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Long- and Short-Term Warming in the Paleogene
The most recent episode of moderate to extreme global warming occurred during the late Paleocene and early Eocene. In detail, this warming was comprised of several events that occurred on different time scales. Over a 4-m.y. period (58-54 Ma), mean global temperatures increased gradually, reaching a peak in the early Eocene that was sustained for roughly 2 m.y. (Fig. 4) (Zachos et al., 1994). Superimposed on this long-term trend was a brief but more extreme episode of warming at the Paleocene/Eocene boundary (~55.5 Ma) (Kennett and Stott, 1991; Bralower et al., 1995; Thomas and Shackleton, 1996), known as the PETM. This event lasted for ~200 k.y. (Röhl et al., 2000).

The exact cause(s) of the long- and short-term warm episodes remains enigmatic. Several pieces of geochemical evidence point toward greenhouse forcing. These include changes in the mean ocean 13C and alkalinity (Shackleton, 1986; Kennett and Stott, 1991; Zachos et al., 1993; Thomas and Shackleton, 1996). Here, we outline specific questions concerning the nature and causes of these warm episodes that can be addressed by Leg 198 drilling.

The Paleocene-Eocene Thermal Maximum (PETM)
In terms of the rate and degree of warming, the PETM is unprecedented in Earth's history. The deep-sea and high-latitude oceans warmed by 4° and 8°C, respectively. The warming, in turn, led to profound changes in precipitation and continental weathering patterns (Gibson et al., 1993; Robert and Kennett, 1994). The climatic changes also affected biota on a global scale, triggering rapid turnover of benthic and planktonic organisms in the ocean (e.g., Thomas, 1990; Kelly et al., 1996) and a sudden radiation of mammals on land (Koch et al., 1992).

The carbon isotopic composition of the ocean decreased by 3‰ to 4‰ coeval with the warming event, suggesting a massive perturbation to the global carbon cycle (Fig. 5) (Kennett and Stott, 1991; Bains et al., 1999). The large magnitude and rate (~-3‰ to -4‰ per 5 k.y.) of the carbon isotope excursion (CIE) is consistent with a sudden injection of a large volume of isotopically depleted carbon into the ocean/atmosphere system. Dickens et al. (1995, 1997) suggested that the largest source of depleted carbon was the vast reservoir of methane clathrates stored in continental shelf sediments. They hypothesized that gradual warming of deep waters during the late Paleocene would have destablized shelf and slope clathrates, triggering a catastrophic release of CH4. Much of this methane would have quickly converted to CO2, stripping O2 from deep waters and lowering alkalinity, thus contributing to the benthic extinction. Both CO2 and CH4would have immediately contributed to greenhouse warming.

The Leg 198 depth transect will help us address the following questions.

Paleogene Deep Water Circulation
Several investigators have suggested that early Cenozoic global warming would have altered deep-ocean circulation patterns by reducing the density of surface waters in high latitudes (Kennett and Shackleton, 1976; Wright and Miller, 1993; Zachos et al., 1993). This in turn would permit increased downwelling of highly saline but warmer waters in subtropical oceans. Such reversals or switches in circulation probably occurred suddenly rather than gradually. In fact, it has been suggested that a sudden change in intermediate water circulation patterns may have occurred just prior to the PETM, possibly triggering the dissociation of clathrates (Bralower et al., 1997b). There may have been additional abrupt warming intervals in the late Paleocene and early Eocene (Thomas and Zachos, 1999; Thomas et al., 2000). These "hyperthermals" were characterized by changes in the assemblage composition of benthic foraminifers corresponding to negative shifts in planktonic and benthic foraminiferal 18O and 13C values. The ultimate cause of the hypothermals may be similar to the late Paleocene Thermal Maximum (LPTM), driven by release of greenhouse gas.

