basement character and basin formation in gorontalo bay, sulawesi, indonesia

Basement character and basin formation in Gorontalo Bay, Sulawesi,

Indonesia: new observations from the Togian Islands

M. A. COTTAM1*, R. HALL1, M. A. FORSTER2 & M. K. BOUDAGHER-FADEL3

1SE Asia Research Group, Department of Earth Sciences, Royal Holloway University of London,

Egham, Surrey, UK2Department of Earth Sciences, Australian National University, Canberra, ACT, 0200, Australia

3Department of Earth Sciences, University College London, London, UK

*Corresponding author (e-mail: [email protected])

Abstract: We present a new stratigraphy for the Togian Islands, Sulawesi, and interpret the age,character and evolution of Gorontalo Bay. At its western end the bay is underlain by continentalcrust. The central part is underlain by Eocene to Miocene oceanic and arc rocks, although thearea south of the Togian Islands could have continental crust of the Banggai-Sula microcontinentthrust beneath this and the East Arm ophiolite. Gorontalo Bay was not a significant deep bathy-metric feature before the Miocene. Field relationships indicate a latest Miocene to Pliocene agefor inception of the basin. Medium-K to shoshonitic volcanism in the Togian Islands is not dueto subduction but reflects crustal thinning and extension in the Pliocene and Pleistocene, causingthe underlying mantle to rise, decompress and melt. Extension is continuing today and is probablythe cause of volcanism at Una-Una. Volcanic activity migrated west with time and volcanic pro-ducts have been offset by dextral strike-slip displacement along the Balantak Fault. Extension andsubsidence was driven by rollback of the subduction hinge at the North Sulawesi Trench with apossible contribution due to flow of the lower crust.

Gorontalo Bay is one of the most enigmatic basinsin East Indonesia. It is relatively deep with waterdepths up to 2000 m, and Hamilton (1979) showedup to five kilometres of sediment in its westerndepocentre. It is surrounded by land on three sidesand receives large volumes of sediment fromnearby mountains up to three kilometres high.Miocene carbonates are widespread in these areas(van Leeuwen & Muhardjo 2005) and suggest thatthe deep basin formed since their deposition butthe timing and mechanism of basin inceptionremain unclear.

The nature and age of the crust beneathGorontalo Bay is also unknown. To the north, theNorth Arm of Sulawesi is interpreted as a volcanicarc built on Eocene oceanic crust (Taylor & vanLeeuwen 1980; Elburg et al. 2003; van Leeuwen& Muhardjo 2005). In contrast, at the western endof Gorontalo Bay, there are two kilometre highmountains with young metamorphic ages and evi-dence of continental crust, Miocene extension andcore complex formation (Sukamto 1973; Elburget al. 2003; van Leeuwen et al. 2007). To the south,the East Arm of Sulawesi comprises ophiolitic rocksof the East Sulawesi Ophiolite (Simandjuntak 1986;Monnier et al. 1995; Bergman et al. 1996; Parkinson1998; Kadarusman et al. 2004). Silver et al. (1983b)suggested that Gorontalo Bay was a fore-arc basin,

underlain by ophiolitic crust equivalent to the EastArm ophiolite, situated in front of the North Armvolcanic arc that has been thrust south onto theBanggai-Sula microcontinent.

The Togian Islands, situated in the centre ofGorontalo Bay (Fig. 1), offer a unique opportunityto investigate aspects of the basin’s origin and evol-ution. The archipelago forms a broadly WSW–ENEtrending ridge that continues to the west as a sub-marine feature. Geological maps of the islandsshow igneous rocks and contrasting interpretationsof them. Kundig (1956) reported andesitic intrusiverocks in the central islands, and older ophioliticrocks in the eastern islands – suggesting a possiblelink to the East Sulawesi Ophiolite. In contrast,Rusmana et al. (1982) reported widespread tuffsand sedimentary formations of Mio-Pliocene age.The volcanic rocks could therefore be part of theophiolite, could form part of the North Arm volcanicarc, or could be subduction-related products thatpredate the collision (Garrard et al. 1988; Davies1990) of the Banggai-Sula microcontinent with theEast Arm.

The Togian Islands are also close to the isolatedactive volcano of Una-Una, just NW of the Togianarchipelago, which has a K-rich chemistry anderupted violently in 1983 (Katili et al. 1963; Katili& Sudradjat 1984). It is not a typical subduction

From: Hall, R., Cottam, M. A. & Wilson, M. E. J. (eds) The SE Asian Gateway: History and Tectonicsof the Australia–Asia Collision. Geological Society, London, Special Publications, 355, 177–202.DOI: 10.1144/SP355.9 0305-8719/11/$15.00 # The Geological Society of London 2011.

volcano in position (about 200 km above theBenioff zone) and, if related to this subduction, isunusual in being the only volcano.

We present a stratigraphy for the Togian Islandsbased on new field observations and dating. In manycases dating was restricted by the intense tropicalweathering typical of SE Asia, and/or a lack ofdatable material. We combine these new data withearlier studies and observations of the physio-graphy, bathymetry and seismicity of the northernSulawesi region, to elucidate the Cenozoic historyof Gorontalo Bay.

Tectonic setting

Sulawesi comprises a complex association of mag-matic arcs, metamorphic rocks (varying in gradefrom low to high), ophiolites and microcontinentalfragments that have been variously assembled anddeformed during the Late Mesozoic and Cenozoic(e.g. Audley-Charles 1974; Hamilton 1979; Hall2002). It has been subdivided into four tectonostra-tigraphic terranes separated by major faults (e.g.Hamilton 1979). The composition of the terranessurrounding the study area is described below.Following recent studies (e.g. Calvert 2000; vanLeeuwen & Muhardjo 2005; van Leeuwen et al.2007) we do not use the term Western SulawesiPlutono-Volcanic Arc Terrane. Instead, we follow

previous authors in separating this ‘terrane’ in totwo different entities based on the recognition ofsignificant differences in age and character ofrocks (e.g. Taylor & van Leeuwen 1980; Calvert2000; Elburg et al. 2003). We adopt the termsWestern Sulawesi Province and Northern SulawesiProvince (e.g. van Leeuwen et al. 2007; seeFig. 1). The position of the boundary between theseprovinces remains uncertain (Elburg et al. 2003).

Western Sulawesi Province

The Western Sulawesi Province (Fig. 1) represents acontinental margin segment (van Leeuwen et al.2007). It has a metamorphic basement that includesthe Malino and Palu Metamorphic Complexes,exposed at the NW and SW corners of GorontaloBay respectively (Elburg et al. 2003; van Leeuwenet al. 2007). These rocks form part of an arcuatezone of dismembered accretionary complexes(Parkinson 1998) and continental fragments, meta-morphosed in the mid-Cretaceous during emplace-ment along the SE margin of Sundaland by NWdirected subduction (Parkinson 1998).

The basement is overlain by weakly metamor-phosed Upper Cretaceous sedimentary rocks of theLatimojong Formation, which are in turn overlainby a sequence of weakly metamorphosed Palaeo-gene sedimentary rocks and subordinate volcanicrocks belonging to the ‘Older Series’ of Elburg

Fig. 1. Tectonostratigraphic provinces of Sulawesi. Modified after Hall & Wilson (2000), Calvert (2000) andvan Leeuwen & Muhardjo (2005).

M. A. COTTAM ET AL.178

et al. (2003). The exact nature of the contact(depositional or faulted) is not known. Close to thestudy area these rocks include the Tinombo For-mation (Brouwer et al. 1947), fore-arc basin sedi-ments characterized by a transition from syn-riftsedimentation to platform carbonates and deepermarine sedimentation between the Late Eoceneand Middle Miocene (Coffield et al. 1993; Wilson& Bosence 1996; Calvert 2000). The contempora-neous Tinombo Formation volcanic rocks (c. 51 to17 Ma) range from basalt to rhyolite and includedykes, volcanic piles and co-magmatic intrusivestocks (Elburg et al. 2003).

Intrusive and extrusive rocks of the ‘YoungerSeries’ (Elburg et al. 2003) include an acidichigh-K calc-alkaline (CAK) suite of plutons(Kavalieris et al. 1992) and comagmatic volcanicrocks (van Leeuwen et al. 1994; Elburg et al.2003), and a high-K calc-alkaline, shoshonitic andultra-potassic alkaline (HK) suite of dykes, smallstocks and less common extrusive rocks (Elburget al. 2003).

Northern Sulawesi Province

The Northern Sulawesi Province (Fig. 1) comprisesa dominantly tholeiitic Tertiary volcanic arc builton Eocene oceanic crust (Taylor & van Leeuwen1980; Elburg et al. 2003; van Leeuwen & Muhardjo2005). Volcanism was driven by the northward sub-duction of Indian Ocean lithosphere beneath theNorth Arm (e.g. Hall 1996, 2002; Rangin et al.1997). The Papayato Volcanic rocks are the pro-ducts of this arc, a bimodal suite of mafic andfelsic volcanic rocks cut by co-magmatic stocks ofgabbro and diorite (Trail et al. 1974; Kavalieriset al. 1992; van Leeuwen et al. 1994; Elburg et al.2003) belonging to the ‘Older Series’ of Elburget al. (2003). Limited isotopic and palaeontologicalages suggest a Middle Eocene to earliest Mioceneage (van Leeuwen et al. 2007) making them thebroad age equivalent of the Tinombo Formation inthe Western Sulawesi Province. However, contrast-ing volcanic–sedimentary proportions suggest thatthey were formed in different tectonic environments(van Leeuwen et al. 2007).

The Papayato Volcanic rocks are overlain by athick series of Neogene volcanic rocks and volcani-clastics of calc-alkaline composition and cut byco-magmatic intrusives (‘CA Suite’ of Polve et al.1997), which are accompanied by marine sedimen-tary rocks (Kavalieris et al. 1992) that include well-bedded shallow marine sediments and limestones ofEarly to Middle Miocene age (e.g. Sukamto 1973;Norvick & Pile 1976; Ratman 1976). All theserocks are cut by Neogene volcanic rocks belongingto the ‘Younger Series’ (Elburg et al. 2003). Theyinclude andesitic and dacitic stocks, dykes and

epiclastic rocks of the calc-alkaline ‘CA Suite’,and associated Early–Middle Miocene marine sedi-ments (Elburg et al. 2003).

