation

(i) Claystone facies: grey to brownish grey mudstones, commonly calcareous and locally silty; partly well bedded, in partly massive; single mudstone intervals 5 cm to 25 m thick; no evidence of bioturbation was observed; planktonic foraminifera on the other hand are relatively abundant. (ii) Thin bedded sandstone facies: composed of sandstone, 2– 50 cm thick (average 10 cm), with thin mudstone intercalations; lower bed contacts sharp and locally erosive; fining upwards is a common feature of the sandstone beds; the thin bedded sandstone facies shows incomplete Bouma sequences (truncated, base–cutout, or a combination of the two). (iii) Graded sandstone facies: characterized by well bedded sandstones showing graded bedding; can be divided into two subfacies, A and B, which have beds that are <1 m and >1 m thick respectively; subfacies A sandstone beds are medium, coarse or very coarse grained in the lower part of an individual bed and locally pebbly; subfacies B sandstones are coarse to very coarse grained and generally pebbly; both subfacies fining upwards of the sandstone beds is a common feature with the uppermost part of the bed becoming siltstone (maximum 15 cm); this upper siltstone layer is not always present, maybe because it was not formed or was eroded by the overlying bed; parallel lamination commonly developed, in particular in the fine-grained portions. The sandstones in both the thin bedded and graded facies are composed of volcanic particles and carbonate. The latter occurs as fragments, matrix and cement. In a few cases the carbonate content exceeds 50 volume%. (iv) Bouldery mudstone facies: represented by chaotic, predominantly matrix supported breccias; matrix consists of clayey to subordinate sandy material; breccias thickly bedded to massive; breccia intervals vary from 60 cm to 30 m in thickness; clast lithologies include mudstone, andesitic-basalt lava, volcanic breccia, carbonate rocks, and volcaniclastic sandstone; some beds volcanic components dominate, in

others mudstone is the dominant clast phase, in yet others carbonate fragments (including bioclastic limestone and nummulitic limestone) are most common; breccia beds are poorly sorted, with a complete spectrum of clast sizes ranging from granulite to boulder (up to 2.5 m in diameter) in individual beds; finer grained breccias (fragments < 2 cm) locally present; larger fragments are subangular to angular, whereas smaller fragments tend to be subrounded to subangular; some clasts aligned parallel to the bedding but more commonly they show random orientations; bed contacts sharp and in places erosive.

Appendix D. Appendix 4 Rock types Member C, Baco Formation exposed in S. Kalupang Sandstones are the dominant rock type; typically show spheroidal weathering patterns, very hard when fresh; fine to very coarse grained, poorly to moderately well sorted; both texturally and compositionally immature with the texture being epiclastic, fragments angular to subangular consisting of variably altered andesite-basalt, pyroxene and feldspar, and subordinate hornblende, olivine, biotite, and opaque minerals; cement mineralogies include secondary quartz, zeolite, chlorite, and iron-oxides. Bed thicknesses vary from 10 cm to 10 m; thicker beds commonly pebbly, with the volume of clasts varying from rare to common; increase in average bed thickness upwards the sequence accompanied by a decrease of fine-grained interbeds, bed contacts difficult to recognize, but where observed beds possess erosive bases; vertical amalgamation of beds typical feature, with only rare intervals of fine-grained sediment preserved in beds up to 5 cm thick, consisting of mudstone and siltstone; rare wave ripples in siltstone; locally water escape and soft sediment deformation structures, including convolute bedding; other internal structures include vague bedding formed by parallel aligned clasts and alternating finer and coarse layers, normal graded bedding, and parallel lamination; latter particularly common where sandstone beds alternate with thin mudstone beds, which are either massive or also show parallel lamination. Interbedded volcanic breccias and subordinate conglomerates are massive, or more rarely, show internal stratification; sharp erosive-based contacts with underlying sandstone beds; breccias clast to matrix supported, poorly sorted, either monomict or polymict; clasts consist of andesitic to basaltic rocks and minor limestone, generally <40 cm in diameter; interbedded breccia and conglomerate in lower part of sequence as thin beds and lenses, increasing in thickness (up to 30 m) and frequency upwards; mudstone and siltstone intercalations mostly thin. References Adams, C.G., 1970. A reconsideration of the East Indian Letter Classification of the Tertiary, British Museum of Natural History. Geology 19 (3), 87–137. Ascaria, N.A., Harbury, N.A., Wilson, M.E.J., 1997. Hydrocarbon potential and development of Miocene knoll-reefs, South Sulawesi. In: Proceedings of the Petroleum Systems of SE Asia and Australasia Conference, May 1997, pp. 569– 584. Barber, A.J., Simandjuntak, T.O., 1996. Report on Geological Fieldwork in South Sulawesi. Southeast Asia Research Group, University of London, Report No. 154, p. 38. Ben-Avraham, Z., Nur, A., Jones, D., 1982. The emplacement of ophiolites by collision. Journal of Geophysical Research 87, 3861–3867. Bergman, S.C., Coffield, D.Q., Talbot, J.P., Garrard, R.J., 1996. Tertiary tectonic and magmatic evolution of Western Sulawesi and the Makassar Strait, Indonesia: evidence for a Miocene continent-continent collision. In: Hall, R., Blundell, D. (Eds.), Tectonic Evolution of Southeast Asia. Geological Society Special Publication 106, pp. 365–389. Berry, R.F., Grady, A.E., 1987. Mesoscopic structures produced by Plio-Pleistocene wrench faulting in South Sulawesi, Indonesia. Journal of Structural Geology 9, 563–571.

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