Leg 198 will provide the samples needed to assess regional-global circulation changes during the Paleogene. Major changes in the sources of waters bathing Shatsky might be reflected in the spatial and vertical distribution of carbon isotope ratios in bottom waters as well as in benthic foraminiferal assemblage patterns. Several studies have shown that throughout the late Paleocene and early Eocene, the most negative deep-ocean carbon isotope values were consistently recorded by benthic foraminifers from Shatsky Rise (Miller et al., 1987a; Pak and Miller, 1992; Corfield and Cartlidge, 1992). Such a pattern is similar to that in the modern ocean, implying older nutrient-enriched waters in the Pacific and younger nutrient-depleted waters in the high latitudes. Although Site 577 is discontinuous across the Paleocene/Eocene (P/E) boundary, isotope data from Site 865 on Allison Guyot in the equatorial Pacific suggest a possible reduction, if not reversal, in the 13C gradient between the shallow Pacific and most of the ocean (Bralower et al., 1995). If this were true, it would be consistent with increased production of intermediate waters in low latitudes. In summary, Leg 198 samples will help address the following questions.

Eocene-Oligocene Paleoceanography
The Eocene to Oligocene represents the final transition from a "greenhouse" to an "icehouse" world. Although this transition occurred over a period of 18 m.y., stable isotopic records reveal that much of the cooling occurred over relatively brief intervals in the late early Eocene (~52 Ma) and earliest Oligocene (~33 Ma) (Fig. 4) (e.g., Kennett, 1977; Miller et al., 1987b, 1991; Stott et al., 1990; Zachos et al., 1996). Furthermore, small ephemeral ice sheets were probably present on Antarctica some time after the first event (Browning et al., 1996). The first large permanent ice sheets became established much later, most likely during the early Oligocene event (Zachos et al., 1992a). Current reconstructions of ocean temperature and chemistry for the Eocene and Oligocene, however, are based primarily on pelagic sediments collected in the Atlantic and Indian Oceans (Miller et al., 1987b; Zachos et al., 1992b, 1996). Very few sections suitable for such work have been recovered from the Pacific (Miller and Thomas, 1985; Miller and Fairbanks, 1985). As a consequence, we still lack a robust understanding of how global ocean chemistry or circulation evolved in response to high-latitude cooling and glaciation.

Eocene-Oligocene sediments are present on Shatsky Rise. Unfortunately, early coring efforts on the rise failed to recover continuous and undisturbed sequences (e.g., Site 305), whereas later coring efforts encountered major unconformities (e.g., Site 577). Leg 198 advanced hydraulic piston core (APC) drilling at or near Site 305 should provide a nearly continuous sequence of well preserved Eocene to Oligocene sediment. Recovery of similar sediments at one or two of the deeper Leg 198 sites would provide a vertical depth transect and the first opportunity to reconstruct, in a third dimension, the evolution of ocean chemistry and temperature during this important climatic transition. Leg 198 drilling will allow us to address the following questions.

Mid- and Late Cretaceous Climate
The mid-Cretaceous (Barremian-Turonian [125-85 Ma]) Earth experienced some of the warmest temperatures and lowest thermal gradients of the entire Phanerozoic Eon. This time interval, therefore, represents one of the best ancient approximations of "greenhouse" climate. The Late Cretaceous was characterized by significant global cooling, but available oxygen isotopic records differ on the exact timing of the end of the greenhouse conditions. Records from DSDP Site 511 on the Falkland Plateau, South Atlantic (Fig. 6) (Huber et al., 1995) and the chalk from England (Jenkyns et al., 1994) suggest that peak warmth occurred in the early Turonian, about 90 Ma. Data from Shatsky Rise DSDP sites (e.g., Douglas and Savin, 1975; Savin 1977), however, indicate that peak greenhouse conditions existed in the Albian, ~105 Ma. In addition, these stratigraphies differ on whether peak warming was immediately followed by long-term cooling (English Chalk) or sustained warmth then cooling beginning in the mid Campanian (Site 511 data). Differences between the various records may reflect real latitudinal climatic variations or diagenetic alteration of stable isotope proxies.