East Sulawesi Ophiolite

The East Sulawesi Ophiolite (Fig. 1) comprises asequence of dunite, lherzolites and harzburgites,ultramafic cumulates, layered gabbros, isotropicgabbros, sheeted dykes and basaltic pillows andlavas (e.g. Simandjuntak 1986; Parkinson 1991,1998). Field mapping (Kadarusman et al. 2004)and geophysical studies (Silver et al. 1978) sug-gest an abnormally large reconstructed strati-graphic thickness of at least 15 km. The origin ofthe East Sulawesi Ophiolite has been variouslyattributed to a typical mid-oceanic ridge (e.g.Soeria-Atmadja et al. 1974; Simandjuntak 1986),supra-subduction zone (Monnier et al. 1995;Bergman et al. 1996; Parkinson 1998) and oceanicplateau settings (Kadarusman et al. 2004). K–Ardating of the ophiolite ranges in age from Cretac-eous to Eocene (Simandjuntak 1986). They areinterpreted to reflect Cretaceous, specificallyCenomanian, ocean floor with younger seamounts(Simandjuntak 1986). K–Ar dating (Parkinson1998) has been interpreted to suggest intra-oceanicthrusting of the ophiolite at c. 30 Ma.

Microcontinental fragments

The Banggai-Sula block (Fig. 1) has a basement ofPalaeozoic or older metamorphic rocks intrudedby Permo-Triassic granites associated with acid vol-canic rocks. These rocks are overlain by undated,probably Lower Jurassic, terrestrial sediments andby Jurassic and Cretaceous marine shales and lime-stones. In the western parts of the islands there areEocene to Neogene limestones (Garrard et al.1988; Supandjono & Haryono 1993; Surono &Sukarna 1993).

The block is a continental fragment derived fromnorthern Australia (e.g. Audley-Charles et al. 1972;Hamilton 1979; Pigram et al. 1985) which collidedwith a subduction margin represented by the ophio-lites and associated rocks of East Sulawesi.Hamilton (1979) suggested it was sliced from NewGuinea and carried westward along a strand of theSorong Fault system and this view has becomewidely accepted and incorporated in many tectonicmodels (e.g. Pigram et al. 1985; Garrard et al.1988; de Smet 1989; Daly et al. 1991; Smith &Silver 1991; Hall et al. 1995; Hall 1996, 2002).The collision is generally thought to have occurredin the Neogene (Simandjuntak & Barber 1996) buta wide range of ages has been suggested includ-ing Late Oligocene or Early Miocene (Milsomet al. 2001), within the Miocene (Hamilton 1979),

TOGIAN ISLANDS AND GORONTALO BAY 179

Early to Middle Miocene (Bergman et al. 1996),Middle Miocene (Sukamto & Simandjuntak 1983;Simandjuntak 1986), Middle Miocene to Pliocene(Garrard et al. 1988) and Late Miocene (Silveret al. 1983b; Davies 1990; Smith & Silver 1991;Parkinson 1998). Buton–Tukang Besi has beensuggested to be another microcontinental fragment(Hamilton 1979) that collided in the Early orMiddle Miocene (Fortuin et al. 1990; Smith &Silver 1991), after strike-slip faulting sliced it fromNew Guinea.

Although these microcontinents are small, theircollisions are often interpreted to be responsiblefor widespread deformation in Sulawesi andBorneo. Westward thrusting of the central Sulawesimetamorphic belt, a foreland fold and thrust belt inwest Sulawesi, deformation in the Makassar Straits,deformation in the Meratus Mountains, and inver-sion in the Kutei basin have been attributed to thecollision (e.g. van de Weerd & Armin 1992;Coffield et al. 1993; Simandjuntak & Barber 1996;Pubellier et al. 1999; McClay et al. 2000). Manyauthors suggest the collision, or collisions, followedwestward subduction of ocean lithosphere (e.g.Garrard et al. 1988) interpreted to have produced amagmatic arc in West Sulawesi (e.g. Hamilton1979; Parkinson 1991) or alternatively post-collisional magmatism (e.g. Bergman et al. 1996;Polve et al. 1997; Elburg et al. 2003).

The age of collision is difficult to determine andcould vary within Sulawesi. It requires dating of ter-restrial clastic rocks (‘Celebes Molasse’) that restunconformably on deformed sedimentary, meta-morphic and ophiolitic rocks. In the East ArmUmbgrove (1938) reported a Lower Mioceneunconformity, Brouwer et al. (1947) recorded iso-clinal folding of Early to Middle Miocene age, andKundig (1956) interpreted a Middle Miocene oro-genic phase followed by molasse sedimentationand later Pliocene folding. Hamilton (1979)reported that ‘lower Miocene strata are fullyinvolved in the imbrication and upper Mioceneclastic rocks were derived from the thrust belt’.Other authors have reported Middle Miocenefolding and thrusting (e.g. Audley-Charles et al.1972; Audley-Charles 1974; Katili 1978; Parkinson1991). Surono (1995) suggested that conglomeratesfrom the SE Arm are the oldest Lower to MiddleMiocene parts of the Langkowala Formation whichrests unconformably upon the ophiolite. In Buton,Smith & Silver (1991) interpreted a deformedcomplex including Upper Eocene or Lower Oligo-cene pelagic limestones to be overlain by LowerMiocene conglomerates, but because of the lack ofophiolite detritus interpreted the conglomerates tobe the product of erosion associated with slicing ofthe block from New Guinea rather than collision.They suggested that separate microcontinents may

have collided with East Sulawesi or that a singlelarge microcontinent may have been fragmentedduring oblique collision.

Recent work has cast doubt on the existence of asubduction-related volcanic arc in West Sulawesiduring most of the Palaeogene and Neogene(Polve et al. 1997; Elburg et al. 2003). There isalso little evidence for a collision that affectedWest Sulawesi (Hall & Wilson 2000; Calvert &Hall 2007), and it is now known that the NorthBanda basin formed by oceanic spreading duringthe Middle Miocene (Hinschberger et al. 2000).Spakman & Hall (2010) have proposed a tectonicmodel for the Banda and Sulawesi region that recon-ciles these and other observations with earlierinterpretations, and offers an alternative to the pre-viously accepted idea of slicing of continentalslivers from New Guinea. There was an EarlyMiocene collision of the Sula Spur with the NorthArm volcanic arc and East Arm ophiolite, and thiscontinental area was then fragmented during exten-sion caused by subduction rollback into the Bandaembayment.

Celebes Molasse

Pre-Miocene rocks of the different provinces areunconformably overlain by the Celebes Molasse –a weakly to moderately consolidated associationof interbedded sedimentary formations that is wide-spread across Sulawesi (Sarasin & Sarasin 1901;van Bemmelen 1949). Sediments include conglom-erate, quartz sandstone, greywacke and mudstonewith subordinate intercalations of breccia, marland coral limestone (e.g. van Bemmelen 1949; vanLeeuwen et al. 2007). They have been interpretedto reflect deposition in a coastal alluvial plainenvironment situated along the flanks of rapidlyuplifting and eroding mountains (Calvert 2000).The Celebes Molasse was originally interpreted torelate to a single Miocene collision (Kundig1956). More recently it has been suggested to bediachronous across Sulawesi, representing severaltectonic events (Hall & Wilson 2000). WithinWest Sulawesi and the East Arm it is interpretedto represent latest Miocene to Plio-Pleistoceneuplift and erosion (Hall & Wilson 2000).

Stratigraphic observations

We present a new stratigraphy (Fig. 2) for thewestern, central and eastern Togian Islands(Fig. 3). Based on new field observations and lab-oratory analyses, we define three new units, theWalea Formation, Peladan Formation and BentengIntrusives, and integrate them with the previouslyrecognized Lamusa Formation (Rusmana et al.


1982, 1993), Bongka Formation (Rusmana et al.1993) of the Celebes Molasse (Sarasin & Sarasin1901; van Bemmelen 1949), Lonsio Formation

(Rusmana et al. 1982, 1993) and Luwuk Formation(Garrard et al. 1988). Our new stratigraphy rangesin age from possible Mesozoic basement rocks

Fig. 2. Schematic Neogene stratigraphy of the western, central and eastern Togian Islands, incorporating the age rangesderived in this study. Age (Ma) from Gradstein et al. (2004); PZ, Planktonic Foraminiferal biozones fromBouDagher-Fadel (2008); LS, Far East Letter Stages from BouDagher-Fadel (2008). Note that the timescale isnot linear.


Fig. 3. Simplified geological map of the Togian Archipelago, modified from Rusmana et al. (1982, 1993) based on new field observations. Island names in bold italics; populationcentres in regular. Open circles and/or underlined dip measurements indicate locations examined in this study. Other structural information from Rusmana et al. (1982, 1993).Arrows and bold numbers (all prefixed RTG-) highlight the location of samples explicitly discussed in the text, for which GPS locations (decimal degrees) are listed in the inset table.

M.

A.

CO

TT

AM

ET

AL

.182

through to Quaternary deposits (Fig. 2). The mostcomplete section is seen in the eastern islands(Walea Kodi and Walea Bahi; Fig. 3) where base-ment rocks, possibly of Eocene to Oligocene age,are overlain by Middle Miocene, Pliocene andQuaternary strata. The central and western islandsexpose more restricted sections dominated by Plio-cene volcaniclastics and clastics respectively.

Lamusa formation

Indurated sedimentary rocks of different types occurin several small exposures at the southern end of thechannel between the islands of Batu Daka andTogian (Fig. 3). The rocks are weakly bedded anddip to the north. Lithologies include calcareoussandstones, interbedded with non-calcareous sand-stones and dark mudstones, and dark, fine-grainedrecrystallized limestones. They are heavily brec-ciated and crushed. All lithologies are cut by smallextensional faults. No fossils or sedimentary struc-tures were identified. The formation has a min-imum thickness of 3 m, but neither the top, nor thebase was seen. Following Rusmana et al. (1993)we assign these rocks to the Mesozoic LamusaFormation. Their highly indurated and veined char-acter is consistent with the Mesozoic age suggestedby Rusmana et al. (1993) and suggests that they mayform part of the basement of the Togian Islands.

