BACKGROUND

Physical Features and Gross Structure of the Plateau
The Ontong Java Plateau covers an area >1.5 x 106 km² (roughly the size of Alaska or six times that of the United Kingdom) and consists of two parts: the main or high plateau in the west and north and the eastern lobe or salient in the east and south (Fig. 1; Kroenke, 1972). The plateau surface rises to depths of about 1700 m below sea level (mbsl) in the central region of the high plateau but generally lies at water depths of between 2 and 3 km. The plateau is bounded by the Lyra Basin to the northwest, the East Mariana Basin to the north, the Nauru Basin to the northeast, and the Ellice Basin to the southeast. Along its southern and southwestern boundaries, the plateau has collided with the Solomon Islands arc and now sits at the junction of the Pacific and Australian plates. Much of the high plateau's surface is relatively smooth, although several large seamounts have been built on it. In many areas, the basement crust is covered with pelagic sediments >1 km thick. The eastern lobe consists of a large but unnamed northern ridge and the Stewart Arch, which are separated by the ~300-km-wide Stewart Basin; at its southeastern end, the Stewart Basin merges with the Ellice Basin (Kroenke and Mahoney, 1996). Physiography around the margins of the plateau is complicated. In the north and northeast, numerous horst-and-graben structures appear to predate much of the sediment cover (e.g., Kroenke, 1972; Berger et al., 1992). Faulting and deformation along the Ontong Java Plateau's southern and southwestern margins are associated with the plateau's collision with the Solomon arc (e.g., Petterson et al., 1997). An extensive fold belt, the Malaita Anticlinorium, embraces the island of Malaita and the northern half of Santa Isabel.

Crustal thickness on much of the high plateau is considerable, even in comparison to other plateaus. Seismic and combined seismic and gravity evidence indicates that crustal thickness is generally in the 25- to 37-km range (e.g., Furumoto et al., 1976; Hussong et al., 1979; Miura et al., 1996; Gladczenko et al., 1997; Richardson et al., 2000). Over much of the high plateau, the depth to the top of Layer 3A (i.e., to the base of the seismically defined upper crust) is 10-16 km (see review of Neal et al., 1997). Lower crustal seismic-wave velocities suggest a granulite-grade gabbroic lower crust, whereas sub-Moho P-wave velocities of 8.4-8.6 km/s detected in the northwest and southwest portions of the plateau may indicate the presence of eclogite at depth (Saunders et al., 1996; Neal et al., 1997). The maximum extent of Ontong Java-related volcanism may go well beyond the plateau proper, as the Early Cretaceous lava flows filling the Nauru Basin and similar flows in the East Mariana and Pigafetta basins to the north appear likely to be related closely to the plateau (e.g., Castillo et al., 1994; Neal et al., 1997; Gladczenko et al., 1997).

Tectonic Setting and Age of Emplacement
The original plate-tectonic setting of the Ontong Java Plateau is open to some question because well-defined magnetic lineations have not been found on the plateau. However, block-faulting structures along the eastern margin of the high plateau, interpreted as roughly north northeast-trending fracture zones, led to proposals that the plateau formed at a west northwest-trending ridge (Hussong et al., 1979) and possibly at a triple junction (Winterer, 1976; Hilde et al., 1977). Preliminary isotopic study of Ontong Java basement lava flows suggested a hot-spot connection and that the plateau may have formed at a ridge-centered or near-ridge hot spot (Mahoney, 1987). Subsequent geochemical work indicated that plateau basement lavas were formed from a hot-spot-type mantle source by large percentages of partial melting (estimated at 15%-30%), consistent with the plateau having formed on relatively thin and young lithosphere (Mahoney et al., 1993; Tejada et al., 1996; Neal et al., 1997). From bathymetry and satellite derived gravity fabric, Winterer and Nakanishi (1995) inferred that a north-northeast-trending spreading axis ran through the plateau, whereas Neal et al. (1997) argued that the north northeast-trending fabric represents fracture-zone orientation. M-series magnetic lineations adjacent to the plateau in the Nauru and Lyra basins run east-northeast to west-southwest. Coffin and Gahagan (1995) reviewed the available geophysical evidence and concluded that it weakly favors emplacement of most of the plateau in an off-ridge location.

Richards et al. (1991), Tarduno et al. (1991), and Mahoney and Spencer (1991) all favored the starting-plume head of the Louisville hot spot (now at ~50°S) as the source of the Ontong Java Plateau. However, 0- to 70-Ma lava flows dredged from sites along the Louisville Ridge, the plume-tail seamount chain formed by the hot spot, are isotopically distinct from Ontong Java basalt (Mahoney et al., 1993). Moreover, a recent plate reconstruction suggests that the plateau formed 10°-15° north of this hot spot's current location (Neal et al., 1997).

As noted above, the plume-head model predicts that plateaus are emplaced in massive eruptive events lasting only a few million years or less. Surprisingly, however, ⁴⁰Ar-³⁹Ar ages of Ontong Java Plateau lava flows in the Solomon Islands and the pre-Leg 192 drill sites revealed a sharply bimodal distribution (Fig. 3), with ages of 122 ± 3 Ma and 90 ± 4 Ma (total ranges). Thus, most of the plateau may have formed in two relatively brief episodes (Mahoney et al., 1993; Tejada et al., 1996; 2000; Parkinson et al., 1996). Because sampling over the plateau's huge area was very limited, the relative importance of these two episodes remained unclear. However, Tejada et al. (1996) argued that the 122-Ma event was significantly larger than the 90-Ma event. On the basis of abundant 90-Ma lavas (and some dikes) in Santa Isabel and Cenomanian-Coniacian ash layers at DSDP Site 288, they suggested that the 90-Ma episode may have been focused on the eastern salient; shortly thereafter, 90-Ma basalts were also found to be abundant on San Cristobal (Birkhold-VanDyke et al., 1996). An alternative possibility, however, was that further sampling (and dating) could show that eruptions on the plateau actually occurred over a span of 30 m.y. or more (e.g., Tejada et al., 1996; Birkhold-VanDyke et al., 1996; Ito and Clift, 1998).

