OAEs represent major disruptions to the ocean system defined by massive deposition of organic carbon in marine environments (Schlanger and Jenkyns, 1976; Jenkyns, 1980; Herbert and Fischer, 1986; Arthur et al., 1990). Despite the fundamental role that OAEs are widely hypothesized to have played in the evolution of Earth's climatic and biotic history, very little is really known about the causes and effects of these events. Arguably, between two and six OAEs occurred during the midLate Cretaceous (OAEs 1a1d, 2, and 3) (Jenkyns, 1980; Arthur et al., 1990; Erbacher and Thurow, 1997) (Fig. F5), and these are particularly important because they have left records, not merely in shallow seas but also in the deep oceans.
The 13C records from the Western Interior, English Chalk, and Italian Scaglia appear to confirm the initial designation of OAE 3 for the late Coniacian, but current resolution of Atlantic records is insufficient to determine the existence of additional events in the late TuronianSantonian (Jenkyns, 1980; Jenkyns et al., 1994). Similarly, until recently, comparatively little was known about the Albian OAEs (OAEs 1b1d), but two new studies demonstrate the potential to improve constraints on the origin of different OAEs when diagenetically uncompromised microfossils become available from modern ocean drilling. Data from ODP Site 1049 suggest that pronounced water column stratification instigated OAE 1b (Erbacher et al., 2001), whereas records from nearby Site 1052 indicate that OAE 1d was triggered by the total collapse of upper ocean stratification, intense vertical mixing, and high oceanic productivity (Wilson and Norris, 2001). These antipodal hypotheses for the proximal causes of two OAEs within the same Cretaceous stage emphasize the utility of targeting sections that we know to contain records of multiple OAEs.
The two most prominent mid- to Late Cretaceous black shale events are the late early Aptian "Selli Event" (OAE 1a; ~120 Ma) and the Cenomanian/Turonian boundary ("Bonarelli Event") (OAE 2; ~93.5 Ma) (Fig. F5). Both OAEs 1a and 2 have sedimentary records in all ocean basins (Arthur et al., 1985, 1988, 1990; Bralower et al., 1994; Thurow et al., 1992), and the Aptian event is now known to have been truly cosmopolitan; its sedimentary expression extends even to the extremely shallow waters of mid-Pacific atolls (Jenkyns and Wilson, 1999). These findings and recent improvements to 13C records from land sections (both in outcrop and drill core) and the mid-Cretaceous seawater 87Sr/86Sr curve reveal three important factors concerning the possible origin of OAE 1a (Bralower et al., 1997; Menegatti et al., 1998; Erba et al., 1999; Jahrens et al., 2001; Jones and Jenkyns, 2001):
The global presence of laminated sediments and a variety of geochemical records demonstrate that the response of the carbon cycle during OAE 2 was somehow related to dysoxiceuxinic conditions at the sediment/water interface (e.g., Sinninghe Damsté and Koster, 1998). However, the cause and dimensions of O2 deficiency remain unclear and controversial. The substantial positive 13C excursion of seawater at the time of OAE 2 (Scholle and Arthur, 1980; Schlanger et al., 1987; Jenkyns et al., 1994) is widely attributed to increased global oceanic productivity and rates of Corg burial, but recent treatment of the problem using a simple model of the global carbon cycle indicates that this paradigm view requires more thorough investigation (Kump and Arthur, 1999). The process of sedimentary sequestration of Corg is hypothesized to act as a rapid negative feedback mechanism for global warming via drawdown of atmospheric carbon dioxide (Arthur et al., 1988; Kuypers et al., 1999).
The following scientific questions can be addressed using the sediments recovered from the Demerara Rise transect:What is the history of OAEs in the tropical Atlantic as recorded on Demerara Rise?
Currently, a wide range of hypotheses invoke changes in ocean circulation and/or stratification to explain OAEs, but virtually no reliable geochemical data exist to constrain changes in the basic physical properties (temperature and salinity) of the water masses involved. These competing hypotheses can be tested using 18O, trace element, and biomarker records using well-preserved microfossils and Corg from Demerara Rise.