Walea formation

Arc-related volcanic and volcaniclastic rocks areobserved in exposures along the western coast ofWalea Bahi and eastern coast of Walea Kodi.They include volcanic breccias, pillow lavas andarc-derived volcanogenic sediments. Well-beddedvolcanogenic sedimentary rocks are exposed as alarge, possibly fallen, block on the west coast ofthe southern peninsula of Walea Bahi. Medium-grained, feldspar-rich, grey-brown beds are inter-bedded with green and blue-grey units with a fine-grained green matrix on a scale of c. 5 cm. Allshow internal stratification and possible grading.Further north, just south of a large coastal embay-ment, a larger outcrop exposes an in-situ sectionof gently dipping (19–298 to the east) volcanogenicsediments (Fig. 4a) including interbedded sandsand silts, some of which are calcareous. Mostlybeds are laterally persistent with normal grading,parallel and cross-lamination and ripple crossbedding. Bedding parallel bioturbation and waterescape structures are evident in the more sandylayers. Finer-grained siltstones dominate the upperpart of the exposed sequence. The volcanogenicsedimentary rocks are interpreted to have beendeposited as turbidites and debrites in a deepwaterarc-related setting.

Volcanic rocks occur along the west coast ofthe southern peninsula of Walea Bahi and in coastaloutcrops along the channel between Walea Bahiand Walea Kodi. They include breccias, pillowlavas and more massive and layered lavas of basal-tic to andesitic composition. The rocks have a fine-grained groundmass of feldspar, pyroxene andaltered olivine + phenocrysts of plagioclasefeldspar + amygdales (up to 1 cm) of zeolite and/or calcite. Large blocks (1 � 0.6 m) of breccia areexposed in the beach along the west coast of thesouthern peninsula of Walea Bahi. There are large,sub-rounded, clasts of dark grey (c. 10 cm) andgreen (c. 6 cm) material within a light grey matrix.The clasts have within them feldspar phenocrystsand amygdales of low-grade epidote-rich alterationproducts. Further north, pillows are exposed inseveral outcrops along the channel between WaleaBahi and Walea Kodi, often forming small head-lands. Pillows are grey greenish in colour, weather-ing to grey brown. In most places they are heavilyweathered and altered with late-stage alterationalong fractures. Where relatively fresh, pillowsshow spectacular teardrop shapes (around 30 cmacross), picked out by dark, glassy chilled rims ofbetween 0.5 and 3 cm and fine-grained interpillowmaterial (Fig. 4b), which provide right way-upcriteria. Pillows contain abundant zeolite and/orcalcite amygdales up to 1 cm in size; chilled rimscontain small (5 mm) amygdales and alter to rustycoloured skins where weathered. More massive,layered lavas are also present; individual flows aremarked by craggy tops and brecciated areas.

Rusmana et al. (1982) reported similar pillowlavas, breccias, conglomerates and sandstonesfrom Poh Head (Fig. 1), at the east end of the EastArm, and within the eastern Togian Islands, assign-ing them both to the Miocene Malik Formation.Simandjuntak (1986) assigned basaltic rocks fromPoh Head to the basalt zone of the Balantak Ophio-lite, and suggested a Late Cenomanian to Eoceneage based on K–Ar ages. In a later revision,Rusmana et al. (1993) assigned these rocks to theCretaceous Mafic Complex, whilst those in theeastern Togian Islands were reassigned to the Mio-Pliocene Lonsio Formation (see below).

Based on new observations we assign the basal-tic lavas and volcanogenic sedimentary rocks ofthe Togian Islands to the Walea Formation, a newformation named from the type localities on theislands of Walea Bahi and Walea Kodi. Neitherthe top, nor the base, of the Walea Formation isobserved but there is a minimum thickness of 5 mof pillows and 7 m of volcanogenic sediments.The total thickness of the formation is probablymuch greater. The exact age of the Walea Formationis unknown, but it is the stratigraphically oldest andstructurally lowest unit seen in the eastern islands.


Volcanic arc sedimentary rocks have not beenreported from the East Arm ophiolite. Their associ-ation with basaltic lavas is more similar to the oldestrocks known from the North Arm which formed inan intra-oceanic arc between the Middle Eoceneand earliest Miocene.

Peladan formation

Hard, indurated limestones occur on (at least two)small islands situated around 250 m off the centralwest coast of Walea Bahi. Lithologies includemicritic wackestones and packstones with

30 cm

20 cm

20 cm ~2 m

a b

c

d

Fig. 4. Field photographs of the Walea Formation and Lonsio Formation. (a) Arc-related (?) volcaniclastic sediments ofthe Walea Formation. (b) Basaltic pillows of the Walea Formation exposed on the west coast of Walea Bahi. (c)Well-bedded tuffs of the Lonsio Formation. Coarser tuff units (centre of image beneath pen) show rough stratification,dewatering and cross bedding. Finer tuff units (upper and lower sections of photograph) are more massive, haveirregular bases and show an increase in joint density towards the upper boundary (lower section of photograph). (d)Syn-sedimentary folding and faulting within the Lonsio Formation.


planktonic and benthic foraminifera and fine-grained volcanogenic material (Table 1). These out-crops define the type section for the new PeladanFormation. Benthic and planktonic foraminiferaindicate shallow inner platform or fore-reef shelfand deeper inner platform environments (Table 1).The sequence is well-bedded on a decimetre scale,up to a maximum of around 1 m (mode c. 30 cm)and dips gently towards the north. The sequencehas a minimum observed stratigraphic thickness ofaround 12 m. The top and base of the sequence isnot seen and the true thickness may be muchgreater. No other structure (folding/faulting) wasobserved. In places the beds have a rubbly textureinterpreted to reflect re-working of componentsprior to deposition. Some thin (c. 10 cm), finergrained horizons appear not to have been reworked.In places the limestone are partially dolomitized.

Early–Middle Miocene limestones of a similarage and character are reported from the North andEast Arms of Sulawesi (e.g. Sukamto 1973; Norvick& Pile 1976; Rusmana et al. 1982; Garrard et al.1988; van Leeuwen & Muhardjo 2005). In theTogian Islands Rusmana et al. (1982) previouslyassigned these rocks to the Salodik Formation andsuggested a Late Paleocene to Early Miocene age.Later, Rusmana et al. (1993) reassigned them tothe Lonsio Formation tuffaceous units.

Micropalaeontological analyses of larger fora-minifera and planktonic foraminifera were per-formed on five samples of the Peladan Formation(Table 1). Nannofossil dating was not attempted.We correlate the standard Planktonic Foraminiferalbiozones (PZ) with the ‘Letter Stages’ (LS) of theFar East (as defined by BouDagher-Fadel 2008),relative to the geological timescale of Gradsteinet al. (2004). Analyses indicate a late MiddleMiocene age (PZ: Late N12 – Early N17; LS ofthe Far East: Late Tf2 – Early Tg). Based on theirlithology and age we assign these rocks to the newPeladan Formation, named for one of the twoislands on which they were observed.

Bongka formation (Celebes Molasse)

Weakly to moderately consolidated interbeddedsediments with characteristic lithic-rich horizonsoccur in heavily weathered outcrops along the chan-nel between the islands of Batu Daka and Togian.They are sub-rounded, green-brown, medium-grained sandstones with bands of coarser, angularlithic fragments, medium-grained sandstones witha slabby, bedded character, and brecciated materialwith possible ultrabasic content. Petrographicalanalyses reveal a matrix of serpentinite-rich mate-rial. The sequence dips moderately (c. 308) to thenorth. Sediments with a coarser grain size, but ofcomparable composition and structure (moderate

c. 308 dips to the NE/NNE) are seen in a moreextensive cliff outcrop on a small island east ofthe village of Katupat. Lithologies at this locationinclude laminated siltstones and sandstones, andpebble conglomerates containing well-rounded peb-bles (up to 2 cm) dominated by ophiolitic material(basalts, dolerites, gabbros and serpentinite) withsome chert and limestones. The silts and sandscontain abundant, highly oxidized, plant mate-rial. The formation has a minimum thickness, asobserved in outcrop, of 15 m but neither the topnor the base of the unit is seen.

Similar deposits, but coarser still in grain size,were observed in roadside outcrops on the northerncoast of the East Arm of Sulawesi, west of the townof Bunta (Rusmana et al. 1982, 1993; this study).Here they comprise coarse, massive, sandstoneswith pebble-rich horizons that include large clasts(up to 3 cm) of red chert and cobbles (up to15 cm) of basalts, dolerites, gabbros, metagabbrosand serpentinite with some limestones. Again, thesequence dips north at moderate angles of c. 308.We observed a minimum stratigraphic thickness ofaround 20 m, although neither the top nor the baseof the unit was seen. Based on strong lithologicaland compositional similarities between these rocksand those within the Togian Islands, we followRusmana et al. (1993) in assigning all of theserocks to the Bongka Formation of the CelebesMolasse. In northern Sulawesi palaeontologicaldating of the Celebes Molasse suggests a LateEarly Pliocene to Mid Pleistocene age (Norvick &Pile 1976; Ratman 1976; Hadiwijoyo et al. 1993;Chamberlain & Seago 1995). Late Miocene–Pliocene ages have been reported for the East Arm(Surono & Sukarna 1996).

The Celebes Molasse has been interpreted asalluvial fan and coastal fan delta deposits thatreflect the deposition of locally sourced sedimentin alluvial plain environments with a marginalmarine influence (Calvert 2000). In contrast, theophiolitic material observed in the Togian Islandshas no local source, and such material can onlyhave been derived from the East Arm Ophiolite.Based on the relative grain size and shared structuralcharacteristics (gentle north dip), we suggest thatoutcrops of the Bongka Formation within the EastArm and the Togian Islands represent proximal(coarser) and distal (finer) alluvial fan depositsrespectively, both having been transported northfrom the interior of the East Arm.