Between 124 and 100 Ma, the plateau appears to have been positioned close to the Pacific-plate Euler pole, so that it would have moved little relative to the inferred hot-spot source (see Neal et al., 1997). At ~100 Ma, plate motion changed from a northwestward to a more northward trajectory, which continued until ~85 Ma. At ~90 Ma the southeastern corner of the plateau may have been situated rather close to the 122-Ma position of the central high plateau. Following the 90-Ma eruptive episode, rifting and seafloor spreading may have occurred for several million years within the plateau's eastern salient, forming the Stewart Basin in conjunction with spreading in the Ellice Basin to the east (Kroenke and Mahoney, 1996; Neal et al., 1997). An ⁴⁰Ar-³⁹Ar age of 83 Ma was measured by Duncan (1985) for an ocean ridge-type basalt from the eastern Ellice Basin.

Relatively localized Tertiary volcanism is recorded in Malaita in the 44-Ma alkalic lavas of the Maramasike Formation (Tejada et al., 1996; Petterson et al., 1997). Malaita is also peppered with small intrusions of 34-Ma alnöites (e.g., Davis, 1977; Nixon and Neal, 1987). In San Cristobal, a sequence of tholeiitic basalts with ages of ~61 and ~36 Ma overlie the ~90-Ma basalt and are compositionally similar to it (Birkhold-VanDyke et al., 1996). The causes of these later magmatic events are uncertain.

After a long period of northward and northwestward motion, the Ontong Java Plateau collided with the old Solomon arc during the early Neogene, initially in a diachronous "soft docking" without significant deformation. Following a reversal of subduction direction, the intense deformation of the Malaita Anticlinorium occurred in the late Miocene through Pliocene (see Petterson et al., 1997). The bulk of the plateau appears to be more or less unsubductible (Cloos, 1993; Abbott and Mooney, 1995), but the post-Miocene removal of a portion of the lower Ontong Java Plateau between Santa Isabel and San Cristobal is evident from recent seismic surveys (Mann et al., 1996).

Results from Previous Sampling of Cretaceous Igneous Basement
Ontong Java Plateau basement at all previously drilled sites and in the Solomon islands of Malaita, Santa Isabel, Ulawa, Ramos, and San Cristobal consists of pillowed or massive flows of basalt averaging ~10 m in thickness. Dikes are rare in the island exposures, and, hence, the eruptive vents for most of the lava flows may be rather distant. All of the basalt flows appear to have been emplaced well below sea level and are overlain by bathyal or abyssal pelagic marine sediments (see Neal et al., 1997, and references therein). However, all of the locations studied prior to Leg 192, except Site 289, were at the margins of the plateau; thus it remained possible that the central regions of the high plateau and eastern lobe were formed under shallow-marine or even subaerial conditions. Basement of the 122-Ma age group comprises lava flows from Sites 289 and 807, Malaita, Ramos, and part of Santa Isabel, whereas the 90-Ma flows are found at Site 803, in Santa Isabel, and in San Cristobal (Fig. 3) (Mahoney et al., 1993; Tejada et al., 1996, 2000; Parkinson et al., 1996; Birkhold-VanDyke et al., 1996). Also, volcanic ash layers of late Cenomanian to Coniacian age (i.e., in roughly the ~95- to 87-Ma range) are present at DSDP Site 288 (which did not reach basement) on the southern edge of the high plateau (Andrews, Packham, et al., 1975). A glass-shard-rich interval in Aptian limestone at Site 288 and several late Aptian ash layers above basement at Site 289 (Andrews, Packham, et al., 1975) may indicate fairly prolonged shallow or subaerial volcanism in some areas on the crest of the plateau following eruptions at 122 Ma (early Aptian).

Lava flows drilled at all pre-Leg 192 sites are composed of unmetamorphosed, moderately evolved, low-K tholeiite (Fig. 4). They have relatively flat primitive mantle-normalized incompatible element patterns (intermediate between those of normal ocean-ridge basalt and most oceanic island or continental tholeiites) (Fig. 5) and a narrow range of ocean-island-like Nd-Sr-Pb isotopic ratios (Fig. 6). Major and trace element modeling indicates that the basalt represents high degree partial melts (Mahoney et al., 1993; Tejada et al., 1996; Neal et al., 1997). Two geochemically and stratigraphically distinct groups of 122-Ma lava flows are apparent in the thick basement section on Malaita (Tejada et al., 1996, 2000) and in the much thinner one at Site 807 (Mahoney et al., 1993). The upper 46 m of lava flows at Site 807 (Unit A) (1) is isotopically and chemically closely similar to those of the Singgalo Formation, which comprises the upper 750 m of flows in central Malaita, and (2) tentatively has been correlated with them by Tejada et al. (2000). The lower basalt units at Site 807 (Units C-G) and the single flow encountered at Site 289 resemble the flows forming the lower 2.7 km of the volcanic pile on Malaita, termed the Kwaimbaita Formation, with which they have been correlated. The 90-Ma lava flows of Site 803, Santa Isabel, and most of those of San Cristobal are isotopically similar to the 122-Ma Kwaimbaita Formation basalt. Thus, an isotopically ocean-island-type mantle source containing (at least) two distinct components was involved in magma generation at the northern and southern margins of the plateau at 122 Ma. Further, the mantle source of most 90-Ma lava flows was similar to that of the Kwaimbaita Formation, the stratigraphically lower of the two 122-Ma basalt groups.