There is also significant disparity as to exactly how much cooling occurred in the Late Cretaceous, especially in the tropics. Savin (1977) and D'Hondt and Arthur (1996) concluded that the Maastrichtian was characterized by surprisingly cool tropical SSTs (20°-21°C) based on 18O analyses of planktonic foraminifers, i.e., the "cool tropics paradox" (D'Hondt and Arthur, 1996). Wilson and Opdyke (1996), on the other hand, measured 18O values on rudists recovered from Pacific guyots and concluded that tropical SSTs in the same interval were extremely warm (between 27° and 32°C). The climate history of the Cretaceous is based on a limited number of data points from few sites with little information from the tropics. In fact, Shatsky Rise Site 305 (Douglas and Savin, 1975) is among a handful of low-latitude sites that form the basis of most Cretaceous thermal gradient estimates that are used as inputs in climate models (e.g., Barron and Peterson, 1990).

There is a limited understanding of the evolution of bottom-water circulation in the mid and Late Cretaceous, in particular, how and when the transition from low-latitude (e.g., Brass et al., 1982) to high-latitude (e.g., Zachos et al., 1993) deep-water sources took place. Benthic foraminiferal 18O records are even sparser than those based on planktonic foraminifers, and there are very few benthic data from the entire Pacific. Thus, the role of this giant basin in the evolution of deep waters during the mid and Late Cretaceous is poorly understood.

The long-term cooling of the Late Cretaceous was interrupted by a dramatic event in the mid Maastrichtian, when the source of deep waters appears to have changed abruptly from low- to high-latitude sources (e.g., MacLeod and Huber, 1996; Barrera et al., 1997; Frank and Arthur, 1999). This event appears to have coincided with the extinction of the inoceramid bivalves (MacLeod et al., 1996) and possibly also the rudistid bivalves (Johnson et al., 1996). Growing evidence, however, suggests that this benthic event is distinctly diachronous (MacLeod et al., 1996). The change to high-latitude deep-water sources appears to have been long lived, lasting until the PETM. However, more benthic data are required to accurately characterize Late Cretaceous and Paleocene deep-water properties.

Drilling of relatively shallow burial-depth Upper Cretaceous sections on Shatsky Rise will help address the following questions.

Mid-Cretaceous Oceanic Anoxic Events
The beginning of "greenhouse" climate conditions in the mid Cretaceous (Barremian Turonian) was associated with widespread deposition of organic carbon (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. 7). 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, 1990). 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 Aptian-early Albian (e.g., Weissert, 1989; Bralower et al., 1998). Short-lived negative 13C 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 disocciation of methane hydrates (Jahren and Arens, 1998; Jahren et al., 2001).

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, the Manihiki Plateau, the Mariana Basin, and the accreted oceanic limestone from the Franciscan Complex along the western margin of North America (Sliter, 1984).

Of particular importance to understanding the paleoceanography of OAEs are the changes in oceanic and atmospheric chemistry associated with the onset of black shale deposition in the latest Barremian to earliest Aptian. It was recently documented that the deposition of mid Cretaceous Corg-rich sediments were coincident with a world-wide pulse in ocean crustal production (Fig. 8) (Larson, 1991; Tarduno et al., 1991; Arthur et al., 1991; Erba and Larson, 1991). The initial massive pulse of this volcanic episode was associated with doubling of ocean-crust production rates at about 5 Ma. The pulse peaked in the middle to late Aptian, tapered gradually through the rest of the mid Cretaceous, and dropped significantly at the end of Santonian time. This, the largest volcanic episode possibly in the past 250 m.y. of Earth history, included increased seafloor spreading rates and increased rates of formation of LIP oceanic plateaus, seamount chains, and continental flood basalts. The first large-scale plateau eruptions were in the Pacific Basin where eruptions in the Aptian began forming the Ontong Java and Manihiki Plateaus, the Mid-Pacific Mountains, and Hess Rise, among other features.

