20 m.y. is:We cored the sedimentary sequence at Site 1183 from 328.0 to 452.7 mbsf and from 752 mbsf to basement at 1130.4 mbsf. Aptian through Miocene sediment at Site 1183 is composed of foraminifers and calcareous nannofossils with variable amounts of chert and volcaniclastic material. Bioturbation is pervasive. Diagenetic changes include progressive pressure-solution lithification, partial silicification and chert formation, and redox-mediated migration (mobilization) of iron and manganese. The depositional record has at least two hiatuses of Late Cretaceous age that appear to be regional in extent.
The first section below summarizes the general features of each lithologic facies or subunit. The second section discusses the main trends in the sedimentation history. The third section reviews different interpretations of postdepositional diagenetic alteration.
The sedimentary succession at Site 1183 is dominated by carbonate facies and was divided into three lithologic units with seven subunits (Fig. F6; Table T3) on the basis of the following criteria:
Stratigraphic divisions in the upper 950 m (middle Eocene-Miocene) generally parallel those used at DSDP Sites 289 and 586 and ODP Sites 803 and 807 on the main Ontong Java Plateau. This stratigraphic framework is also applied to Leg 192 Sites 1185, 1186, and 1187. Our subdivisions of the Cretaceous and Paleogene section at Site 1183 can be recognized in the approximately correlative sediments at DSDP Site 289 and ODP Site 807, where broader divisions were employed. This consistency suggests that these lithologic divisions represent regional trends in paleoceanography, volcaniclastic influxes, and diagenetic processes.
Unit I consists predominantly of foraminifer nannofossil ooze to chalk. Priority given to basement basalt objectives, coupled with good recovery of a sedimentary succession in relatively nearby DSDP Sites 289 and 586, resulted in the decision to core only two portions of this thick unit at Site 1183. The ooze-to-chalk transition, which defines the transition between Subunits IA and IB, is at ~337.6 mbsf. The lowest subunit, Subunit IC, is characterized by volcanic ash beds within the chalk. Similar volcanic ash-rich horizons are in coeval sediments at Sites 289 and 807; therefore, our Subunit IC is a regionally recognizable interval. The contact between Subunits IB and IC was not cored in Hole 1183A.
Hole 1183A was washed down to 328.0 mbsf, and Core 192-1183A-2R contained carbonate with an ooze texture. The ooze is white to bluish white and homogenized by bioturbation. Foraminifers and calcareous nannofossils have approximately equal abundance by volume, and there are lesser amounts of radiolarians, diatoms, and sponge spicules.
Overprinting the bioturbated texture are faint, colored, horizontal streaks at ~5-cm intervals. These color bands cut across burrows and are interpreted as diagenetic Liesegang banding. Darker patches, possibly of disseminated pyrite or ferromanganese oxide staining, are scattered throughout the ooze.
Subunit IB consists of foraminifer nannofossil chalk with minor amounts of siliceous microfossils and sponge spicules. The transition from ooze to chalk is between Cores 192-1183A-2R and 3R at 337.6 mbsf. We cored 115 m of Subunit IB (Cores 192-1183A-3R through 14R), but the underlying 300 m was drilled without coring. The sediment below the coring gap is also chalk but contains common volcanic ash-rich beds and is designated lithologic Subunit IC. The coring gap precludes precise designation of the boundary between "pure" chalk Subunit IB and underlying volcaniclastic-rich Subunit IC.
The chalk of Subunit IB ranges from white to bluish white to very light greenish gray. Carbonate ranges from 90 to 95 wt%, and the noncarbonate component is predominantly biogenic silica. The low magnetic susceptibility (Fig. F6) suggests that no significant volcaniclastic materials are present. Sediments are faintly mottled throughout, suggesting pervasive bioturbation. Discrete burrows include Zoophycos and Planolites. Mottling aside, color is relatively constant in Cores 192-1183A-3R through 9R. In contrast, Cores 192-1183A-10R through 14R (414-445 mbsf) display subtle color alternations from white to light green at ~75-cm spacing. Burrows in the upper portion of each color type are filled with sediment of the overlying color. Subunit IB contains abundant, thin, crosscutting color bands and streaks of green, blue, and purplish red (Fig. F7).
The lower half of Core 192-1183A-11R includes a thick (3.4 m), homogeneous interval without obvious burrows or color streaks. This interval is unique in having a uniform texture spanning almost 4 m. It appears to contain elongate white "clasts" or filled burrows (1-2 cm long) with a preferential dip of 30° to 45°, in addition to scattered, equant, white clasts. Shipboard biostratigraphy indicates that this bed overlies a hiatus between the early and the middle Miocene, but sedimentological evidence is inadequate to confirm a redeposition or slump origin. Core 192-1183A-12R has similar, but thinner, intervals that display partial bioturbation.
The common presence of volcaniclastic material distinguishes Subunit IC from the ash-free chalk of Subunit IB. Core 192-1183A-15R, recovered below the 300-m coring gap, contains significant amounts of ash; therefore, the upper contact of this subunit is placed within the coring gap. The contact with the underlying Unit II is placed at the uppermost significant chert at 838.6 mbsf (192-1183A-24R-1, 43 cm). Thus, Subunit IC is at least 86 m thick.
The dominant lithology in Subunit IC is nannofossil foraminifer chalk with small amounts of siliceous microfossils. The carbonate content ranges from 90 to 95 wt%, slightly decreasing downward, with a very low value (49%) recorded from an ash layer (Sample 192-1183A-19R-4, 19 cm; 794.79 mbsf) (Table T4). The lowermost 50 m (below Core 19R at 790 mbsf) is limestone rather than chalk. Thin sections indicate that the average chalk is a packstone containing 50% foraminifers forming a grain-supported fabric with a nannofossil matrix (Fig. F8).