The tropics are widely viewed as an environment in which physicochemical factors and thus biotic compositions are inherently stable. Yet many low-latitude species have low environmental tolerances, thereby suggesting that relatively small climate changes may result in a substantial biological response (Stanley 1984). The so-called Cretaceous and Paleogene greenhouse was characterized by a series of significant marine and terrestrial biotic turnovers. Most of these events seem to be linked to major changes in Earth's climate (EoceneOligocene transition and P/E boundary), paleoceanography, and/or the geochemical carbon cycle (Cretaceous OAEs and mid-Maastrichtian Event). Many of these events also produced synchronous turnovers in both terrestrial and marine biotas. The causes of most of these turnovers are poorly known because of the absence of expanded sections in the deep sea, where paleontological and isotopic studies can be carried out at high temporal resolution.
Widespread and presumably related isotopic, sedimentological, and paleontological changes are concentrated in the middle portions of the Maastrichtian (e.g., Barrera and Savin, 1999; Frank and Arthur, 1999; MacLeod and Huber, 2001). However, there are currently no established criteria for defining this interval. Some associated changes are graded over millions of years (e.g., high-latitude cooling), others are markedly diachronous (e.g., the last occurrence of bathyal inoceramids), others are not expressed in all areas (e.g., carbon isotopic excursion among benthic foraminifers), and still others are poorly dated (e.g., the collapse of rudist reefs). These uncertainties notwithstanding, in the subtropical North Atlantic and Tethys, the decline in abundance among inoceramids can first be resolved in the upper portion of Chron 31R, and their disappearance (except Tenuipteria) occurs in Chron 31N. This interval coincides with diversification among Tethyan planktonic foraminifers, with increased isotopic gradients among planktonic foraminifers on Blake Nose, an increase in the intensity of bioturbation at ODP Site 1052 and the Basque region, and proposed changes in the North Atlantic CCD. Thus, during Leg 207, mid-Maastrichtian changes are expected within the upper portion of Chron 31, in Zone CC25, and near the first occurrence datum of the planktonic foraminifer Abathomphalus mayaroensis.
The biotic turnovers of the mid-Cretaceous OAEs (OAEs 1b, 1d, and 2) are broadly comparable to one another even if the detailed causal factors are thought to have been different (Leckie, 1987; Erbacher and Thurow, 1997; Premoli Silva et al., 1999). A faunal crisis in nannoconids is well documented in the Aptian (Erba, 1994). Similarly, the early Albian OAE 1b strongly influenced the evolution of both planktonic foraminifers and radiolarians as did the other OAEs. Some events not only influenced planktonic groups but also benthic foraminifers, ammonites, bivalves, and even angiosperms. OAE 2 ranks as the eighth largest mass extinction in Phanerozoic Earth history (Sepkoski, 1986). Extension of the oxygen minimum zone (OMZ) and a rapid eutrophication of the oceans have been linked to extinction and a subsequent radiation of plankton and benthos alike (e.g., Hart, 1980; Caron and Homewood, 1983; Kaiho eand Hasegawa, 1994; Erbacher et al., 1996; Leckie, 1989). Documentation of fossilized photosynthetic green sulfur bacteria from Cenomanian/Turonian boundary black shales points to the existence of euxinic conditions in the protoNorth Atlantic (Sinninghe Damsté and Koster, 1998; Kuypers et al., 2002). But interpretations differ as to whether the associated Corg burial was caused by OMZ expansion and Corg preservation or to enhanced productivity (Sinninghe Damsté and Koster, 1998).
The 13C excursions around three events (OAE 1b, 1d, and 2) have been interpreted in terms of increases in oceanic productivity, and this mechanism has been invoked to explain wide-scale carbonate platform drowning events in the Tethyan realm (Erbacher and Thurow, 1997; Weissert et al., 1998). On the other hand, results from the Pacific suggest that high tropical SSTs rather than eutrophication were responsible for platform drowning (Wilson et al., 1998; Jenkyns and Wilson, 1999). Cretaceous OAEs and extreme climates of the Paleogene (K/T and P/E boundaries and middlelate Eocene refrigeration) led to profound changes in plankton and benthos within the oceans (Thomas, 1998; Aubry, 1998).
