Several research programs in the last decade have significantly improved our understanding of the regional geological and geophysical setting of rifting into the Papuan Peninsula:
Much of the Papuan Peninsula is 1-3 km above sea level and is underlain by a crust 25-50 km thick (Finlayson et al., 1976). Major orogenic thickening of the crust occurred following the northeast subduction and partial accretion of a thick sequence of dominantly Cretaceous to Eocene strata beneath a late Paleocene-early Eocene island arc that includes the Papuan Ultramafic Belt (PUB), Milne Basic Complex, and Cape Vogel boninites (Davies and Jaques, 1984; Davies et al., 1984; Rogerson et al., 1987, 1993). Collision of the Australian (Papuan) continental margin plateau caused subduction to cease and uplifted the accretionary complex (Owen Stanley Metamorphics) by the early Miocene (Rogerson et al., 1987). Metabasites in the Emo metamorphics and Suckling-Dayman massif (the latter with a cover of Maastrichtian micrites) have been exhumed from 7 to 12 kb (25-35 km) and may represent slivers of the subducted Cretaceous oceanic crust (Unit 4B in Fig. F2; Davies, 1980; Worthing, 1988).
The pre-Miocene geology of the islands on the Pocklington and Woodlark Rises is similar to that of the Papuan Peninsula, with Owen Stanley Metamorphics in the south (Misima, Tagula, and Rossel Islands; Fig. F1) and Milne Basic Complex outcrops on Woodlark Island (Davies and Smith, 1971; Ashley and Flood, 1981). Likewise, upper Paleocene volcanics are present at the base of the Nubiam-1 well (Stewart et al., 1986; Fig. F1), west of the Trobriand Islands. The metamorphic core complexes on the D'Entrecasteaux Islands have a core of Owen Stanley Metamorphics and a cover of metamorphosed ultramafics, unmetamorphosed gabbros, and lower Miocene and younger volcanogenic sediments and limestone (Davies and Warren, 1988; Fig. F2).
Superimposed on this Paleogene basement is widespread middle Miocene-Holocene arc magmatism (Fig. F3; Smith and Milsom, 1984) associated with southwest subduction of the Solomon Sea Basin at the Trobriand Trough. The main volcanic arc is calc-alkaline, whereas minor volcanism on the outer forearc high (Lusancay and Woodlark Islands) is high-K (shoshonitic). Historical arc volcanism and a deforming accretionary prism are compatible with present-day slow subduction at the Trobriand Trough (Hamilton, 1979; Davies and Jaques, 1984; Cooper and Taylor, 1987; Lock et al., 1987; Pegler et al., 1995), although this remains controversial given the small number of intermediate-depth earthquakes beneath the region (Abers and Roecker, 1991) and the lack of 10Be in the lavas of Mount Lamington (Gill et al., 1993). The geochemistry of the volcanics reveals melting and mixing of at least three magma sources: (1) subduction-modified mantle supplied the calc-alkaline arc volcanism; (2) this mantle, with the addition of partial melts from upwelling aesthenosphere, produced the comenditic (transitional basalt-peralkaline rhyolites) series around Dawson Strait (between Fergusson and Normanby Islands); and (3) contamination by lower crust of Australian affinity formed minor high-K trachytes (see Smith, 1976; Hegner and Smith, 1992; Stolz et al., 1993; and references therein).
The Cape Vogel Basin (Fig. F4) is the western part of the Neogene Trobriand forearc basin that parallels the Trobriand Trough to the south of an outer forearc high (with an associated 150-200 mGal free-air gravity anomaly) topped by the Lusancay-Trobriand-Woodlark Islands (Figs. F1, F2, F4). The stratigraphy is known from onshore exposures and the Kukuia wells on the Cape Vogel Peninsula, and especially from the offshore wells (Goodenough-1 and Nubiam-1; Figs. F1, F4, F5) drilled by Amoco and partners in 1973 following 2 yr of regional basin studies, including the collection of 2250 km of seismic reflection data (Figs. F6, F7; Tjhin, 1976; Pinchin and Bembrick, 1985; Stewart et al., 1986; and Francis et al., 1987). It is characterized by early(?)-middle Miocene subsidence, volcanism, and deep-marine sedimentation; late Miocene uplift and erosion (1-2 km) of the margins; and Pliocene coarse clastic (from uplift of the Papuan Peninsula and D'Entrecasteaux Islands to the south) and Quaternary carbonate, shallow-water sedimentation during broad subsidence (Tjhin, 1976; Stewart et al., 1986; Francis et al., 1987; Davies and Warren, 1988).