Color varies from white to light greenish gray. Smear slide comparisons suggest that intervals of light greenish gray contain more siliceous microfossils. Pervasive bioturbation includes Zoophycos and Planolites burrows. Chert-filled burrows (e.g., a 4-cm-long, 1-cm-wide occurrence in Section 192-1183A-18R-1, 49 cm) are rare. As in Subunits IA and IB, color bands and laminae of greenish gray and other colors cut across the bioturbated fabric.
Gray layers with volcaniclastic material are intercalated with the chalk and limestone throughout this subunit (Fig. F9). These layers are rich in glass shards. Thicker ash layers are preserved as discrete bands, but thinner beds have been blurred by burrowing and form diffuse zones of mixed chalk and glass shards. Increased magnetic susceptibility and reduced color reflectance indicate that these discrete and diffuse ash layers are more abundant toward the lower part of this subunit (Fig. F6). Greenish intervals containing a relative abundance of siliceous microfossils are commonly seen above ash-rich horizons.
The intermixed chalk-ash bands display a well-developed array of small-scale textural features from wispy flaser (chalk lenses partially outlined by ash-rich seams) through flaser-nodular (chalk lenses embedded within an ash-rich matrix) to clusters of compacted ash-filled burrows. We interpret these flaser textures in Subunit IC and in other subunits as diagenetic features, and they are discussed in the final part of "Postburial Diagenetic Features".
The top of Unit II is marked by the uppermost chert layer (Section 192-1183A-24R-1, 43 cm; 838.6 mbsf). A similar unit boundary was used in both DSDP Site 289 and ODP Site 807. The highest chert layer in Hole 1183A is 1.5 m above the lowest volcanic ash (Section 192-1183A-24R-2, 90 cm; 840.1 mbsf) typical of overlying Subunit IC. We subdivided Unit II into a relatively chert-rich upper interval (Subunit IIA; 838.6-958.3 mbsf) and a relatively chert-poor lower interval containing common zeolite-rich beds (Subunit IIB; 958.3-986.6 mbsf). The base of Unit II is marked by the lowest zeolite-rich bed at 986.6 mbsf (Section 192-1183A-39R-4, 15 cm).
Subunit IIA spans 120 m and coincides approximately with the Eocene. Recovery of this subunit was generally <20%, with some cores containing only 10 cm of rock fragments. Recovery of correlative intervals in DSDP Site 289 and ODP Site 807 was similarly low, thus precluding a reliable characterization of the lithologic succession. For example, estimates from wireline-log velocity data at ODP Site 807 indicated a chert content of only 20%, although the recovered sediments were 50% chert (Kroenke, Berger, Janecek, et al., 1991).
Limestone intervals recovered from Subunit IIA are white with faint burrow mottling. The limestone contains an average of 40% foraminifers with a grain-supported fabric in a matrix of calcareous nannofossils and micrite. The limestone from the upper transition between Subunits IIA and IC contains 50%-60% foraminifers in a grain-supported fabric with little intergranular micrite. Even though chert is common, siliceous microfossils are not preserved in the limestone. Carbonate content is 95-99 wt% (Table T4), and the low magnetic susceptibility (Fig. F6) indicates an absence of volcaniclastic material.
The chert is typically light olive-gray to dark olive, but some pinkish gray chert is present in Cores 192-1183A-33R through 35R (Fig. F10). Chert fragments commonly display white rinds, and incompletely silicified, white limestone is incorporated within the chert. White silicified limestone and porcellanite are also present.
Subunit IIB spans only 28 m and contains numerous zeolite-rich beds. Upper and lower boundaries are the highest and lowest zeolitic chalk horizons at 958.3 mbsf (Section 192-1183A-36R-4, 71 cm) and 986.6 mbsf (Section 39R-4, 15 cm), respectively. Similar zeolite-rich beds are present within the Paleocene limestone at DSDP Site 289 and ODP Site 807.
The dominant lithology in this interval is white to very light gray limestone with variable concentrations of foraminifers. In thin section, the limestone is predominantly foraminifers with a grain-supported fabric in a nannofossil matrix (Fig. F11). The limestone is moderately to commonly bioturbated, especially with Planolites burrows, with less abundant Zoophycos and vertical burrows. Carbonate content is 94-98 wt% (Table T4). In a few intervals, the undersides of burrows are outlined in bluish gray (e.g., interval 192-1183A-37R-3, 135 cm, through 37R-4, 50 cm; Fig. F12). Dark brown chert is a minor lithology within the limestone. The chertification is incomplete, and the chert is spotted with residual limestone and white porcellanite.
Gray 0.5- to 5-cm-wide zones within the limestone are zeolite rich and contain small amounts of glass and traces of biotite. They are less lithified than the limestone and therefore are classified as chalk. The zeolitic chalk zones have diffuse, bioturbated boundaries with the host limestone. Similar zeolite-rich bands were interpreted as altered volcanic ash layers at DSDP Site 289 and ODP Site 807 (Shipboard Scientific Party, 1975; Kroenke, Berger, Janecek, et al., 1991). The center of a typical zone displays a finely laminated texture if thicker than 0.5 cm; otherwise, the center consists of a single seam or group of interwoven seams (Fig. F13). Agglutinated foraminifers, which include feldspar or other noncarbonate grains in their tests, are found in the zeolite-rich zones but not in the host limestone.
A relatively greater abundance of zeolitic chalk beds in the lower part of Subunit IIB is indicated by an increase in average magnetic susceptibility, a darker average color, and a slightly lower carbonate content (Fig. F6). The lowest zeolite-rich bed (interval 192-1183A-39R-4, 14-15 cm) is 1 cm wide, greenish gray, and indistinctly laminated (Fig. F14). No zeolite-rich layers are present in the underlying white limestone of the Maastrichtian; therefore, this lowest zeolite-rich bed was used to mark the base of Subunit IIB.