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 due to creation of anomalously young and therefore shallow ocean floor (Hays and Pitman, 1973; Schlanger et al., 1981). The signature of this massive volcanic event may be detected in strontium isotope ratios (Jones et al., 1994; Bralower et al., 1997a).

The OAEs had a profound effect on the evolution and extinction of marine nekton, plankton, and benthos (Eicher and Worstell, 1970; Elder, 1987; Leckie, 1987, 1989; Roth, 1987; Bralower, 1988; Premoli-Silva et al., 1989; Erba, 1994). Planktonic foraminifers from each of the Corg-rich episodes show unique patterns of extinction and survival, species diversity, specimen size, morphotype, and subsequent adaptive radiation (e.g., Leckie, 1989; Sliter, 1980). In each case, the survivors were largely small globular forms that resemble modern species that proliferate in areas characterized by increased upwelling. In contrast, the larger, more morphologically complex forms that resemble modern species that live within and below the thermocline suffered extinction or severe reduction in numbers. These faunal changes across the Corg-rich intervals strongly suggest corresponding changes in the chemistry and physical characteristics of the upper water column (e.g., Sliter and Premoli-Silva, 1990; Premoli-Silva and Sliter, 1999).

The Barremian/Aptian boundary interval was a time of major diversification of nannoplankton (e.g., Bralower et al., 1994); however, the causes of this diversification are not understood. The early Aptian OAE was associated with the demise of the nannoconids (Coccioni et al., 1992; Erba, 1994), a group of taxa that existed in rock-forming abundances in the Early Cretaceous. Erba (1994) postulated that the nannoconid demise resulted from wholesale changes in the fertility structure of the oceans, possibly as a result of voluminous plateau-building volcanism.

Variations in productivity, oxygenation, temperature, and density structure have been suggested as potential factors that may have driven evolution and caused extinction in these time periods. However, there is little independent evidence for changes in any of these factors during the OAEs. Clearly, a better understanding in the changes of water column properties during the OAEs will help constrain the driving forces for faunal and floral evolution and extinction that accompany these events.

Recovery of mid-Cretaceous Corg-rich deposits at relatively shallow burial depth from Shatsky Rise will help determine relationships between the conditions that favored the deposition of Corg-rich sediments, climate, and volcanism. The following specific questions will be addressed.

General Early Cretaceous Paleoceanography
Microplankton Evolution
The Early and mid Cretaceous were critical times in the evolution of planktonic foraminifers and calcareous nannoplankton (e.g., Roth, 1987; Leckie, 1989). Nannoplankton underwent dramatic radiations close to the Jurassic/Cretaceous and Barremian/Aptian boundaries (e.g., Bralower et al., 1989, 1994). Both of these events have been documented in the Atlantic and Tethys but not yet from the Pacific. Pacific sites recording these diversification events would help provide an understanding of their causes.

Planktonic foraminifers appear to have evolved in the Bajocian (Middle Jurassic), but their occurrence is sporadic below the Lower Cretaceous. The diversification of this group was, until recently, thought to have occurred in the early Aptian. Coccioni and Premoli-Silva (1994), however, found the evolutionary appearance of a number of taxa far below their previous ranges in the lower Valanginian of the Rio Argos section of Spain. Documentation of this diversification event in other locations and oceanographic settings will help our understanding of its causes.

Shatsky drilling will help us determine: (1) How did the evolution of nannoplankton correlate to changes in ocean thermal structure and circulation? (2) Is there evidence for diversification of planktonic foraminifers in the early Valanginian as in Spain, and if so, did this event correlate with any obvious changes in circulation or climate?