The basin experienced late Miocene inversion in the northwest, but its center continues to subside in the southeast (Figs. F4, F5, F6; Pinchin and Bembrick, 1985). Several key unconformities are regionally correlatable on reflection seismic lines within the basin and were penetrated by the exploration wells. Of particular importance to Leg 180 results is unconformity A and the overlying Ruaba unit. The Ruaba unit is a distinctive facies characterized by sandstones, shales, and claystones with occasional lignite and conglomerate, and it is interpreted to have been deposited in a paralic environment, including nonmarine deltas to mid-neritic prodeltas (Stewart et al., 1986; Francis et al., 1987). The top of the unit is dated as NN10/NN11 at Goodenough-1, and there, as well as at Nubiam-1, the unit is younger than the underlying Nubiam shale whose upper section belongs to biozones NN10 and N16 (Fig. F5). Thus, the regional unconformity A is definitely younger than 9.63 Ma (the base of Zone NN10) and possibly younger than 8.6 Ma (the top of Zone NN10) (Berggren et al., 1995). That the forearc basin filled to near sea level in the late Miocene is inferred from the seismic sequences that prograde northward across the basin beneath unconformity A to a depocenter just south of the outer forearc high (e.g., Fig. F7).
The Trobriand Trough, the outer forearc structural and gravity high, and the Trobriand forearc basin terminate near Woodlark Island. Farther east, the Woodlark and Pocklington Rises were not a Pliocene-Pleistocene arc-forearc system. Rather, the northern edge of the eastern Woodlark Rise was a transform margin. Seismicity and sidescan data indicate that it is still an active right-lateral fault (Fig. F1). Thus, the Woodlark Basin east of 153.5ºE did not originate as a backarc basin in that the eastern Woodlark and Pocklington Rises were not active island arcs (Weissel et al., 1982). Nevertheless, the locus of present rifting (see "Pliocene-Pleistocene Rifting") does, and did since the latest Miocene bisect an active continental volcanic arc with inherited crustal asymmetry, having a Neogene forearc basin to the north and a Paleogene accretionary/collision complex to the south.
Shallow seismicity, with extensional and strike-slip focal mechanisms having northerly tension axes, is concentrated between Fergusson and Woodlark Islands and extends westward into the Papuan Peninsula at 9ºS-10ºS to about 148ºE (see Fig. F2 in the "Leg 180 Summary" chapter; Weissel et al., 1982; Abers, 1991; Abers et al., 1997). Pliocene-Pleistocene extension has produced three flooded grabens (Mullins Harbour, Milne Bay, and Goodenough Bay) on the eastern extremity of the Papuan Peninsula with associated rift-flank subaerial uplift to over 500 m (Smith and Simpson, 1972). Metamorphic core complexes on the D'Entrecasteaux Islands and in the Suckling-Dayman massif on the Papuan Peninsula were also exhumed during the Pliocene-Pleistocene (Davies, 1980; Davies and Warren, 1988; Hill, 1990). The best structural studies, geothermometry, geobarometry, and age dating of these complexes have been conducted on Goodenough and Fergusson Islands (Davies and Warren, 1992; Hill et al., 1992, 1995; Hill and Baldwin, 1993; Baldwin et al., 1993; Lister and Baldwin, 1993; Hill, 1994; Baldwin and Ireland, 1995). There, normal movement along a 0.3- to 1.5-km-thick ductile mylonitic shear zone resulted in the uplift of deep metamorphic rocks and the juxtaposition of unmetamorphosed cover rocks. Granodioritic intrusion then focused uplift on several domes, offset by strike-slip faults. The ductile shear zones were brecciated and truncated by brittle faults late in their history. These dome structures are juxtaposed along strike with regions of significantly less unroofing, such as the low-grade (greenschist) eastern halves of both Normanby and Misima islands.