Lonsio formation

Volcaniclastic rocks are extensively exposed incoastal outcrops on the northern peninsula of TalaTeoh, the north coast of Togian and the west coastof Walea Kodi. They are grain-supported rocks


Table 1. Biostratigraphical age, facies and palaeoenvironmental analyses of the Peladan formation

Sample ID Depositionalenvironment

Microfacies Components Age (PZ/LS)* (basedon first appearance)

RTG 18 A Shallow innerplatform/fore-reefshelf

Micritic packstone of planktonic andbenthic foraminifera. Micriticpatches reworked into the matrix.

Benthic foraminifera: Cycloclypeus indopacific, Katacycloclypeusmartini, Amphistegina spp., Cycloclypeus pillaris, Cycloclypeusspp., Sphaerogypsina spp., Lepidocyclina spp., Lepidocyclina(Nephrolepidina) spp., L. (Nephrolepidina) angulosa Planktonicforaminifera: Sphaeroidinellopsis spp., Globorotaliapraemenardii, Globigerinoides spp., Globorotalia peripheroacuta,Globorotalia praefohsi, Globoquadrina altispira, Planorbulinellasolida Globoquadrina spp., Globoquadrina dehiscens, Echinoidspp., fragments of rodophyte algae.

Late N12/Late Tf2

RTG 18 B Shallow innerplatform/fore-reefshelf

Micritic packstone of larger benthicforaminifera

Benthic foraminifera: Cycloclypeus spp., Cycloclypeus pillaria,Cycloclypeus carpenteri, Amphistegina spp., Discogypsina discus.Textularia spp., Carpenteria spp., Katacycloclypeus annulatus,Planorbulinella spp. Planktonic foraminifera: Dentoglobigerinaaltispira, Globigerinoides primordius, Globigerina spp.,Globigerinoides quadrilobatus, Orbulina suturalis, Globorotaliapraemenardii, Echinoid spp., fragments of rodophyte algae andcorals, Gastropods, fragments of bryozoa.

Late N12/Late Tf2

RTG 18 C Shallow innerplatform/fore-reefshelf

Micritic packstone of recrystallizedalgae and benthic foraminifera.Micritic patches reworked into thematrix.

Benthic foraminifera: Cycloclypeus spp., Amphistegina spp.,Textularia spp., Miliolid spp., Sphaerogypsina spp. Planktonicforaminifera: Globigerinoides quadrilobatus, Orbulina spp.,Globigerinoides spp., Globorotalia menardii, fragments ofrodophyte algae, Lithophyllum spp., Lithothamnium spp.,Gastropods, Echinoid spp., rare fragments of bryozoa.

N12 and younger/Tf2and younger

RTG 18 D Shallow innerplatform/fore-reefshelf

Micritic packstone of foraminiferaand algae.

Benthic foraminifera: Cycloclypeus pillaria, Planorbulinella solida,Gypsina spp., Sphaerogypsina spp., Elphidium spp., Nodosariaspp. Planktonic foraminifera: Globoquadrina spp.,Globigerinoides trilobus, Globigerinoides spp., Orbulina suturalis,Globorotalia conoidea, Globorotalia menardii, Globorotaliascitula, Gastropod spp., fragments of bryozoa, fragments of coral.

Late N12 – EarlyN17/Tf3 – EarlyTg

RTG 18E Relativelydeeper innerplatform

Micritic wackestone of foraminifera.Reworked patches of micrite arealso present.

Benthic Foraminifera: Lepidocyclina spp., Carpenteria spp.,Cycloclypeus spp., Cycloclypeus pillaria, Operculina spp.,Heterostegina spp., Gypsina spp., Planorbulinella larvata, Lagenaspp., Textularia spp. Planktonic foraminifera: Globoquadrinaaltispira, Globorotalia spp., Globorotalia scitula, Globoquadrinadehiscens, Globorotalia menardii, Globoquadrina dehiscens,Globorotalia fohsi, Ostracod spp., Gastropod spp.

Late N12 – EarlyN13/Late Tf2 –Early Tf3

*We correlate the standard Planktonic Foraminiferal biozones (PZ) with the ‘Letter Stages’ (LS) of the Far East (as defined by BouDagher-Fadel 2008), relative to the biostratigraphical timescale (as defined byGradstein et al. 2004)

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with a carbonate (dominantly sparry) matrix. Micro-fossils and algal fragments are also embeddedwithin the matrix; their abundance varies betweenunits (Table 2).

The volcaniclastics are well-bedded, commonlyat the decimetre scale, with a maximum bed thick-ness of around 3 m. Two main bed types alternateat a range of scales. Individual beds appear laterallypersistent at the outcrop scale. Stratified beds aretypically around 10 to 30 cm in thickness andshow parallel lamination of fine to coarse sand.They contain rare horizons of small (up to finepebbles) angular lithic fragments (Fig. 4c). Inplaces the beds show spectacular dewatering struc-tures, and may be wholly or partly cross-bedded,producing an irregular upper surface. The base ofthe beds is almost universally planar. Stratifiedunits are overlain by fine-grained cream colouredmaterial, which show little variation in grain sizeor internal structure (Fig. 4c). Beds range in thick-ness from cm scale to a maximum of 3 m. Theirbases are commonly irregular, reflecting the topo-graphy of the stratified layer below, and theydisplay an increase in joint density towards theupper boundary, which is characteristically planar.In places the finer beds may be very thin, or entirelyabsent from the sequence. Overall, the sequencedips gently in various directions. Locally, therocks dip steeply and show intense syn-sedimentaryfolding and faulting (Fig. 4d), interpreted toreflect soft sediment deformation. The sequencehas a minimum stratigraphic thickness of around20 m, however, the top and base of the sequenceis not seen and the true thickness is probablymuch more.

Comparable volcaniclastic rocks are observed onPoh Head, where they include thick sequences ofcoarse stratified units (this study). Rusmana et al.(1982, 1993) described these rocks as tuffaceoussediments and assigned them to the Lonsio For-mation. Simandjuntak (1986) interpreted similarvolcanogenic sediments from the East Arm asmegacyclic turbidites, and assigned them to theLonsuit Turbidites of the Batui Group. We interpretthese rocks as tuffaceous sediments that reflect rapidaqueous reworking of primary volcaniclastic mate-rial during deposition in a shallow marine environ-ment soon after eruption. Microfossil observationssuggest depths less than 200 m. Stratificationreflects crude sorting of coarse ash during settlingthrough the water column; cross-bedding may reflectturbidity currents formed by ash initially held insuspension. Finer-grained ash settled more slowlythrough the water column, draping topography inthe underlying coarse units. Pumice is largelyabsent and may have been floated off and not pre-served (e.g. Freundt 2003). The repeated sequenceof coarse and fine tuff may reflect pulses within a

single eruption or input from several eruptions.Dewatering structures within coarse units suggestrapid loading by the subsequent fine units. Jointingpresent near the upper boundary of the finer unitsmay be syn- or post-depositional.

Based on their striking similarity to tuffaceousunits observed on Poh Head (Simandjuntak 1986;A. J. Barber, pers. comm. 2009) we assign theserocks to the Lonsio Formation of Rusmana et al.(1982, 1993). Micropalaeontological analyses oflarger foraminifera and planktonic foraminiferawere performed on five tuff samples from theLonsio Formation (Table 2). Nannofossil datingwas not attempted. Foraminiferal assemblagesrange from N4 and younger (PZ) and Te andyounger (LS), and constrain a Late Miocene toEarly Pliocene age (PZ: N19; LS: Early Th).

Benteng Intrusives

Intrusive rocks of intermediate composition areexposed in isolated outcrops, along the northernand southern coasts of Togian Island. They occuras small intrusions, often forming topographichighs and small islands. We infer the presence ofadditional intrusive bodies within the interior ofTogian Island based on the presence of isolatedsteep-sided topographic highs visible from thecoast as shown on the map of Kundig (1956). Therocks have a fine to medium grained light-greygroundmass with phenocrysts of phlogopite mica(up to 7 mm) + feldspar (6–7 mm) + hornblende(1–3 mm) + mafic xenoliths (up to 2 cm). In placesfeldspar phenocrysts are concentrated into ‘trails’up to 20 cm long. Orthogonal sub-horizontal andsub-vertical joints spaced at around 20 to 50 cm,and resulting in a characteristic blocky appearance,suggest intrusion at shallow depths. In places therocks are cut by east–west trending brittle faults,producing breccia zones around 1 m wide.

These rocks are classified (Fig. 5; Table 3) astrachydacites and trachyandesites on the totalalkalis v. silica (TAS) diagram of Le Maitre (1989)(they are syenites on TAS diagrams adapted for plu-tonic rocks (e.g. Wilson 1989)) and belong to thealkaline magma series (Kuno 1966; Irvine &Baragar 1971). They have an extremely K-richchemistry and plot within the shoshonitic field ofRickwood (1989) on a K2O v. SiO2 diagram. Intru-sive intermediate rocks were first recognized onTogian Island by Kundig (1956), who identifiedrocks of andesitic composition. These were sub-sequently misidentified as basaltic (Rusmana et al.1982) or volcaniclastic (Rusmana et al. 1993) incharacter. We assign these intrusive rocks to thenew Benteng Intrusives, named for the village ofthe same name in south central Togian Island(Fig. 3).


Table 2. Biostratigraphical age, facies and palaeoenvironmental analyses of the Lonsio formation

Sample ID Depositional environment Microfacies Components Age (PZ/LS)* (basedon first appearance)

RTG 25 Inner neritic, planktonic & shelfbenthic foraminifera drifted/reworked into volcanic deposits.

Sparitic packstone of volcanicsediments rich in embeddedplanktonic foraminifera and rarelarger benthic and algaefragments

Globoquadrina altispira, Globoquadrina spp.,Orbulina spp., Globorotalia margaritae,Globorotalia scitula, Sphaeroidinellopsissubdehiscens, Globigerinoides trilobus,Globigerinoides quadrilobatus, Globorotaliaacostaensis, Fragments of rodophyte algae,Elphidium spp.