Unit III spans ~144 m from the contact with the lowest zeolite-rich bed at 986.6 mbsf (Section 192-1183A-39R-4, 15 cm) to the contact with basalt at 1130.4 mbsf (Section 54R-3, 120 cm). The unit is predominantly Cretaceous but includes ~70 cm of lowermost Danian sediments at the top. The upper 102 m of the succession consists of white limestone, and the lower 42 m is mottled grayish and pinkish white limestone with minor clay and volcaniclastic beds. This color difference divides the unit into two subunits, and the subunit contact at 1088.8 mbsf seems to correspond to a major hiatus between Santonian and Campanian strata. Core descriptions indicate that both subunits of Unit III are also present at DSDP Site 289 and ODP Site 807.
Subunit IIIA consists of light-colored foraminifer nannofossil limestone with abundant bioturbation and scattered black stylolites. The subunit spans a 102-m-thick interval from the zeolitic chalk at 986.6 mbsf (Section 192-1183A-39R-4, 15 cm) to a marked color change at 1088.8 mbsf (Section 50R-1, 60 cm). The limestone contains recrystallized foraminifers in decreasing abundance downhole. Carbonate content of the limestone is generally 98-100 wt% (Fig. F6; Table T4). Burrows are dominated by Planolites and vertical types and are filled by white, finer grained limestone. Chert is present as rare replacement nodules or lenses of reddish brown to dark brown (Fig. F15).
The upper half of Subunit IIIA (Core 192-1183A-39R through 44R) consists of white to very light gray limestone with abundant foraminifers. In Sections 192-1183A-39R-4 through 40R-3, the limestone contains bluish burrow mottles, has individual trace fossils up to a centimeter in diameter that penetrate as much as 10 cm, and is a foraminifer limestone with ~40% foraminifers in a grain-supported fabric. In Sections 192-1183A-40R-3 through 44R-CC, the limestone contains ~20%-30% foraminifers. In Sections 192-1183A-40R-3 through 41R-1 and Cores 43R and 44R, the limestone is white, relatively soft, and quite homogeneous with very faint burrow mottling. The limestone in Core 192-1183A-42R and most of Core 41R has a slight greenish tinge, and the burrows are usually parallel to bedding planes.
The lower half of Subunit IIIA (Cores 192-1183A-45R through 49R) displays subtle color alternations at a 30- to 40-cm scale between white and pale yellow limestone (Fig. F16). Bioturbation is especially visible at the color transitions. In thin section, the white limestone contains more foraminifers (10%; Fig. F17) than the pale yellow limestone (<5%).
The Cretaceous/Paleogene boundary is ~70 cm below the top of Subunit IIIA in a piece of white siliceous limestone in interval 192-1183A-39R-4, 87-90 cm. Burrows filled with foraminifers of earliest Danian are present in foraminifer limestone of late Maastrichtian age.
Subunit IIIB consists of ~42 m of varicolored micritic limestone with interbeds of dark ferruginous calcareous claystone and vitric tuff in the lowest 10 m.
There is a complex transition interval in Core 192-1183A-50R, and the subunit boundary was assigned to the highest level of bioturbated reddish limestone at 1088.8 mbsf (Section 50R-1, 60 cm). The overlying 60 cm of Core 50R displays a downward change from white to pinkish tints. The highest gray-mottled limestone that is typical of most of Subunit IIIB is nearly 1.7 m below this level. The intervening 1.7-m interval (interval 50R-1, 60 cm, to 50R-2, 87 cm) was recovered completely and contains three dramatic alternations between bioturbated clayey limestone and 10-cm-thick beds of calcareous claystone. The limestone is intensely pink to reddish brown, whereas the claystone is very dark reddish brown (Fig. F18). The claystone has a high magnetic susceptibility (Fig. F6). Pronounced dark stains extend from the claystone into the adjacent limestone intervals. Biostratigraphy indicates that this 1.7-m transition interval includes early Campanian, Santonian, and late Coniacian strata. The reddish brown bed of late Coniacian age directly overlies a hardground on light pink limestone of latest Albian age (see "Biostratigraphy"). Clasts of this pink Albian limestone are found within the reddish brown Coniacian bed.
Below this transition interval, the Aptian-Albian limestone in Subunit IIIB is characterized by color alternations and mottles of lighter pink and of darker gray, although some intervals are a uniform light yellowish gray. The darker gray intervals seem to have more Chondrites burrows, whereas the lighter pink zones contain more Planolites burrows. Anastomosing pressure-solution seams are present in both colors but seem more prevalent in gray intervals. Color changes are variously sharp to gradational and cut across burrow fillings, pressure-solution seams, and chertification horizons. Patches of lighter pink occur within darker gray intervals and vice versa. Superimposed on the mottles are diffuse black spots.
The carbonate content of the limestone is typically 95 wt%, with no significant difference between the color facies. Foraminifers generally comprise <5% of the limestone, and calcified radiolarians are abundant in the basal 50 cm above the basalt. Magnetic susceptibility displays closely spaced oscillations on a generally elevated background (Fig. F6), indicating that iron oxyhydroxides are a constituent of the abundant fine-grained opaque and brownish semiopaque particles observed in thin sections. The noncarbonate components are concentrated along pressure-solution seams.
A combination of compacted subhorizontal burrows, subhorizontal streaks, discontinuous planar to irregular laminations, and thin elongate lenses imparts a distinctive "woody" texture to most of the Aptian-Albian limestone. This woody texture grades into microflasers and anastomosing pressure-solution seams within the limestone (Fig. F19).
