The beginning of "greenhouse" climate conditions in the mid-Cretaceous (BarremianTuronian) was associated with widespread deposition of Corg-rich sediments, informally known as "black shales," in the oceans. These Corg-rich deposits are known to occur primarily at specific stratigraphic horizons, namely, the lower Aptian, the uppermost Aptian to lowermost Albian, the upper Albian and in the upper Cenomanian, close to the Cenomanian/Turonian boundary (e.g., Jenkyns, 1980; Schlanger et al., 1987; Sliter, 1989; Arthur et al., 1990; Bralower et al., 1993) (Fig. F9). Schlanger and Jenkyns (1976) hypothesized that these OAEs resulted from the vertical expansion of oxygen minimum zones linked to transgressive sea-level pulses and the reduced oxygenation of bottom waters. Others have theorized that oxygen depletion and the deposition of Corg-rich sediments instead was the consequence of other paleoceanographic changes such as salinity stratification (Ryan and Cita, 1977; Thierstein and Berger, 1978) and increased flux of Corg from surface productivity or terrestrial sources (e.g., Dean and Gardner, 1982; Parrish and Curtis, 1982; Pedersen and Calvert, 1991).
The deposition of mid-Cretaceous Corg-rich sediments coincided with a worldwide pulse in ocean crustal production (Fig. F10) (Larson, 1991a; Tarduno et al., 1991; Arthur et al., 1991; Erba and Larson, 1991). The release of mantle CO2 from this enormous volcanic episode may have directly caused mid-Cretaceous "greenhouse" warming. The increased preservation and production of organic carbon may have resulted from this warming (e.g., Arthur et al., 1985) combined with increases in nutrients, while sea level rose as the result of the creation of an anomalously young, and therefore shallow ocean floor (Hays and Pitman, 1973; Schlanger et al., 1981).
Regardless of their origin, the burial of Corg-rich sediments enriched in 12C led to significant positive 13C excursions. These have been documented for the Cenomanian/Turonian boundary (Scholle and Arthur, 1980), the early Aptian, and the late Aptianearly Albian (e.g., Weissert, 1989; Bralower et al., 1999). Short-lived negative 13C excursions at the onset of the events may be related to input of mantle CO2 during volcanic events (e.g., Bralower et al., 1994) or to the dissociation of methane hydrates (Jahren and Arens, 1998; Opdyke et al., 1999; Jahren et al., 2001). OAEs are known to be times of rapid turnover among marine biotas as a result of complex changes in habitats (Coccioni et al., 1992; Erba, 1994; Premoli Silva and Sliter, 1999; Premoli Silva et al., 1999; Leckie et al., in press).
Complicating the development of paleoceanographic models are apparent differences in the stratigraphic extent and paleobathymetry of Corg-rich deposits from the Pacific compared to the Atlantic and Tethys Oceans. In the Atlantic and Tethys, Corg-rich deposits occur mostly in basinal settings characterized by major inputs of terrestrial Corg by turbidity currents that led to vertically widespread, long-lived episodes of deep-water anoxia (e.g., Arthur and Premoli-Silva, 1982; Arthur et al., 1984; Stein et al., 1986). Terrestrial Corg-rich deposits in the Atlantic and Tethys occur in intervals besides the OAEs (e.g., Bralower et al., 1993). The record of carbonaceous strata in the Pacific is concentrated in the OAEs, dominated by marine Corg, and almost exclusively restricted to paleobathymetric highs (e.g., Dean et al., 1981; Thiede et al., 1982). However, our understanding of the Pacific record is based on scattered occurrences of carbonaceous strata from Shatsky, Hess, and Magellan Rises, the Mid-Pacific Mountains (MPM), the Manihiki Plateau, the Mariana Basin, and the accreted oceanic limestone from the Franciscan Complex along the western margin of North America (Sliter, 1984).
Recovery of mid-Cretaceous Corg-rich deposits at relatively shallow burial depth from Shatsky Rise will help determine
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