N19/Early Th

RTG 26 Inner neritic Sparitic packstone of volcanicsediments with rare embeddedplanktonic foraminifera

Globigerinoides spp. N4 and younger/UpperTe and younger

RTG 27 Sparitic packstone of volcanicsediments with rare embeddedplanktonic foraminifera

Globigerinoides spp. N4 and younger/UpperTe and younger

RTG 30 Inner neritic, planktonic & shelfbenthic foraminifera drifted/reworked into volcanic deposits.

Sparitic packstone of volcanicsediments rich in embeddedplanktonic foraminifera and rarelarger benthic and algaefragments

Catapsydrax spp., Orbulina universa,Globoquadrina dehiscens, Pulleniatinaprimalis, Globoquadrina altispira,Globorotalia globosa, Globorotaliahumerosa, Globorotalia mayeri, Globorotaliascitula, Globigerinoides sacculifer,Globigerinoides quadrilobatus, Elphidiumspp., Amphistegina spp., Heterostegina spp.,Asterigerina spp.

N19/Early Th

RTG 36 Inner neritic Sparitic packstone of volcanicsediments with rare embeddedplanktonic foraminifera

Orbulina universa, Globigerinoides spp.,Globigerinoides quadrilobatus,Globoquadrina spp.

N4-N19/Upper Te –Early Th

*We correlate the standard Planktonic Foraminiferal biozones (PZ) with the ‘Letter Stages’ (LS) of the Far East (as defined by BouDagher-Fadel 2008), relative to the biostratigraphical timescale (as defined byGradstein et al. 2004)

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Five high-purity mica separates from four sam-ples of the Benteng Intrusives were dated using40Ar/39Ar techniques. Samples were crushed,graded using disposable nylon cloth sieves in abrass collar and separated using conventional elec-tromagnetic techniques. High-purity mineral separ-ates were handpicked from the 63–250 mm fraction,and for RTG-12 from the .250 mm fraction, thusany contamination in the analyses is assumed to bedue to intra-grain alteration and/or contaminants.All analyses were undertaken in the Argon Labora-tory of the Research School of Earth Sciences, TheAustralian National University, using the furnacestep-heating technique (Table 4). Samples wereirradiated at the McMaster Nuclear Reactor,McMaster University, Canada using Sanidine 92–176 from Fish Canyon Tuff, Colorado (K/Ar refer-ence age 28.10 + 0.04 Ma) as the Fluence Monitor(Spell & McDougall 2003). Ages were calculatedusing the 40K abundances and decay constants ofSteiger & Jager (1977). Uncertainties in isotopicratios and ages are quoted at the 1s level.

For all samples plots of 36Ar/40Ar v. 39Ar/40Ardemonstrate the presence of one main gas popu-lation, with varying amounts of contaminants (suchas excess argon), and a large atmospheric argoncomponent – particularly in the coarser grainedsamples (Fig. 6). The oldest ages are preservedin the high-temperature heating steps of coarse-grained (.250 mm) biotite from samples RTG12(2.40 + 0.01 Ma; MSWD (mean square of weighteddeviation) 1.58) and RTG31 (2.02 + 0.01 Ma;MSWD 0.01) (Fig. 6). However, significant atmos-pheric argon contents, and evidence of argon lossand possible younger events render the meaningof these ages ambiguous. Analysis of fine-grained(250–63 mm) biotite from sample RTG12 containssignificantly less atmospheric argon than thecoarser-grained biotite and produced a reliable, con-sistently flat spectrum of 1.80 + 0.01 Ma (MSWD3.95) (Fig. 6). This analysis provides the best agefor this sample and the most robust age for theBenteng Intrusives. Analyses of fine-grained micafrom two other samples gave robust Pleistoceneages. Despite disturbance during the initial heat-ing steps (linked to variation in Ca), over 50% ofthe gas emitted from RTG08 produced a strongplateau with an age of 1.52 + 0.02 Ma (MSWD0.3) (Fig. 6). Except for several contaminated inter-vening steps, analysis of fine-grained biotite fromRTG09 would have produced a similar plateau,giving an age of 1.68 + 0.09 Ma (MSWD 3.9)with a younger age of 1.37 + 0.02 Ma evident(Fig. 6).

Based on our new field observations and labora-tory analyses we interpret these rocks as shallowlevel stocks and dykes of Late Pliocene to EarlyPleistocene age. The observed and inferred intrusive

Table 3. Major element data (weight %) for samplesof the Benteng Intrusives analysed in this study

Sample ID RTG08 RTG09 RTG12 RTG31

SiO2 63.36 63.14 58.97 61.39Al2O3 15.52 15.43 15.85 14.54Fe2O3 3.92 3.77 4.09 5.82MgO 2.25 2.37 2.97 2.24CaO 3.67 3.77 4.38 2.21Na2O 4.41 4.62 3.88 3.44K2O 5.40 5.34 6.51 6.46TiO2 0.34 0.33 0.76 0.69P2O5 0.42 0.41 0.45 0.43MnO 0.09 0.09 0.06 0.12

Total 99.38 99.26 97.92 97.34

Fig. 5. Major element classification diagrams for thevolcanic rock samples analysed in this study. (a) Totalalkalis (K2OþNa2O) v. silica (SiO2) diagram. Fieldboundaries are those of Le Maitre (1989): 1, andesite; 2,dacite; 3, trachyandesite; 4, trachydacite. Subdivisioninto alkaline and sub-alkaline series: dashed curved line– Irvine & Baragar (1971); solid curved line – Kuno(1966). (b) K2O v. SiO2 diagram. Series boundaries andnomenclature: dashed lines and bold italics, Le Maitre(1989); solid lines and nomenclature in parentheses, afterRickwood (1989).


Table 4. 40Ar/39Ar step heating analyses

Temp Ar36 err Ar37 err Ar38 err Ar39 err Ar40 err Ar40* Ar40*/ Cumulative Calculated age Ca/K(8C) (mol) (%) (mol) (%) (mol) (%) (mol) (%) (mol) (%) (%) Ar39(K) Ar39(%) Ma + 1s.d.

Sample RTG-08 (R1) Biotite600 4.19E-16 0.87 2.04E-15 3.90 1.79E-16 3.38 6.52E-15 0.24 1.30E-13 0.25 4.60 0.92 0.71 2.87 + 0.78 0.59650 1.47E-16 1.46 1.89E-15 6.16 1.40E-16 3.73 6.08E-15 0.56 4.77E-14 0.64 8.90 0.70 1.37 2.20 + 0.36 0.59700 2.48E-16 0.73 1.96E-15 3.02 2.12E-16 1.75 1.28E-14 0.20 8.16E-14 0.21 10.00 0.64 2.76 2.00 + 0.14 0.29750 4.99E-16 0.48 2.07E-15 8.62 3.92E-16 0.29 2.52E-14 0.11 1.57E-13 0.16 5.50 0.34 5.49 1.08 + 0.16 0.16800 1.02E-15 0.24 2.88E-15 2.87 7.11E-16 1.25 4.07E-14 0.09 3.24E-13 0.11 6.60 0.53 9.92 1.66 + 0.09 0.14840 1.28E-15 0.38 4.70E-15 3.48 8.54E-16 0.21 4.97E-14 0.09 4.00E-13 0.11 5.30 0.43 15.32 1.35 + 0.13 0.18890 2.90E-15 0.26 7.08E-15 2.45 1.66E-15 0.91 9.11E-14 0.11 9.10E-13 0.13 5.80 0.58 25.22 1.83 + 0.12 0.15930 2.91E-15 0.41 9.44E-15 3.38 2.09E-15 0.54 1.27E-13 0.11 9.24E-13 0.18 6.60 0.48 39.02 1.51 + 0.14 0.14970 2.49E-15 0.34 7.38E-15 0.88 2.93E-15 0.63 1.99E-13 0.12 8.37E-13 0.17 11.60 0.49 60.61 1.53 + 0.06 0.071020 1.94E-15 0.48 5.66E-15 3.74 2.42E-15 0.47 1.66E-13 0.09 6.55E-13 0.11 12.10 0.48 78.64 1.50 + 0.08 0.061070 8.81E-16 0.67 6.81E-15 1.57 1.67E-15 0.64 1.21E-13 0.13 3.11E-13 0.20 15.50 0.40 91.83 1.25 + 0.06 0.111140 2.98E-16 0.73 4.17E-14 0.64 9.95E-16 0.72 7.28E-14 0.08 1.19E-13 0.14 28.10 0.46 99.74 1.44 + 0.06 1.091200 5.94E-17 2.41 1.04E-13 1.06 4.91E-17 0.80 2.00E-15 0.50 1.25E-14 0.53 45.10 2.94 99.95 9.20 + 0.90 103.001350 4.00E-17 3.20 1.85E-14 5.07 1.27E-17 20.02 4.88E-16 0.61 1.04E-14 0.70 4.20 0.92 100.00 2.89 + 4.44 74.20Total 1.51E-14 2.16E-13 1.43E-14 9.20E-13 4.92E-12 0.49 1.53 + 0.10