Chert is present as irregular bands and as nodules partially replacing limestone. Red chert is associated with pinkish limestone, and dark brown chert with grayish limestone. Pressure shadows are present adjacent to some chert nodules (e.g., Section 192-1183A-52R-1, 97 cm), indicating that initial silicification preceded the final stages of compaction and lithification.
Two beds of laminated calcareous claystone lie at ~10 m and 4 m above the base of Subunit IIIB. The upper bed (interval 192-1183A-53R-3, 115 cm, to 53R-4, 85 cm) is a 1.2-m-thick, reddish brown to black, ferruginous nannofossil claystone containing scattered horizontal stringers and lenses of packed foraminifer tests and fine sand-size clasts (Fig. F20). The lower bed (Section 192-1183A-54R-1, 50-60 cm) is a very dark brown claystone that was poorly recovered. In sharp contrast to the pervasive burrowing of the adjacent limestone pieces, these claystone beds display virtually no bioturbation. Both beds have high magnetic susceptibility (Fig. F6). Clay- to silt-size, brown, semiopaque particles of goethite and perhaps other iron oxyhydroxides are predominant in smear slides, thin section, and X-ray diffraction (XRD) data (Table T4). Carbonate content is ~40 wt%. The scattered laminations are concentrations of calcite-filled foraminifers and phosphatic fish debris; these components are also dispersed within the clay-rich matrix. Foraminifer tests and fragments are partially replaced and filled by opaques, which appear to be iron oxyhydroxides that may have replaced pyrite or precipitated in place. Planktonic foraminifers are common throughout the claystone, but benthic foraminifers are conspicuously absent.
The basal 5 m of limestone has a variety of primary and secondary colors, including white, gray, yellow, pink, reddish brown, and olive-brown. Darker shades form faint halos around some Chondrites burrows.
The lowest 2 m of Subunit IIIB contains two intervals of olive-gray to reddish brown vitric tuff (Fig. F21). The lowermost of the two intervals is separated from the underlying basalts by a 25-cm-thick limestone bed. The main component of the tuff, identified in thin section, is basaltic ash consisting of partly glassy to tachylitic fragments with abundant plagioclase microlites. Texturally, many of the fragments are similar to aphanitic pillow margins in the underlying basalts. Most fragments are nonvesicular, but some have vesicles or scalloped margins. Altered brown glass shards are also present; most are blocky and nonvesicular, but some are moderately vesicular. The tuff is composed of at least eight normally graded beds, several of which have scoured bases. The uppermost layer grades up through parallel-laminated to cross-laminated beds, indicating deposition by turbidity currents or reworking by currents. There is a thin layer of bioturbated limestone above the fourth depositional layer.
The base of Subunit IIIB is placed at the top of the uppermost basalt unit at 1130.4 mbsf (Section 192-1183A-54R-3, 120 cm), but the sediment/basalt contact was not recovered. The lowermost sediment recovered above the basalt (interval 54R-3, 95-120 cm) is the typical bioturbated nannofossil limestone of Subunit IIIB (Fig. F22).
Basement consists of eight pillow-basalt units with thin limestone interbeds. The basalt is described in detail in "Igneous Petrology," "Alteration," and "Structural Geology". The limestone is partially recrystallized and light yellow to very pale brown. The limestone beds recovered between pillow-basalt flows comprise Piece 2A above hyaloclastite in Section 192-1183A-54R-4 (Fig. F23), two pieces at the top of Core 55R, one piece in Section 55R-2 at 80 cm, and one piece at the top of Section 56R-1.
The sediment is thermally altered ferruginous micrite limestone with foraminifers and radiolarians. Calcite-filled foraminifers and radiolarians are in a micrite matrix that has been partially recrystallized to fine-grained spar and that contains abundant fine-grained opaque and brown semiopaque particles. Sedimentary structures are difficult to distinguish. XRD data show that the limestone fragment in Core 192-1183A-60R includes glauconite.
The Maastrichtian through Pleistocene sedimentation record is quite consistent across the main Ontong Java Plateau (e.g., Berger et al., 1991). The succession is predominantly calcareous sediments composed of foraminifers and calcareous nannofossils. Deposition was punctuated by episodes of relative enrichment in siliceous microfossils and in volcaniclastic material and by major regional hiatuses. Only the Aptian-Campanian stratigraphy displays significant local variability.
Noteworthy aspects of Site 1183 sedimentation history include the following:
The Aptian-Albian section is relatively thick at Site 1183, where it is represented by 42 m of mottled grayish and pinkish white limestone (Subunit IIIB) and by rare pockets of ferruginous micrite limestone within the underlying pillow basalt flows. This subunit at Site 1186 has a similar thickness (see "Lithostratigraphy" in the "Site 1186" chapter). In contrast, only a thin veneer of basal limestone overlain by noncalcareous radiolarian claystone was found at Site 807 on the plateau's northern flank.
Following the last flow of pillow basalt at Site 1183, a minimum of 25 cm of bioturbated limestone was deposited before deposition of at least eight beds of vitric tuff. The minor, moderately vesicular basaltic glass shards in these tuff beds indicate formation by relatively shallow submarine eruptions, whereas the partly glassy basaltic ash that constitutes the dominant component could have been derived from shallow-water to subaerial hydroclastic eruptions or by erosion of a volcanic source somewhere in the summit region of the main Ontong Java Plateau. These beds and possibly a vitric tuff of similar age just above basement basalt at DSDP Site 289 (Andrews, Packham, et al., 1975) are the only evidence that a portion of the main plateau was relatively shallow and possibly subaerial.
The Aptian-Albian limestone is thoroughly bioturbated and light pink (with secondary gray staining, as explained later). These characteristics indicate oxygenated waters at this depth in the Pacific basin. However, there are two important exceptions.