The following questions concerning Cretaceous and Paleogene biotic turnover will be addressed using microfauna recovered from Demerara Rise:
A wide range of biotic observations suggest that substantially higher mid-latitude and polar temperatures relative to today prevailed during certain intervals of Earth history (e.g., mid-Cretaceous and early Paleogene), with tropical temperatures throughout the past ~150 m.y. probably at least as warm as today (Adams et al., 1990; Crowley and North, 1991). The 18O paleothermometry in deep-sea foraminiferal calcite supports the existence of these past "warm climates" (Fig. F6). These data show that deep and surface waters in the Cretaceous Antarctic during these intervals were significantly warmer than today (e.g., ~15°C for SSTs) (Huber et al., 1995). In contrast, broadly contemporaneous SSTs estimated in this way for the tropics are generally no warmer and sometimes much cooler (a minimum of ~12°18°C) than today (Shackleton, 1984; Barrera, 1994; D'Hondt and Arthur, 1996). Such cool tropical SSTs contradict not only biotic observations but also basic theories of tropical ocean-atmosphere dynamics (Crowley, 1991). Attempts to simulate Cretaceous climates using numerical general circulation models have consistently demonstrated that (1) high levels of atmospheric CO2 (four times present amounts) are needed to explain the warm polar SSTs derived from 18O paleothermometry and (2) this level of greenhouse forcing also yields increases in tropical SSTs beyond those indicated by 18O data sets (e.g., Manabe and Bryan, 1985; Barron, 1995; Bush and Philander, 1997; Poulsen et al., 1999; Otto-Bliesner et al., 2002). Explanations for the apparent paradox of the "cool tropical greenhouse" fall into two basic categories: (1) models of past warm climates fail to account adequately for polar ocean and/or atmospheric heat transport and (2) tropical 18O SST estimates are misleading.
Many artifacts plague existing records of tropical SST, including their extremely low resolution misidentification of true surface-dwelling species of foraminifers and the susceptibility of epipelagically secreted calcite to early diagenetic alteration in favor of artificially low SSTs (Douglas and Savin, 1975; Killingley, 1983; Schrag et al., 1995). Recent studies demonstrate that, under the correct geological circumstances, ancient carbonates (even highly metastable minerals) can be remarkably well preserved and yield 18O SSTs for the tropics that are significantly warmer than those provided by diagenetically suspect material (Wilson and Opdyke, 1996; Norris and Wilson, 1998; Wilson and Norris, 2001; Wilson et al., 2002; Norris et al., 2002; Pearson et al., 2001). These studies show that foraminifers recovered from sections with clay-rich lithologies and/or shallow burial depths exhibit a distinctive "glassy" taphonomy similar to fauna recovered from modern-day sediment traps. This Cretaceous material includes epipelagic fauna that yield tropical 18O SSTs that match or, in some cases, exceed those measured today, thereby suggesting a thermal response to greenhouse forcing in the tropics.
The concept of a greenhouse mid- to Late Cretaceous Period is well supported by models of Earth's tectonic history. These models indicate that the mid- to Late Cretaceous was a time of exceptional rates of seafloor spreading and intraplate volcanism. This pulse in global oceanic crustal production is hypothesized to have caused increases in the levels of atmospheric carbon dioxide and global sea levels via increases in global oceanic ridge volumes, magmatic outgassing, and metamorphic decarbonation reactions (Schlanger et al., 1981; Larson, 1991; Berner, 1994; Larson and Erba, 1999). Fundamental problems, however, remain in terms of our understanding of these Cretaceous environments and their Paleogene equivalents.
One problem concerns our understanding of Cretaceous climate change at tectonic timescales. Maximum rates of CretaceousCenozoic ocean crust cycling and, therefore, inferred atmospheric carbon dioxide levels are thought to have occurred during Aptian/Albian time (Larson, 1991; Larson and Erba, 1999), but this significantly predates Cretaceous climatic optima as perceived from mineralogical evidence and existing 18O paleothermometric records (Fig. F6.). One explanation for this discrepancy, consistent with the timing of peak CretaceousCenozoic sea level (also Turonian), is that subducted crust and/or dating problems within the Cretaceous magnetic superchron (Superchron C34n) act to obscure a "hidden" Turonian pulse in ocean crust cycling (Wilson et al., 2002). Alternatively, the mismatch is real evidence of some other factor influencing CO2 and/or SST (e.g., higher rates of CaCO3 subduction) (Kump, 2002) during the Turonian relative to the AptianAlbian (Wilson et al., 2002).