The Woodlark Basin has been opening since 6 Ma (Chron 3An.1n), at which time the westernmost spreading segment was located east of 157ºE (Taylor, 1987; Taylor and Exon, 1987; Goodliffe, 1998; Taylor et al., 1999). The pole of early opening was at 9.3ºS, 147ºE, with a rotation rate of 4.234º/m.y., and the spreading center migrated westward through time (Goodliffe, 1998; Taylor et al., 1999). The spreading center grew discontinuously, stepping across Simbo Transform (156.5ºE) at ~3.6 Ma and Moresby Transform (154.2ºE) at ~2 Ma, to reach 151.8ºE by 0.7 Ma. The 500-km-long spreading axis reoriented synchronously ~80 ka, producing the present discordance of the spreading axes with the Brunhes Chron boundary, following a change in the pole of opening (recorded by the fracture zones) at 0.52 Ma to 12ºS, 144ºE, 2.437º/m.y. (Goodliffe et al., 1997; Goodliffe, 1998; Taylor et al., 1999). This pole is located within the error ellipse of the position of the current Australia-Woodlark Euler pole (10.8ºS, 145.2ºE, 1.86º/m.y.) determined geodetically by Tregoning et al. (1998) but predicts opening rates ~30% faster over the last 0.52 m.y. than at present.
The sidescan and geophysical data show that the rifting-to-spreading transition involves both nucleation of discrete spreading cells and organized ridge propagation (Taylor et al., 1995). Two ridge propagation events into the margin at 153ºE formed continental slivers surrounded by oceanic crust (Fig. F1). Spreading is about to propagate into this margin again. Rifting of the conjugate margins continues for ~1 m.y. after spreading has separated them. Extension does not immediately localize to the ridge axis, as shown by the present overlap between spreading and seismogenic margin faulting, and by inwardly curved seafloor fabric and magnetic anomalies that require nonrigid margin reconstructions. The continental breakup proceeds by successive phases of rifting localization, spreading center nucleation, spreading center propagation (sometimes stalling), and then a jump across an offset to the next site of localized rifting (Taylor et al., 1999). As a result of spreading propagation and, most recently, synchronous reorientation, there is no simple relation between seafloor spreading segmentation and margin rift segmentation. The major spreading center offsets all began as overlapping and/or propagating nontransform offsets and many continue in that mode. The Moresby and 155º10'E transforms initiated by crosscutting previous rift structures to link spreading segments that had nucleated in, and/or propagated into, offset continental rifts (Taylor et al., 1999). They did not evolve from transfer/transform faults or aulocogens in the rifted margins. There is no evidence for thermal uplift or igneous underplating of the Moresby Transform continental margins as spreading axes pass by them, including where spreading segments abut the transform continent/ocean boundary segments today.
Seismic reflection data indicate a very sharp (<5 km wide at the surface) transition from rifted crust to oceanic crust (Taylor et al., 1999). There are no dipping reflector sequences indicative of excessive lava production and high degrees of mantle partial melting, but there are small volcanoes a few kilometers in diameter that often erupt along margin faults. Indeed, the initial seafloor spreading lavas indicate low degrees of partial melting: basalts from the youngest spreading segment, just east of Moresby Seamount, have Na8 = 3.1 (Binns and Whitford, 1987; Danyushevsky et al., 1993; Dril, 1997). Other young axial lavas include Fe-Ti basalts, and low- and high-Si andesites with evidence for both mantle heterogeneity and crustal contamination.
The rifting region just ahead of the apex of spreading has been imaged by several seismic reflection and swath bathymetry surveys (Mutter et al., 1993, 1996; Goodliffe et al., 1993; "Geophysical Data and Processing" in the "Data Report: Marine Geophysical Surveys of the Woodlark Basin Region" chapter; Taylor et al., 1996, in press). Continental breakup is focused currently on an asymmetric rift basin bounded on the south by a low-angle (27º ± 3º), curvi-planar, normal detachment that is imaged over the full depth extent of the seismogenic zone (3-9 km; Fig. F4 in the "Leg 180 Summary" chapter; Taylor et al., 1999). The detachment forms the north slope of Moresby Seamount. The rift basin has >2 km of sedimentary fill beneath ~3 km of water and is bounded on the north by an antithetic normal fault that dips ~45ºS. The margin further north, including a prerift (Trobriand forearc) sedimentary basin and basement sequence, is gently flexed down to the south and unconformably onlapped by synrift sediments. The seismic stratigraphy can be reasonably jump-correlated to that in the Trobriand Basin and is interpreted to be a Pliocene-Pleistocene synrift sequence lying unconformably above a Miocene forearc basin sequence on Paleogene volcanic (and metamorphic?) basement (see Figs. F3, F5; both in the "Leg 180 Summary" chapter). To the south of Moresby Seamount, high-relief rotated fault blocks are commonly overlain by only minor ponded sediments.