Above the Aptian/Albian boundary interval is a 1.2-m-thick, reddish brown to black, ferruginous nannofossil claystone (interval 192-1183A-53R-3, 115 cm, to 53R-4, 85 cm; Fig. F24). Another, but poorly recovered, very dark brown claystone is ~6 m lower, within the middle Aptian (interval 192-1183A-54R-1, 50-60 cm). The moderate abundance of planktonic foraminifers indicates depositional depths above the foraminifer lysocline.
Some components of these intervals suggest a period of condensed sedimentation. The abundance of phosphatic fish debris and enrichment in fine-grained noncalcareous components suggest a slow rate of sediment accumulation. Iron oxyhydroxides and clay (identified as goethite and nontronite in XRD patterns) form abundant, brown, semiopaque particles within the claystone (Fig. F25). These may be related to red-brown semiopaque objects (RSOs) that are common in pelagic brown clays at deeper Pacific sites (e.g., Yeats, Hart, et al., 1976; Shipboard Scientific Party, 1990). RSOs are considered to be iron oxyhydroxides formed during intense early diagenetic processes accompanying low sedimentation rates (Karpoff, 1989, and references therein). Other opaque particles, which may be pyrite replacements that were later altered, partially fill some foraminifer tests and have shapes similar to radiolarian and foraminifer tests.
Other observations suggest that the Aptian and Albian environment on the plateau during these episodes was unfavorable for bottom life. No bioturbation disrupts the numerous stringers and laminae of concentrated planktonic foraminifers and fish debris (Fig. F26). The laminae of concentrated foraminifers may represent periodic winnowing by bottom currents or surface productivity blooms. Despite the abundance of planktonic foraminifers, there are no benthic foraminifers. Possible causes of unfavorable environments include low oxygen levels on the bottom and/or an inadequate food supply.
Four oceanic anoxic events (OAEs) have been proposed as global or widespread episodes during the Aptian and Albian (e.g., Schlanger and Jenkyns, 1976; Bralower et al., 1993). These events are a mid-early Aptian OAE1a (Selli level), an Aptian/Albian boundary OAE1b (Paquier level), an early late Albian OAE1c, and a latest Albian OAE1d. Organic-rich laminated shale is the typical signature of OAEs in the Atlantic-Tethys seaway. However, except for the OAE1a event, which is recognized in the Mid-Pacific Mountains as three laminated organic carbon-rich intervals (Shipboard Scientific Party, 1981; Bralower et al., 1993), the record of these events in the Pacific basin is patchy. High-resolution biostratigraphic correlation and carbon isotope stratigraphy may establish whether the laminated ferruginous claystone layers at Site 1183 coincide with any of these postulated global OAEs.
An alternative hypothesis is that low sedimentation rates during de-position of the claystone were a result of very low surface productivity (hyperoligotropic). A low flux of planktonic tests and associated organic material to the bottom would limit benthic communities. Under this scenario, the broad summit region of the Ontong Java Plateau became a biological "desert" during part of the Aptian because of reduced circulation of nutrient-rich water.
The Cretaceous sediment succession overlying basalt basement on the main Ontong Java Plateau has been recovered from a total of seven sites during DSDP Leg 30 and ODP Legs 130 and 192. These sites span the lower flank to the summit of the plateau and record different ages both for the termination of calcareous sediment deposition during the Aptian-Albian interval and for the return to carbonate deposition during the Campanian-Maastrichtian interval. We interpret this pattern as the result of a rise and fall of the CCD relative to the depositional surface of the plateau.
We obtained an estimate of relative CCD trends by applying a simple backtracking and subsidence procedure (e.g., Berger and Winterer, 1974; Tucholke and Vogt, 1979) uniformly to all basement sites on the main Ontong Java Plateau to estimate the paleodepth of the sediment surface through time (Fig. F27). This procedure incorporates progressive sediment accumulation with time as the underlying basement undergoes thermal subsidence and sediment loading.
The compensation for sediment loading assumes a local Airy isostatic depression of the mantle:
in which we used the values of Tucholke and Vogt (1979) for mantle density (3.3 g/cm3), average carbonate sediment density (1.8 g/cm3), and seawater density (1.03 g/cm3). The resulting loading adjustment is ~0.66 of the sediment thickness, implying for each 1 m of average sediment accumulation, the basement is depressed by 0.33 m and the depositional surface rises by 0.66 m.
The backstripped depths of the current sediment-basement interface, adjusted for removal of the total sediment loading, are as follows:
These backstripped basement depths are slightly deeper than the array shown in Figure F37 in the "Leg Summary" chapter because here we apply a slightly lower average density for the typical chalk sediment.
We estimated the initial
depth and progressive deepening of basement by backtracking, following a
generalized version of the global depth and heat flow (GDH1) model of Stein and
Stein (1993). In this empirical fit to global seafloor depth data, the
lithosphere behaves as the cold upper boundary layer of a cooling half-space for
the first 20 m.y., followed by an asymptotic approach to observed maximum depths
for old crust. The equation for the total subsidence of crust
20 m.y. is:
where d(t) = total subsidence (in meters), and t = elapsed time (in m.y.) since the initial thermal age.
Application of this equation, assuming that the initial thermal age is the same as the age of the erupted basalts, would imply 2968 m of total subsidence in the past 122 m.y. This would place the uppermost basalt at Hole 1183A at an initial altitude of 405 m above sea level. This result is incompatible with the submarine, nonvesicular nature of the pillow basalts. Similar discrepancies in subsidence models have been noted for sites on the Kerguelen Plateau (Coffin, 1992), Mid-Pacific guyot carbonate platforms (Röhl and Ogg, 1996) and other volcanic edifices constructed on older crust. Detrick and Crough (1978) proposed that the subsidence of a volcanic edifice emplaced on older oceanic crust behaves as if the initial thermal age is intermediate between the edifice age and the age of the underlying and partially cooled oceanic lithosphere. This concept fits the accumulation history of Aptian-Albian carbonate platforms of the Mid-Pacific Mountains and MIT guyot (Röhl and Ogg, 1996) where subsidence rates (rates of shallow-water carbonate accumulation) were constrained by coring during Legs 143 and 144.