Valanginian Greenhouse Event
A major change in stable carbon isotope ratios of marine carbonates and organic matter has been observed in the Valanginian (e.g., Lini et al., 1992). The event appears to correlate with a major burial event of Corg, an increase in atmospheric CO2, and global warming, perhaps the earliest indications of the Cretaceous "greenhouse" climate (Lini et al., 1992). Increased crustal production rates at this time (e.g., Larson, 1991) suggest that the event may have a volcanic origin. Warming in the Valanginian is at odds with the evidence of Stoll and Schrag (1996) and others for glaciation in this part of the Cretaceous. Recovery of high-quality stratigraphic sections from additional locations will help resolve this issue. Shatsky coring will help us address how the Valanginian carbon isotope record correlates to indicators of climate change and volcanism and whether there is evidence for warming or cooling in this time interval.

Early Cretaceous CCD Fluctuations
The Early and mid Cretaceous were characterized by major changes in the level of the CCD (e.g., Thierstein, 1979; Arthur and Dean, 1986). These changes likely resulted from changes in fertility, sea level, ocean-floor hypsometry, and ocean circulational patterns. One of the most dramatic events occurred in the early Aptian at around the same time as the massive Pacific volcanic event, suggesting that volcanism played a direct role, perhaps through increased pCO2. The few data that exist for the Pacific suggest a different CCD history from the Atlantic (Thierstein, 1979), and more data will help resolve the history of the Pacific CCD. Shatsky coring will help us address: (1) What was the gradient of carbonate dissolution in the mid Cretaceous Pacific Ocean? (2) What was the history of variation in the lysocline and CCD in the Early and mid Cretaceous? (3) Was a major early Aptian CCD shoaling episode observed for the Atlantic Ocean basins characteristic of the global ocean or were the oceans out of phase as the result of the pattern of deep-water aging?

Nature and Age of Shatsky Rise Basement
LIPs such as Ontong Java Plateau and Shatsky Rise were constructed during voluminous magmatic events that took place over geologically brief (<1 m.y.) time intervals (e.g., Duncan and Richards, 1991; Tarduno et al., 1991; Coffin and Eldholm, 1994). These events are thought to be associated with massive thermal anomalies in the mantle known as "superplumes" (Larson, 1991). A likely possibility is that the voluminous phase of superplume activity was associated with the ascent of a plume "head" and that activity declined as the magma source dried up as the lithosphere rode over the plume "tail." One of the major questions concerning the origin of LIPs, such as Shatsky Rise, is whether they formed in a midplate setting or at a divergent boundary, possibly a triple junction, at times of changing plate geometry (e.g., Sager et al., 1988).

Trace element geochemistry of most samples from Shatsky Rise is close to mid-ocean ridge basalt (MORB) (Tatsumi et al., 1998), indicating that they were generated at a divergent boundary, but a few samples have an affinity closer to Polynesian alkalic basalts, suggesting a midplate origin. The latter result is not unexpected, because Shatsky Rise is thought to have formed in the South Pacific (McNutt and Fischer, 1987) near crust with the distinctive Polynesian chemistry. Additional basement coring at Shatsky will help us address whether it had some kind of a hybrid origin or whether there are other explanations for the few anomalous samples.

Although volcanic basement crops out at several localities on Shatsky Rise (Sliter et al., 1990), the only basement samples obtained are from dredges and these are heavily weathered. Ozima et al. (1970) dated volcanic rocks dredged from Shatsky Rise as Tertiary in age. Either these rocks were derived from late-stage volcanism identified in seismic data from the rise or else possibly their pervasive alteration precludes reliable age determination. Maximum estimates for the age of basement on Shatsky Rise can be obtained by adjacent magnetic anomalies (Nakanishi et al., 1989). These ages range from 148 Ma (Late Jurassic polarity Zone CM21) in the southern high to 136 Ma (Berriasian-Valanginian polarity Zone CM14) in the northern high based on the time scale of Gradstein et al. (1994). Fresh basement samples will provide valuable age information.

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