Shallow (2-10 km) normal and strike slip faults, with northerly T-axes, bound the north side of the rifting-spreading transition (see Fig. F2 in the "Leg 180 Summary" chapter; Abers, 1991; Taylor et al., 1995; Abers et al., 1997). All of the earthquake hypocenters occur within or north of the rift graben, and there are no major extensional structures north of the graben-bounding antithetic fault. Without local seismometers, the teleseismicity has not been definitively associated with the low-angle reflector, but no more likely candidate structure is observed. Furthermore, the seismic stratigraphy cannot be matched without recent faulting on the low-angle reflector.
Two dredges of the northern flank of Moresby Seamount recovered metabasic greenschists, metagabbro, pelitic schist, and minor siliceous phyllite and microgranite (i.e., similar material to the low-grade [greenschist] metamorphics on eastern Normanby and Misima Islands), not the core complex amphibolite metamorphics as on Goodenough, Fergusson, and northwest Normanby (H. Craig, pers. comm., 1986; Binns et al., 1987; J. Hill, pers. comm., 1996). In contrast, a dredge from 541 to 1211 m (~0.7-1.6 s two-way traveltime [TWT]) on the upper southern flank of the Moresby Seamount recovered mainly late Pliocene (Zone N21 = 1.6-3.35 Ma) clastic sedimentary rocks of equivalent facies to the Awaitapu Formation of the Trobriand region in the Cape Vogel Basin (Francis et al., 1987). Benthic foraminifers indicate sediment deposition in water depths of 340-800 m (J. Resig, pers. comm., 1995). These rocks are equivalent to those that we infer lie near the base of the synrift cover sequence in the rift basin and perhaps on the northern margin.
A simplified cross section consistent with the seismic and dredge data available precruise is drawn in Figure F3 in the "Leg 180 Summary" chapter, (Taylor et al., 1996). At the end of the Miocene, the Paleogene basement and a forearc basin filled with Miocene sediment were being eroded at or near sea level (Fig. F7). Pliocene rifting formed sediment-filled grabens in the southern orogenically thickened arc province, accompanied by gradual subsidence of the thinner, colder (and therefore stronger) forearc to the north. Stretching during the Quaternary was localized on a low-angle normal fault. The antithetic hanging-wall faults accommodated little additional extension. East of the Trobriand reef front at 151º10'E (Fig. F1) the northern margin flexed down southward and was onlapped by sediments delivered in part via submarine channels incising northward (Fig. F7 in the "Leg 180 Summary" chapter). We infer that the reef front formed on the unloaded footwall of a north-south tear fault bounding the down-flexed margin to the west. Recently, continued extension on the low-angle normal fault variably collapsed the hanging-wall graben.
This interpretation predicts ~12 km of heave on the low-angle fault. This compares with 130-145 km of total extension in the 4 m.y. before spreading at this longitude, calculated from the pole of opening derived from seafloor spreading magnetic anomalies (Taylor et al., 1996). If this opening is extrapolated from 4 Ma back to the beginning of spreading at ~6 Ma, then 220 km of total extension is predicted at this longitude (Taylor et al., 1999). Given only minor extension of the northern, flexed margin, we infer that the locus of current extension must be the northernmost of a series of similar structures that extended weak crust to the south, forming the block-faulted Pocklington Rise. The regional estimates predict that this rugged province of mainly inactive faults accommodated ~200 km of total strain as it collapsed from heights comparable to the 3-km-high Owen Stanley Ranges that form the backbone of the Papuan Peninsula. Given that this southern province is only 200 km wide, then, if these regional models are correct, the total strain equals the width of the southern province. Substantial intrusion of arc magmas and/or exhumation of lower crustal material from beneath the northern margin may be required to explain these calculations.