An initial thermal age ~2.5 m.y. older than the assumed 122-Ma age of the uppermost basalt flows on the main Ontong Java Plateau would mean that the lithosphere would have subsided by ~600 m (using d(t) = 365t1/2 for the initial 2.5-m.y. period) before the emplacement of the plateau (see Stein and Stein, 1993). This would satisfy both of the following:
This depth zone is consistent with the calculated 1100-m bathymetric offset between the compensated basement depths of Holes 807C and 1183A, as computed above (see also Fig. F37 in the "Leg Summary" chapter). However, the implied 200-m paleodepth of the final eruptive phase of Site 1183 conflicts with the nonvesicular texture of the basalts, which would generally indicate an eruptive water depth >800 mbsl (Moore and Schilling, 1973).
Implicit in applying this model to the Ontong Java Plateau is the simplifying assumption that no later underplating or reheating episodes (e.g., tectonic model of Ito and Clift, 1998) caused deviations from a thermal subsidence curve typical of an edifice emplaced on older oceanic crust.
The accumulated sediment thickness at each successive 5-m.y. interval was added to the computed subsidence of basement and isostatically compensated, as explained above. The resulting trends of estimated sediment surface paleodepth display a rapid initial deepening, followed by a semiconstant or even slightly shallowing depth as the rapid sedimentation rates in the Cenozoic balance thermal subsidence (Fig. F27).
We assumed that the major hiatuses in the Cretaceous carbonate record were produced when the seafloor at each of the sites was below the CCD and connected these age-depth intersections to derive an estimated history of the Cretaceous and Paleogene CCD. Although the general trends produced by this model may be indicative of the CCD history, the absolute estimates of paleodepths may be in error.
According to this subsidence model and the carbonate successions at the different sites, the CCD on the Ontong Java Plateau rose during the Aptian through Albian, then descended during the Campanian and Maastrichtian. Contributing to the apparent relative rise of the CCD during the Aptian-Albian was the thermal subsidence of the main Ontong Java Plateau. A similar major rise of the CCD during the Aptian-Albian, followed by rapid descent during the Campanian-Maastrichtian, is recorded in the North Atlantic and Indian Ocean basins (e.g., van Andel, 1975; Tucholke and Vogt, 1979; Thierstein, 1979). The CCD oscillations during the Cretaceous may be related to global paleoceanographic and biological trends (e.g., Berger, 1979; Sclater et al., 1979; Barrera and Savin, 1999; Frank and Arthur, 1999; MacLeod and Huber, 2001). Shore-based studies of carbon isotope data coupled with characterization of the planktonic assemblages might provide clues to changing surface-water conditions and circulation patterns on the Ontong Java Plateau.
In Hole 1183A, burrows of earliest Danian (planktonic foraminifer Zone P0) penetrating chalk of middle-late Maastrichtian age represent the Cretaceous/Paleogene boundary (Sample 192-1183A-39R-4, 87-90 cm). A depositional hiatus or end-Maastrichtian erosion episode truncated the final 2 m.y. of the Cretaceous record (see "Biostratigraphy").
An abundance of zeolite-rich horizons characterizes Subunit IIB, which spans only 30 m. Similar zeolite-rich bands within the Paleocene limestone at DSDP Site 289 and ODP Site 807 have been interpreted as altered volcanic ash layers (Kroenke et al., 1993). The lowest zeolite-rich layer is 70 cm above the Cretaceous/Paleogene boundary at interval 192-1183A-39R-4, 81-85 cm (987.1 mbsf). A similar coincidence of a lithified volcanic ash just above the Cretaceous/Paleogene boundary is present at Site 807 (interval 130-807C-54R-3, 135-136 cm).
A potential source of the ash beds is the Paleogene volcanic arc that developed behind the Papua-New Caledonia-Norfolk Ridge subduction zone ~2000 km southwest of the Ontong Java Plateau (Kroenke, 1984, 1996). Activity along this subduction system may have begun when the Pacific plate changed direction at ~65 Ma (polarity Chron C29) and continued until ~40 Ma (Duncan and Clague, 1985; Kroenke et al., 1993). Alternatively, the volcaniclastic material may have had a closer source, perhaps the formation of one or more of the seamounts on top of the Ontong Java Plateau, such as the edifice below the immense Ontong Java atoll (see Fig. F1). Because the glass shards in the original ash have been pervasively altered to zeolite, no information is available from shard morphology or color to help distinguish between these possible sources.
Eocene sedimentation in the world's oceans is characterized by a high accumulation rate of biogenic silica (from diatoms and radiolarians) and abundant chert (e.g., Lancelot, 1971; Jansa et al., 1979; Riech and von Rad, 1979; Emery and Uchupi, 1984). Increased siliceous volcanic activity during the Eocene is a common hypothesis invoked to explain the surge in silica deposition and chert formation (e.g., Gibson and Towe, 1971; Mattson and Pessagno, 1971; Tucholke and Mountain, 1979; Mélières et al., 1981; Emery and Uchupi, 1984). Volcanism is postulated to stimulate siliceous plankton productivity by supplying the nutrients Si and Fe and to enhance preservation by increasing the silica saturation in deep waters. However, this postulated relationship between volcanic activity and elevated silica and chert abundance in sediments is contradictory to the trends recorded on the Ontong Java Plateau.