Lambda K40 ¼ 5.5430E-10 J ¼ 1.7413E-3 +0.413

Sample RTG-09 (R2) Biotite600 9.22E-16 0.55 2.25E-15 5.24 3.05E-16 0.99 7.96E-15 0.28 2.78E-13 0.31 2.10 0.72 0.32 2.26 + 0.94 0.54650 5.83E-16 0.63 1.99E-15 2.74 2.54E-16 0.60 1.06E-14 0.21 1.78E-13 0.26 2.90 0.49 0.75 1.53 + 0.35 0.36700 6.71E-16 0.45 3.90E-15 2.72 4.28E-16 0.66 2.20E-14 0.13 2.08E-13 0.16 4.30 0.41 1.65 1.27 + 0.15 0.34750 1.09E-15 0.72 4.56E-15 2.40 7.13E-16 1.27 4.12E-14 0.38 3.45E-13 0.43 6.70 0.56 3.32 1.77 + 0.23 0.21800 4.73E-15 0.36 1.40E-14 4.79 2.73E-15 0.69 1.52E-13 0.09 1.48E-12 0.11 5.40 0.52 9.50 1.64 + 0.13 0.17840 4.03E-15 0.37 1.01E-14 1.43 2.31E-15 0.26 1.28E-13 0.12 1.26E-12 0.13 5.10 0.50 14.70 1.56 + 0.18 0.15930 3.69E-15 0.56 1.43E-14 1.65 3.58E-15 0.59 2.38E-13 0.22 1.23E-12 0.34 11.10 0.58 24.37 1.80 + 0.11 0.11970 1.92E-15 0.39 1.14E-14 2.18 4.60E-15 0.91 3.39E-13 0.10 7.54E-13 0.19 23.60 0.53 38.13 1.65 + 0.03 0.061020 1.40E-15 0.60 1.07E-14 2.75 4.10E-15 0.33 3.07E-13 0.11 5.74E-13 0.13 26.70 0.50 50.58 1.57 + 0.03 0.071070 2.84E-15 0.40 4.04E-14 1.27 9.22E-15 0.40 7.00E-13 0.06 1.16E-12 0.11 26.40 0.44 78.98 1.37 + 0.02 0.111140 2.07E-15 0.80 2.64E-13 0.22 5.79E-15 0.35 4.27E-13 0.09 8.00E-13 0.12 25.60 0.48 96.31 1.50 + 0.04 1.171200 7.59E-16 1.23 4.84E-13 0.36 9.02E-16 1.33 4.59E-14 0.21 2.19E-13 0.32 19.80 0.95 98.16 2.99 + 0.20 20.201350 4.57E-16 0.84 1.07E-13 0.60 6.62E-16 1.33 4.55E-14 0.09 1.53E-13 0.14 18.40 0.62 100.00 1.95 + 0.08 4.47Total 2.52E-14 9.67E-13 3.56E-14 2.46E-12 8.64E-12 0.50 1.57 + 0.06

Lambda K40 ¼ 5.5430E-10 J ¼ 1.7378E-3 +0.413

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Sample RTG-12 (coarse-grained) (R3) Biotite600 2.35E-15 0.47 4.36E-16 16.96 4.87E-16 1.48 2.39E-15 0.25 6.94E-13 0.26 -0.20 0.00 0.08 0.00 + 4.90 0.35650 1.90E-15 0.63 4.40E-16 22.82 4.18E-16 0.35 4.43E-15 0.32 5.60E-13 0.33 -0.10 0.00 0.23 0.00 + 2.69 0.19700 2.57E-15 0.45 2.14E-15 5.57 6.49E-16 0.88 1.20E-14 0.08 7.64E-13 0.10 0.60 0.40 0.65 1.27 + 1.05 0.34750 4.70E-15 0.45 5.49E-15 1.01 1.24E-15 1.91 2.49E-14 0.30 1.41E-12 0.33 1.20 0.68 1.50 2.15 + 1.08 0.42800 1.15E-14 0.28 6.25E-15 2.98 2.93E-15 0.83 5.73E-14 0.11 3.43E-12 0.12 1.20 0.70 3.46 2.20 + 0.58 0.21840 8.35E-15 0.38 5.59E-15 5.75 2.80E-15 0.28 9.55E-14 0.17 2.50E-12 0.21 1.10 0.30 6.74 0.94 + 0.36 0.11900 1.19E-14 0.50 2.25E-15 5.51 5.12E-15 0.45 2.24E-13 0.35 3.61E-12 0.45 2.10 0.34 14.40 1.08 + 0.38 0.02980 1.51E-14 0.43 3.31E-15 10.36 9.69E-15 0.98 5.49E-13 0.17 4.79E-12 0.24 6.70 0.58 33.22 1.83 + 0.13 0.011020 6.52E-15 0.29 2.40E-15 11.44 6.72E-15 0.51 4.40E-13 0.08 2.22E-12 0.14 12.80 0.65 48.30 2.04 + 0.05 0.011060 3.49E-15 1.03 2.18E-15 11.95 5.04E-15 0.53 3.52E-13 0.19 1.28E-12 0.23 18.30 0.67 60.37 2.09 + 0.12 0.011100 2.81E-15 0.98 2.34E-15 4.76 5.28E-15 0.76 3.82E-13 0.08 1.14E-12 0.16 25.90 0.77 73.47 2.42 + 0.07 0.011200 2.76E-15 0.43 1.71E-14 0.79 9.57E-15 0.26 7.21E-13 0.06 1.38E-12 0.08 39.70 0.76 98.19 2.40 + 0.02 0.051350 1.94E-16 2.36 2.49E-15 5.92 6.83E-16 1.77 5.27E-14 0.53 9.48E-14 0.64 38.50 0.69 100.00 2.17 + 0.09 0.09Total 7.41E-14 5.25E-14 5.06E-14 2.92E-12 2.39E-11 0.65 2.03 + 0.13

Lambda K40 ¼ 5.5430E-10 J ¼ 1.7442E-3 +0.426

Sample RTG-12 (fine-grained) (R4) Biotite600 3.70E-16 1.23 2.29E-15 12.90 1.88E-16 2.11 8.05E-15 0.25 1.18E-13 0.29 7.40 1.09 0.32 3.39 + 0.51 0.54650 3.83E-16 0.85 4.55E-15 1.32 2.74E-16 2.01 1.62E-14 0.18 1.23E-13 0.22 7.90 0.60 0.97 1.87 + 0.19 0.53700 3.24E-16 1.14 6.82E-15 3.16 4.00E-16 0.98 2.74E-14 0.20 1.14E-13 0.25 15.60 0.65 2.07 2.02 + 0.13 0.47750 7.46E-16 1.34 2.17E-14 4.21 1.59E-15 2.31 1.15E-13 0.17 2.92E-13 0.19 24.20 0.62 6.65 1.92 + 0.10 0.36800 2.83E-16 0.87 7.56E-15 1.05 9.40E-16 0.45 7.21E-14 0.39 1.26E-13 0.45 32.60 0.57 9.53 1.78 + 0.04 0.20850 3.30E-16 1.68 7.82E-15 3.81 1.60E-15 0.54 1.21E-13 0.10 1.72E-13 0.14 41.70 0.59 14.37 1.85 + 0.04 0.12890 3.08E-16 1.10 5.27E-15 4.20 2.46E-15 0.78 1.85E-13 0.16 2.03E-13 0.23 52.90 0.58 21.75 1.81 + 0.02 0.05930 2.46E-16 0.90 3.26E-15 4.00 3.56E-15 0.29 2.72E-13 0.12 2.37E-13 0.17 66.30 0.58 32.62 1.80 + 0.01 0.02970 3.62E-16 1.10 2.77E-15 13.41 4.57E-15 0.44 3.50E-13 0.34 3.17E-13 0.38 63.50 0.58 46.61 1.80 + 0.01 0.021020 5.43E-16 0.74 3.14E-15 0.94 4.57E-15 0.34 3.51E-13 0.08 3.71E-13 0.10 54.20 0.57 60.66 1.79 + 0.01 0.021070 5.92E-16 0.72 6.44E-15 2.47 4.14E-15 0.25 3.15E-13 0.07 3.58E-13 0.10 48.90 0.56 73.27 1.74 + 0.01 0.041140 9.20E-16 0.59 4.58E-14 0.58 6.96E-15 0.18 5.33E-13 0.13 5.91E-13 0.15 52.40 0.58 94.61 1.81 + 0.01 0.161200 2.96E-16 0.73 8.36E-14 0.72 1.71E-15 0.35 1.29E-13 0.09 1.61E-13 0.14 48.90 0.61 99.78 1.90 + 0.02 1.231350 5.45E-17 2.59 6.65E-15 1.25 8.01E-17 4.55 5.50E-15 0.41 1.88E-14 0.44 17.00 0.58 100.00 1.81 + 0.24 2.30Total 5.76E-15 2.08E-13 3.30E-14 2.50E-12 3.20E-12 0.58 1.81 + 0.02

Lambda K40 ¼ 5.5430E-10 J ¼ 1.7305E-3 +0.356

(Continued)

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AL

OB

AY

191

Table 4. Continued

Temp Ar36 err Ar37 err Ar38 err Ar39 err Ar40 err Ar40* Ar40*/ Cumulative Calculated age Ca/K(8C) (mol) (%) (mol) (%) (mol) (%) (mol) (%) (mol) (%) (%) Ar39(K) Ar39(%) Ma + 1s.d.