A second problem with our understanding of Cretaceous climate concerns the short-term stability of the so-called greenhouse state. High-resolution bulk carbonate 18O records from classic land sections in Italy reveal positive excursions that have been interpreted in terms of large-scale mid-Turonian and early Cenomanian glaciations superimposed on the middle of the Cretaceous greenhouse (Stoll and Schrag, 2000). Sedimentological and biotic records show no support for this hypothesis, but these records are of insufficient temporal resolution to provide a categorical test. Similarly, our highest resolution long-term 18O record from deep-sea sites for the Cretaceous comes from diagenetically altered bulk carbonate and has a temporal resolution of ~1 sample/200 k.y. (Clarke and Jenkyns, 1999). The best existing corresponding record from separates of planktonic foraminiferal calcite is also diagenetically suspect (it comes from chertified, deeply buried chalks in the Pacific) and is of very low resolution (<1 sample/m.y.; AptianSantonian) (Barrera, 1994). More recently, high-resolution sequence stratigraphic correlation of mid-Cenomanian sediments in northwest Europe and southeast India have been interpreted in terms of eustatic sea level change at an orbital timescale, possibly of glacial origin (Gale et al., 2002).
A third problem concerns the magnitude and origin of warmth of intermediate and deep waters in the Cretaceous oceans. A recent study of the dynamics of this problem demonstrates that the popular concept, instead of low-latitude "warm salty bottom water" formation, is essentially unsupported (Bice and Marotzke, 2001). A simpler way to explain the warm Cretaceous temperatures recorded by deep-sea benthic foraminifers is by water-mass formation at high latitudes that have warmer SSTs than today (presumably in response to pCO2 forcing). However, the extraordinary magnitude of SST warmth indicated by a recent study for the Turonian high-latitude South Atlantic (Deep Sea Drilling Project [DSDP] Site 511; up to 32°C at ~60°S) (Bice et al., in press) raises severe questions concerning the levels of atmospheric pCO2 forcing required.
The following scientific questions will be addressed using well-preserved microfossils and organic-rich sediments recovered from the Demerara Rise transect:What is the history of changes in atmospheric CO2 levels from the mid-Cretaceous to Paleogene time? Well-preserved microfossils and organic carbon-rich sediments from Demerara Rise will provide an ideal means to evaluate this question using multiple proxies for atmospheric pCO2 (e.g., B isotope geochemistry in foraminiferal calcite and 13C geochemistry of bulk and biomarker organic carbon [e.g., 1989; Kuypers et al., 1999; Pearson and Palmer, 1999]).
The Paleogene record is rife with "critical boundaries" that offer significant opportunities for understanding the dynamics of greenhouse gas release, warm climate stability, biotic turnover associated with climate transitions, and extraterrestrial impacts. For example, the early Eocene warm period (~5053 Ma) is the most extreme interval of global warming in the Cenozoic, but little is known about the number of hyperthermals within it, the range of temperatures, or their effects on biotic evolution (Thomas and Zachos, 1999). The Eocene warm period is succeeded by a long shift toward the lower temperatures and increased ice buildup of the late Eocene and Oligocene (Fig. F6), whose history and consequences for ocean circulation, carbon cycling, and biotic evolution are only vaguely understood. Similarly, extraterrestrial impacts in the early middle Eocene and the late Eocene offer the opportunity to study the climatological and biotic effects of impacts that were too small to precipitate global mass extinctions but were large enough to have engendered global changes in climate. Below, we discuss two events that are particularly well expressed by the sediments recovered from the Demerara Rise transect.
The transient global warming at the end of the Paleocene is one of the best candidates for greenhouse warming in the geologic record. A growing body of evidence implicates a massive release of greenhouse gases into the atmosphere and ocean as a cause for ~5°7°C warming in the Southern Ocean and subtropics, a 35% 50% extinction of deep-sea benthic foraminifers, and widespread carbonate dissolution in the deep-ocean record (e.g., Zachos et al., 1993; Koch et al., 1995; Dickens et al., 1995). Recent studies utilizing high-resolution stable isotope analyses (Bains et al., 1999) and orbitally tuned chronologies (Norris and Röhl, 1999) suggest that carbon release occurred in a series of short steps, lasting a few thousand years, punctuated by catastrophic shifts in 13C and ocean temperature. Although these new data support the idea that the carbon may have been sourced from methane hydrate reservoirs, considerable uncertainty remains about how the carbon was released, what triggered the different phases of release, and what the biotic and climatological response was to the input of large amounts of greenhouse gas. We have little data to constrain fluctuations in the CCD during the P/E Event, a key parameter for understanding changes in CO2 storage resulting from hypothesized methane outgassing. There also remain significant questions about the chronology of the P/E Event. Thus far, cycle-based chronologies are based mostly upon analysis of a single site (ODP Site 1051). Analysis of additional cyclostratigraphic records tied to magnetostratigraphy is needed to evaluate the accuracy of the existing chronology and allow us to calculate changes in accumulation rates, rates of evolutionary and ecological responses to climate change, and biogeochemical fluxes during the PaleoceneEocene interval.