Hole 1183A penetrated 120 m of interbedded Eocene chert and limestone (Subunit IIA). The chert-rich interval of the Eocene is ash poor but is sandwiched between ash-rich, chert-poor Paleocene and Oligocene sediments. This large-scale inverse relationship suggests that increased input of volcanic material is not a major factor contributing to the formation of chert. We favor an alternative hypothesis that silica is enriched in the oceans during periods of warm humid climate, when intensified chemical weathering leaches silica from continental rocks (e.g., Millot, 1964). The early and middle Eocene was the warmest period in the Cenozoic (e.g., review by Crowley and North, 1991) and coincides with this anomalous episode of worldwide siliceous sedimentation.
The top of the chert capping Unit II at all sites on the Ontong Java Plateau is the seismic reflector terminating the "Ontong Java Series" (Berger et al., 1991). This reflector is analogous to the Eocene seismic Horizon Ac capping the chert-rich Bermuda Rise Formation of the North Atlantic basin (Tucholke, 1979). A global cooling trend and associated shift of diatom productivity to circum-Antarctic seas and coastal upwelling zones, plus a reduction in global volcanic activity, contributed to silica "starvation" after 40 Ma in tropical seas (reviewed by Berger et al., 1991). However, on the Ontong Java Plateau, regional volcanic activity increased, rather than decreased, at the top of the chert beds. Therefore, we suggest that regional volcanic input to the Ontong Java Plateau inhibited, rather than aided, chertification.
This inhibition of chert formation by volcanism seems to occur both at the large scale of the lithologic subunits and at the individual bed scale. Within the Paleocene limestone, individual beds of volcaniclastic zeolites and localized chertification are never associated. One contributing factor is the process of chert formation during diagenesis. Zeolite and clay from altered volcanic material dispersed in biogenic sediment are sources for easily exchangeable cations which inhibit chert precipitation (e.g., Millot, 1964; Lancelot, 1971).
The second major Cenozoic episode of ash-rich sediment deposition began in the late Eocene and continued into the Miocene. Chalk that is rich in volcaniclastic material was recovered from similar age intervals at Site 1183, DSDP Site 289 (e.g., Shipboard Scientific Party, 1975), and ODP Site 807 (e.g., Kroenke, Berger, Janecek, et al., 1991, fig. 19 on p. 393). Kroenke et al. (1993) did not include this episode in their tephrachronology for the Ontong Java Plateau because the published Site 289 and Site 807 core description sheets did not give details on the numbers of Oligocene ash horizons.
The first appearance of the volcanic ash beds at ~40 Ma is similar in age to both the basaltic volcaniclastic rocks at Site 1184 (see "Biostratigraphy" in the "Site 1184" chapter) to the southeast of the main plateau and the initiation of Melanesian arc volcanism, when the southern side of the Ontong Java Plateau was converging obliquely with the Melanesian Trench (Kroenke et al., 1993). The Melanesian volcanic arc remained active throughout the Oligocene until the completion of docking of the Ontong Java Plateau with the Melanesian arc at ~23 Ma (Kroenke, 1984, 1996).
Light greenish gray intervals with a relative enrichment of siliceous microfossils are found within the white nannofossil chalk of Subunit IC, where they commonly overlie ash-rich horizons. This association may indicate either a preferential preservation of biogenic silica caused by ash-influenced interstitial water chemistry or an enhanced productivity of siliceous plankton. The influence of volcanic ash may be superimposed on a subtle primary cyclicity in the Oligocene. Portions of the Miocene Subunit IB display alternations from white to light green at ~75-cm spacing, which may correspond to oscillations in paleoceanographic conditions caused by Milankovitch cycles and/or episodic minor influxes of fine volcanic ash.
Following deposition and mixing by bioturbation, the carbonate ooze underwent compaction, progressive pressure-solution lithification, partial silicification, and late-stage redox reactions of iron and manganese compounds. These diagenetic effects created a variety of features. Studies of similar features at DSDP Site 289 and from ODP Leg 130 resulted in different interpretations.
The approximate succession of selected diagenetic effects with time and depth is as follows:
These processes progress with depth within Site 1183 and overlap in their effects (e.g., Fig. F28).
The ooze-to-chalk transition (Subunits IB/IA boundary) in Hole 1183A is based subjectively on the finger-press test, and we assigned this transition to ~337.6 mbsf. Other sites at 2- to 3-km water depths on the Ontong Java Plateau have the ooze-to-chalk transition at a similar depth within the sediment column. In deeper water, the ooze-to-chalk transition is at progressively shallower levels within the sediment column (Berger et al., 1991). For example, at Site 803 in 3410 m of water, this transition is at 217 mbsf. However, many factors besides water depth are involved in this early lithification process (van der Lingen and Packham, 1975).
Overprinting the bioturbated texture of the white chalk and ooze throughout the Oligocene and Miocene (lithologic Unit I) are abundant (~5-cm intervals), thin (~1 mm), faint horizontal bands and streaks of green, blue, or purplish red (Fig. F29). Similar Liesegang color laminae and streaks within coeval chalk at ODP Sites 803 and 807 were interpreted in two ways.
The first interpretation is that these color laminae are a diagenetic enhancement of subtle primary compositional differences (especially organic carbon content) by preferential redox reactions of iron and manganese (e.g., Kroenke, Berger, Janecek, et al., 1991; Berger et al., 1991). The second interpretation is that the color bands develop as minor concentrations of volcanic ash are transformed during diagenesis (Lind et al., 1993). This interpretation was based primarily on comparison of appearance and frequency to color laminae in Oligocene-Pleistocene chalk of Lord Howe Rise. Those laminae contain smectite and have a temporal pattern resembling other records of Cenozoic volcanism in the southwest Pacific (Gardner et al., 1986). In the ODP holes on the Ontong Java Plateau, the greenish laminae are associated with Fe- and Al-bearing silicates, whereas the purplish laminae are caused by finely disseminated iron sulfide (Lind et al., 1993). However, direct evidence of a volcanic component, such as glass or thicker ash horizons adjacent to color bands, was not demonstrated in any of the previous studies.