Sample RTG-31 (coarse-grained) (R6) Biotite470 6.17E-17 4.37 1.04E-16 24.22 2.80E-17 4.16 1.17E-15 0.50 1.82E-14 0.51 -0.30 0.00 0.07 0.00 + 1.83 0.17510 3.02E-16 1.19 3.13E-16 15.37 1.50E-16 4.28 7.16E-15 0.20 9.00E-14 0.24 0.70 0.08 0.49 0.26 + 0.50 0.08550 7.04E-16 0.89 3.85E-16 2.84 3.39E-16 0.32 1.65E-14 0.21 2.08E-13 0.23 -0.40 0.00 1.45 0.00 + 0.36 0.04600 1.68E-15 0.67 1.12E-15 3.27 9.63E-16 0.59 5.15E-14 0.18 5.10E-13 0.21 2.30 0.22 4.46 0.70 + 0.22 0.04650 2.96E-15 0.36 2.58E-15 0.45 1.99E-15 0.56 1.12E-13 0.11 8.98E-13 0.14 2.50 0.20 10.98 0.63 + 0.10 0.04700 4.48E-15 0.35 4.28E-15 2.73 3.23E-15 0.47 1.91E-13 0.05 1.40E-12 0.10 5.30 0.39 22.14 1.20 + 0.09 0.04750 5.25E-15 0.28 4.91E-15 5.91 3.87E-15 0.58 2.26E-13 0.07 1.66E-12 0.12 6.40 0.47 35.36 1.47 + 0.08 0.04790 3.87E-15 0.81 3.03E-15 2.15 2.68E-15 0.79 1.60E-13 0.66 1.23E-12 0.79 6.90 0.53 44.72 1.65 + 0.28 0.04840 4.88E-15 0.73 3.76E-15 5.11 3.48E-15 1.11 1.99E-13 0.15 1.56E-12 0.16 7.40 0.59 56.32 1.83 + 0.17 0.04890 2.65E-15 0.91 2.99E-15 4.52 2.19E-15 0.73 1.35E-13 0.10 8.55E-13 0.16 8.10 0.52 64.20 1.61 + 0.17 0.04950 3.94E-15 0.57 3.62E-15 7.25 3.00E-15 1.10 1.73E-13 0.26 1.28E-12 0.29 8.80 0.65 74.31 2.03 + 0.16 0.041000 3.55E-15 0.73 3.19E-15 3.05 2.32E-15 0.38 1.32E-13 0.17 1.14E-12 0.21 7.50 0.65 82.01 2.01 + 0.19 0.051050 5.05E-15 0.54 4.65E-15 5.84 3.17E-15 1.20 1.69E-13 0.21 1.59E-12 0.31 6.10 0.58 91.90 1.80 + 0.13 0.051100 2.52E-15 0.77 6.00E-15 2.63 1.74E-15 2.32 1.01E-13 0.14 8.18E-13 0.22 8.90 0.72 97.79 2.25 + 0.18 0.111200 5.26E-16 1.56 6.02E-14 0.42 6.73E-16 1.57 3.66E-14 0.15 1.79E-13 0.19 16.30 0.80 99.93 2.49 + 0.21 3.131350 4.62E-17 13.90 1.28E-15 8.90 4.96E-17 1.36 1.26E-15 1.19 1.75E-14 1.21 22.50 3.14 100.00 9.76 + 1.91 1.93Total 4.25E-14 1.02E-13 2.99E-14 1.71E-12 1.35E-11 0.52 1.62 + 0.16

Lambda K40 ¼ 5.5430E-10 J ¼ 1.7286E-3 +0.426

M.

A.

CO

TT

AM

ET

AL

.192

bodies follow a broadly north–south trend throughthe centre of Togian Island, supporting the spatialobservations of Kundig (1956), and indicating apossible structural control on their intrusion.

Luwuk formation

Reefal limestones are found throughout the archipe-lago, and dominate outcrop in the western islands(e.g. Batu Daka). They occur as high cliffs andraised terraces of poorly bedded, rubbly limestonescontaining broken coral fragments. The limestoneshave been uplifted to heights of around 200 mwithin the archipelago and to more than 300 m onthe East Arm (Garrard et al. 1988). FollowingRusmana et al. (1982) we allocate these rocks a

Quaternary age and assign them to the LuwukFormation (Garrard et al. 1988).

Discussion

The Togian Islands offer a unique opportunity toinvestigate Gorontalo Bay. Our new stratigraphyoffers insight into several aspects of the basinincluding the nature of its basement rocks, its ageand its mode of formation.

Basement rocks beneath Gorontalo Bay

Based on geophysical evidence, Silver et al. (1983b)suggested that much of Gorontalo Bay is underlain

RTG-08

0 20 40 60 80 100Cumulative % 39Ar released

0.0

1.0

2.0

3.0

4.0

5.0

RTG-08


0.0

1.0

2.0

3.0

4.0

5.0Ag

e(M

a)

RTG-31


0.0

1.0

2.0

3.0

4.0

5.0


0.0

1.0

2.0

3.0

4.0

5.0

RTG-09


0.0

1.0

2.0

3.0

4.0

5.0

RTG-12 C


0.0

1.0

2.0

3.0

4.0

5.0

Age

(Ma)

RTG-12 F

Fig. 6. 40Ar/39Ar age spectra plots for biotite step-heating analyses performed on four samples from the BentengIntrusives. For sample RTG12 separate analyses were undertaken on coarse (.250 mm; RTG-12 C) and fine(63–250 mm; RTG-12 F) mica.


by basement rocks belonging to the East SulawesiOphiolite (East Sulawesi Ophiolite). Beneath thesemay be continental basement rocks belonging tothe leading edge of the Banggai-Sula microconti-nental block (Silver et al. 1983a; Hall & Wilson2000). Other hypotheses are that the bay is underlainby oceanic crust of the Northern Sulawesi Province(Monnier et al. 1995) or that the basement of Goron-talo Bay comprises a complex amalgamation of atleast two tectonostratigraphic provinces. The posi-tion of the Togian archipelago in the middle ofGorontalo Bay provides an opportunity to test thesehypotheses.

Our new field observations suggest that thecentral part of Gorontalo Bay, including theTogian Islands, is underlain by oceanic and arcbasement of the Northern Sulawesi Province ratherthan the continental basement of the Banggai-Sula Block. The Walea Formation represents thebasement within the Togian Islands. Its age is notknown but it is inferred to be older than theMiddle Miocene limestones against which it isfaulted. The formation comprises an association ofvolcanic rocks and subordinate volcanogenic sedi-ments that we suggest represent the products of asubmerged volcanic arc rather than an ophiolite, aspreviously interpreted (Kundig 1956). A similarassociation of volcanic rocks and subordinate volca-niclastics is reported within the Papayato Volcanicrocks of the North Sulawesi Province (Elburget al. 2003; van Leeuwen et al. 2007), and is con-sistent with the suggestion that the basement ofthis province continues southwards beneath thearchipelago. Volcanic rocks (breccias, pillows andlavas) similar to those of the lower parts of theWalea Formation are also reported from the Cre-taceous Balantak Ophiolite of East Sulawesi(Simandjuntak 1986; A. J. Barber, pers. comm.2009), but they do not show the same associationwith contemporaneous volcaniclastic sediments.Geochemical and/or geochronological analyses ofthe Walea Formation, and comparison with the(Middle Eocene to Early Miocene) Papayato Vol-canic rocks (North Sulawesi Province of Elburget al. 2003) and the (Cretaceous) Balantak Ophiolite(East Sulawesi Ophiolite) would help to resolve thisissue but the rocks are so deeply weathered thatobtaining suitable material has not so far beenpossible.

Field investigations and geochemical analysessuggest that the western end of the bay is underlain

by continental crust (Elburg et al. 2003; vanLeeuwen & Muhardjo 2005; van Leeuwen et al.2007) as far east as 1218E (Fig. 1). This materialforms the eastern margin of Sundaland and is prob-ably of Australian origin (van Leeuwen & Muhardjo2005), but was accreted to Sundaland during themid-Cretaceous (Parkinson 1991; Parkinson et al.1998; Hall 2009) and is not part of the Banggai-Sulu block.

Continental crust probably continues north fromthe Banggai-Sulu microcontinent beneath theMolucca Sea (Silver et al. 1983b; Watkinson et al.2010). Beneath Gorontalo Bay earthquake hypocen-tres (Engdahl et al. 1998) define the southern edgeof the westward-subducting Molucca Sea plate. Thisis a very sharp, almost WNW–ESE, line (Fig. 7a)that we interpret as the former continental–oceanic crust boundary between the Molucca Seaand the Banggai-Sula block. The position of theline implies that continental crust continues northfrom the Banggai-Sula Islands to the centre of theeastern part of Gorontalo Bay. How far westbeneath Gorontalo Bay the continental crust con-tinues is uncertain; the oil that seeps throughthe ophiolite north of the thrust complex in theEast Arm (Kundig 1956) suggests continental base-ment may extend at least west to about 1228E(Fig. 8).

Miocene carbonate platform

Miocene carbonate rocks are widespread in northernSulawesi. They include the Middle Miocene lime-stones of the Peladan Formation reported here, car-bonates of the Buol Beds in NW Sulawesi (Ratman1976; van Leeuwen & Muhardjo 2005), the SalodikFormation within the East Arm (Rusmana et al.1982), and limestones observed around Palu andthe western Toli–Toli region (Sukamto 1973;Norvick & Pile 1976; van Leeuwen & Muhardjo2005). Benthic and planktonic foraminifera indicatedeposition within inner platform/fore-reef shallowmarine environments during the late Early toMiddle Miocene (van Leeuwen & Muhardjo 2005;this study). Jablonski et al. (2007) report submergedcarbonate reefs in Gorontalo Bay based on seismicobservations which they interpreted as Oligoceneto Middle Miocene in age. The distribution ofMiocene carbonate rocks suggests that GorontaloBay was an area of extensive carbonate platformdeposition during the Miocene. It was probably

Fig. 7. Earthquake hypocentres in Eastern Indonesia based on the dataset of Engdahl et al. (1998). (a) Black crossesdenote all hypocentres, those assigned to the westward subducting Molucca Sea Plate are highlighted with blue dots,those assigned to subduction at the North Sulawesi Trench are highlighted in green. Hypocentres associated withvolcanism at the Una-Una volcano are shown in purple. Red box denotes the line of section illustrated in (b).(b) North–south cross section though Gorontalo Bay and the Una-Una volcano. Hypocentres associated withvolcanism at the Una-Una volcano (purple dots) are notably shallower than those related to the downgoing slab.


Fig. 7 (Continued) (c) Earthquake hypocentres assigned to the Molucca Sea Plate coloured based on depth. To aidclarity, hypocentres less than 75 km depth are not shown. Colouration shows that the slab dips gently to the NW butis sharply terminated along its southern edge in a steep upturned lip. Black crosses denote hypocentres at depths greaterthan 75 km elsewhere in the region. They are almost entirely absent in the Banggai-Sula plate.


characterized by contiguous shallow marine plat-forms, but was certainly not a continuous deepbathymetric feature at this time. In west Sulawesicarbonate deposition terminated by the end ofthe Middle Miocene (van Leeuwen & Muhardjo2005).

Rapid Pliocene uplift

The clastic sediments of the Bongka Formationrecord localized rapid uplift and erosion of theEast Arm in the latest Miocene to Pliocene(Surono & Sukarna 1986; Hall & Wilson 2000),instigating the development of the high (in places.3 km) present-day topography. In some cases,the sudden influx of clastic material may havebeen directly responsible for the reduction of car-bonate areas from large platforms to isolated pin-nacle reefs (Jablonski et al. 2007).