Two decades of study have provided considerable information on the causes and evolutionary consequences of the Cretaceous/Paleogene (K/P) mass extinction. There is now widespread agreement that the extinction was precipitated by a large-body impact event that created the ~300-km-wide Chicxulub impact structure on the Yucatan Peninsula in eastern Mexico, nearly 5000 km from Demerara Rise. There is considerable evidence linking the Chicxulub impact event to the K/P extinction, ranging from the dating of impact glass to ~65.5 Ma (the currently accepted age for the K/P boundary) to the presence of impact-generated diamonds, high-pressure minerals, and chondritic ratios of platinum group elements in K/P boundary sections worldwide. Although some have continued to dispute the role of the impact event in the associated mass extinction, arguing that the impact predates the extinction or that there were multiple impacts, nearly all boundary layers that are generally agreed to be stratigraphically complete contain only one horizon rich in impact debris, and that layer is associated intimately with faunal and floral evidence for mass extinction.
Recently, it has become apparent that the impact event led to significant disruption of the stratigraphic record both above and below the boundary. Seismicity produced by the impact (an event with a Richter scale magnitude of ~1013) and the effects of tsunamis produced by the impact eroded or caused mass wasting of Maastrichtian sediments throughout the Gulf of Mexico, the Caribbean, the North Atlantic, and Baja, California. Accordingly, it is not surprising that nearly all cores recovered from these areas display sedimentary disruptions or hiatuses associated with the boundary. The mass extinction, in turn, greatly diminished the supply of biogenic sediment to the seafloor, resulting in a significant drop in the sedimentation rate and, therefore, stratigraphic acuity in the earliest Paleogene. Hence, nearly all known K/P boundary sequences do not have rates of sedimentation sufficient to resolve many of the most pressing issues for analysis of extinction processes such as the climatic conditions that prevailed in the few hundred or even thousand years following the impact.
There are a series of outstanding issues that could be addressed by drilling on Demerara Rise:Where does the K/P boundary fall within Magnetochron C29r? Although the position and age of the boundary has been evaluated using cyclostratigraphic work in the South Atlantic, the cycle counts were not entirely unambiguous (within more than two precession cycles; ~42 k.y.) and require testing through further cyclostratigraphy and magnetostratigraphy work.
The opening of the equatorial Atlantic gateway was driven by the separation of Africa and South America and is widely hypothesized to have had a significant effect on both oceanic circulation patterns and heat transport over wide areas of the Cretaceous Atlantic. Yet, the timing of the opening of this gateway remains poorly constrained. Based on the biogeographic distribution of foraminifers and cephalopods, a shallow-water passage is thought to have been initiated between the North and South Atlantic Oceans at some time during the Albian (Moullade and Guerin, 1982; Förster, 1978; Wiedmann and Neugebauer, 1978; Moullade et al., 1993). Results from ODP Leg 159 on the eastern side of the equatorial Atlantic gateway suggest that a strong relationship existed between stepwise deepening and widening of the gateway and black shale deposition on the west African margin from the Albian to the Turonian (Wagner and Pletsch, 1999). Cessation of black shale deposition in the Late Cretaceous is interpreted to result from increasingly vigorous circulation between the North and South Atlantic. Hence, marking the transition from a Mesozoic longitudinal circulation system through the Tethyian and the central Atlantic Oceans to a more Cenozoic-like oxidizing latitudinal circulation pattern through the Atlantic gateway.
Analysis of the subsidence history of Demerara Rise will contribute to interpretations of the history of the opening of the equatorial Atlantic gateway. High-resolution sampling of distinct time slices across a range of paleowater depths will help to constrain the following questions:What were the timings of the establishment of oceanographically significant throughflow in the equatorial Atlantic gateway (i.e., the onset of throughflow of upper intermediate and deeper water masses)? Results will help to determine the paleoceanographic consequences of connecting the previously restricted South Atlantic to the North AtlanticTethyian realm.
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