Silicification of deep-sea sediments is a progressive process involving phase transitions from biogenic-skeletal opal-A through opal-CT to quartz. The transformation rates depend on host lithology, concentration of siliceous microfossils, elapsed time, and burial depth (e.g., Lancelot, 1971; Hein et al., 1981; Behl and Smith, 1992). The first stage is preferential silicification of burrows; the amount and types of burrowing may control the initial loci of formation of chert nodules and bands (Hein et al., 1981). The shallowest chertification in Hole 1183A occurs as rare chert-filled burrows in the Oligocene (Subunit IC). In addition, the relative concentrations of siliceous microfossils may be affected by periodic oscillations in surface fertility driven by Milankovitch climate cycles (e.g., Ogg et al., 1992).
Formation of chert within host limestone typically requires between 30 and 50 m.y., according to Behl and Smith (1992), but this interpretation may be partially biased by the global pattern of chert formation during the Eocene (35-50 Ma). Chertification of limestone is completed when all biogenic silica has dissolved from zones within the host siliceous carbonate sediment and has reprecipitated as bands or nodules (Fig. F30). Within Eocene Subunit IIA, all siliceous microfossils have been dissolved from the surrounding carbonate. This striking contrast between pure carbonate and silicified limestone was also noted at ODP Site 807.
The limestone from Hole 1183A displays a suite of small-scale features, including wispy flaser (chalk lenses partially outlined by volcaniclastic clay seams), flaser-nodular (chalk lenses embedded within a volcaniclastic clay matrix; Fig. F31), and a woody texture (subhorizontal streaks, thin lenses, burrows, and laminations). Many intervals closely resemble microflaser and nodular limestone structures within the English Chalk and other Upper Cretaceous limestone (e.g., Kennedy and Garrison, 1975; Scholle et al., 1983). These structures are most evident in bands that are enriched in volcanic ash, zeolite, or another noncarbonate component.
Two contrasting interpretations have been proposed for such flaser textures. DSDP Leg 30 sedimentologists inferred bottom currents with a superimposed tidal component. Flaser textures are a common feature of storms and tide-influenced ripples, and tides may influence sedimentation to great depths (Lonsdale et al., 1972a, 1972b). Therefore, microflasers in the limestone implied that the surface of the Ontong Java Plateau was affected by deep-sea tidal currents or periodic major storms during the Aptian-Oligocene (e.g., Klein, 1975; Shipboard Scientific Party, 1975). However, at Site 1183, we observed no evidence of current laminations within the foraminifer chalk lenses to support a ripple origin.
In contrast, the Leg 130 sedimentologists postulated a recrystallization and pressure-solution origin for the microflasers at Site 807 (Kroenke, Berger, Janecek, et al., 1991). Earlier, Bathurst (1987) concluded that stylolites are typical of pressure-solution affecting carbonate having <8%-10% clay, whereas greater proportions of clay will result in dissolution seams. However, even though the limestone throughout the flaser-bearing facies from Hole 1183A (and at Site 807) contains <10% clay, stylolite horizons are rare, yet bands of microflasers surrounded by anastomosing seams are common. Therefore, Lind (1993) proposed that microstylolites and flasers develop if the clay and other noncarbonate minerals are finely disseminated, whereas clay concentrated in a narrow zone tends to give rise to a stylolite.
We propose a modified diagenetic process in which heterogeneity, not uniformity, plays an important role in creating the array of diagenetic features at Site 1183 and other sites. We observed a continuum of textures in Hole 1183A, from differential cementation and compaction of burrow fillings to chalk lenses isolated by anastomosing and braided clay seams, to "floating" chalk nodules embedded within volcaniclastic claystone. If bioturbation has redistributed material from volcaniclastic horizons into the adjacent chalks, then there would be an irregular patchy distribution of the noncarbonate minerals. During pressure solution, carbonate dissolves from relatively clay-rich zones, then reprecipitates in more carbonate-rich zones. The insoluble residue left behind forms wavy clay-rich streaks that act as conduits for continued pressure solution of the adjacent carbonate (Fig. F32). The process continues with accumulating overburden and time and results in the wispy flaser structures within these bands. Pressure solution may eventually concentrate the insoluble ash or zeolites into a central seam or anastomosing set of seams surrounded by microflaser textures in the adjacent chalk, as found within some bands in the Paleocene (Subunit IIB). Under further differential compaction and pressure-solution mediation of lithification of the bioturbated sediment, the microflaser and anastomosing clay-seam features grade into the woody texture of the Aptian-Albian limestone (Fig. F33). A similar pressure-solution process was proposed for the woody texture in the lower Cretaceous clay-rich radiolarian sediments at ODP Site 801 (Shipboard Scientific Party, 1990; Ogg et al., 1992, pp. 591-592).
The Aptian-Albian limestone has pink and gray mottling. The color differences are partially a diagenetic feature because the color changes cut across burrow fillings, clay seams, and chertification horizons; are variously sharp to gradational; and are also present as patches within the opposite color facies. The most striking dark color fronts extend from volcaniclastic claystone bands in Core 192-1183A-50R into the adjacent limestone intervals. These color stains are probably a redox mobilization of manganese or iron from the claystone or altered volcaniclastic material. Similar mottling features in Lower Cretaceous sediment in the Pigafetta Basin were ascribed to mobilization and differential staining by iron-manganese oxyhydroxides (Lancelot, Larson, Fisher, et al., 1990; Ogg et al., 1992). Migrating Fe-Mn compounds, either from the volcaniclastic beds or from the underlying plateau basalt, probably also caused the variety of colors (reddish brown, olive, yellow, etc.) in the basal 5 m of limestone in Hole 1183A.