Uplift has previously been attributed to collisionbetween the Banggai-Sula microcontinent and theEast Arm (e.g. Garrard et al. 1988; Davies 1990;Hall 1996; Calvert 2000; Hall & Wilson 2000; vanLeeuwen & Muhardjo 2005). This interpretationfollowed Hamilton’s (1979) proposal of slivers ofcontinental crust moving west from the Bird’s

Head, with westward subduction implied in frontof them. However, it has also been suggested thatcollision between microcontinental blocks andthe East Arm began earlier, between the latestOligocene and Late Miocene (e.g. Audley-Charles1974; Sukamto & Simandjuntak 1983; Daly et al.1991; Parkinson 1991; Smith & Silver 1991;Bergman et al. 1996; Milsom 2000; Hall 2002;van Leeuwen et al. 2007; Spakman & Hall 2010).If so, collision significantly predated the rapiduplift and erosion of the East Arm during thelatest Miocene to Pliocene, which must have a dif-ferent cause.

Basin subsidence

Seismic surveys (Jablonski et al. 2007) and multi-beam surveys of Gorontalo Bay show present-daywater depths up to 2000 m in the western part ofthe basin and .2700 m in the eastern part (Fig. 8).Sediment thicknesses within these areas may be asgreat as 10 km (Jablonski et al. 2007). There isa bathymetric high area that links the East Armand the Togian Archipelago, with water depthsof between 500 and 1200 m (Fig. 8), which maycontinue across the entire bay to the North Arm.

Fig. 8. Detailed bathymetry of Gorontalo Bay, modified from Jablonski et al. (2007). Topography based on SRTM(Shuttle Radar Topographic Mission) data (courtesy of NASA, NGA & USGS).


This feature appears to have a broadly NW–SE trend.

Seismic data have been used to suggest that thebasin formed in a predominantly extensional tec-tonic environment dominated by east–west trendingextensional faults (Jablonski et al. 2007). It wasinterpreted to have formed in the Eocene as afailed rift arm (Jablonski et al. 2007). We infer amuch younger, Pliocene age of formation of thedeep basin.

We interpret deposits of the Bongka Formation(Celebes Molasse) observed in the Togian Islandsand the East Arm as distal and proximal equivalentsof a Pliocene alluvial fan building out from the EastArm. Seismic data reveal thick (up to 2 secondsTWT (two-way travel time)) north–south trendinglobes of sediment that we infer to be submergedparts of this fan (Fig. 9). Prograding fan deltadeposits of similar age are also interpreted fromelsewhere in the basin (Jablonski et al. 2007).These observations imply that basin subsidence(from close to sea level to present-day waterdepths of 500 to 1500 m) occurred after depositionof the fan. The age of the Celebes Molasse in theEast Arm therefore provides a maximum, latestMiocene to Pliocene age (e.g. Surono & Sukarna

1996; Hall & Wilson 2000) for inception ofthe basin.

Cause of subsidence and uplift

The broadly contemporaneous nature of basin subsi-dence and uplift and erosion at the flanks suggeststhat these two processes are inherently linked.Together, the rapid latest Miocene to Plioceneuplift (c. 3 km) and subsidence (.2 km) in andaround Gorontalo Bay has produced an exceptionaltotal elevation contrast of more than 5 km in lessthan 6 Ma. The thickness of sediment in thecentral part of the bay (up to 10 km) suggest muchgreater differential movements.

The North Sulawesi subduction zone probablydeveloped in the last 5 Ma (Silver et al. 1983a;Surmont et al. 1994). We suggest that palaeomag-netic data (Surmont et al. 1994), seismic data(Silver et al. 1983a; Jablonski et al. 2007) andplate tectonic modelling (Silver et al. 1983b;Hall 1996, 2002) indicate that the region has beenin extension since the Early Pliocene, with theNorth Arm moving away from the East Arm.We interpret Global Positioning System (GPS)measurements of present-day motions (Walpersdorf

Fig. 9. Thickness of the sedimentary fill in Gorontalo Bay, modified from Jablonski et al. (2007). Thickness is based ontwo-way travel-time in seconds (TWT s) between water bottom and basement isochron (Jablonski et al. 2007).Topography based on SRTM (Shuttle Radar Topographic Mission) data (courtesy of NASA, NGA & USGS).


et al. 1998; Vigny et al. 2002; Socquet et al. 2006) toindicate that this extension continues today. There-fore one possible cause of subsidence is extensionof the upper plate that was driven by rollback ofthe subduction hinge at the North SulawesiTrench. However, the extremely rapid rates andlarge amounts of uplift and subsidence in theregion suggest that significant flow of lower crust,from beneath the basin towards topographicallyelevated areas, may also have contributed (Hall2010).

Young volcanism

The rocks of the Lonsio Formation and BentengIntrusives record young volcanism in GorontaloBay during the late Neogene. Although the TogianIslands are in the right position for a subduction-related volcanic arc ahead of a westward-movingBanggai-Sula microcontinent, volcanism does notappear to be subduction related. Such volcanicactivity should have preceded the East Sulawesi–Banggai-Sula microcontinent collision. Theyoungest age suggested for this is end Miocene(c. 5 Ma) but the dates we have for the TogianIslands and Poh Head volcanic activity are Plioceneor younger. The composition of the volcanic rocks isnot typical of most subduction-related volcanism.The Benteng Intrusives are extremely rich in potass-ium (Fig. 5; Table 3), they are shoshonites using thescheme of Rickwood (1989). Earthquake hypocen-tres beneath Una-Una volcano (Fig. 7b) show thatvolcanism is unrelated to subduction beneath theNorth Arm, being much further west and muchshallower than hypocentres related to the down-going slab.

High-K compositions are characteristic of smalldegrees of partial melting of anomalous (metasoma-tized or enriched) material in the upper mantle (e.g.Wilson 1989). We infer a similar origin for theBenteng Intrusives and suggest that rapid exten-sional thinning of the crust beneath Gorontalo Baycaused the upper mantle to rise, decompress andmelt. The resulting K-rich melts were intrudedinto the crust as a series of shallow level stocksand dykes. Present-day high-K volcanism at Una-Una suggests that volcanism has evolved to a rela-tively less K-rich chemistry, possibly reflectingincreased amounts of partial melting, and hasmoved WNW over time.

The tuffaceous rocks of the Lonsio Formationrepresent the products of extrusive volcanism,reworked during deposition in a shallow marineenvironment during the latest Miocene and EarlyPliocene (N19). They are significantly older (asmuch as 3 million years) than the Benteng Intru-sives, and appear to be derived from a different –or unknown volcanic centre.

Post-Pliocene tectonics

Tuffaceous rocks of the Lonsio Formation are alsoknown from the East Arm (Rusmana et al. 1982,1993; this study), around 150 km SE of the TogianIslands. Following Simandjuntak (1986), wesuggest that Poh Head has been offset to the SEalong the Balantak Fault. Based on satelliteimages, field observations and seismic data, weinterpret this structure as a steeply dipping, right-lateral, strike-slip fault that can be traced offshoreto the east, where it terminates in a zone of dextraltranspression (Watkinson et al. 2011). To the westof Poh Head the position of the fault is not known,but it may bend to the north, possibly linking tothe fault that we infer between the islands ofWalea Kodi and Walea Bahi. The distribution andages of the volcanic rocks in the Togian Islandsand Poh Head could therefore be explained by post-depositional dextral faulting, or by westwardmigration of the volcanic centre with time.

Conclusions

We interpret Gorontalo Bay to be underlain by acomposite basement comprising several differenttectonostratigraphic provinces. The western end ofGorontalo Bay is underlain by continental crustadded to the eastern margin of Sundaland in themid Cretaceous. The central part of the bay, includ-ing the Togian Islands, is underlain by oceanic base-ment of the Northern Sulawesi Province. It ispossible that the area south of the Togian Islandshas continental crust at depth, with a thrust contactbeneath the Northern Sulawesi volcanic basementand East Arm ophiolite, as suggested by oil seepsthrough the ophiolite on land.

In the Miocene, Gorontalo Bay was an area ofextensive carbonate deposition, characterized bycontiguous shallow marine carbonate platforms. Itwas not a significant, continuous, deep bathymetricfeature in the Miocene. Instead, broadly contem-poraneous flank uplift and basin subsidence give amaximum latest Miocene to Pliocene age for theinception of the deep basin.

Volcanism in the Togian Islands is unrelatedto subduction that preceded collision of theBanggai-Sula microcontinent. Instead, it recordsrapid extension of the crust in the Pliocene andPlio-Pleistocene, causing the underlying mantle torise, decompress and melt. We interpret GPS obser-vations (Socquet et al. 2006) to indicate extension iscontinuing today and is probably the cause of vol-canism at Una-Una. Volcanic activity has migratedwest towards Una-Una during the Pleistocene anddeposits of the Pliocene volcanic episode mayhave been offset by dextral strike-slip displacementalong the Balantak Fault.


Rapid subsidence associated with crustal thin-ning was driven by rollback of the subductionhinge at the North Sulawesi Trench. The unusualcharacter of volcanism in the Togian Islands is notdue to subduction but reflects crustal thinning andextension. The extreme rates of uplift and sub-sidence observed in and around Gorontalo Bay(producing an elevation contrast of .5 km) suggestflow of lower crust may also have contributed.

The industrial member companies of the SE Asia ResearchGroup Consortium provided financial support for our work.The authors thank Benjamin Sapiie and Alfend Rudyawan(Institute Teknologi Bandung) for facilitating our work inIndonesia. M.A. Forster acknowledges the support of anAustralian Research Fellowship provided by the Austra-lian Research Council (ARC) associated with the Discov-ery grants DP0877274, and additional support from theResearch School of Earth Sciences at The AustralianNational University. eArgon software written by GordonLister. We thank Theo van Leeuwen and Moyra Wilsonfor their reviews of the manuscript.

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basement character and basin formation in gorontalo bay, sulawesi, indonesia

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