Oceanic Anoxic Events (OAEs) mark the contemporaneous deposition of organic-rich marine sediments termed “black shales” in the wide areas of the oceans. An anoxic event that occurred at the Cenomanian-Turonian boundary, OAE-2, has been recognized as one of the largest events in the Cretaceous. Carbon isotopic compositions (δ13C) of sedimentary carbonate and organic matter exhibit a positive excursion across the OAE-2, reflecting an enhanced burial rate of 13C-depleted organic carbon during the event. Here we compile a spatiotemporal distribution of black shales on the basis of their onset timings relative to the δ13C excursion as a time-control reference, and discuss the “spreading patterns” of black shale deposition. The patterns suggest that the deposition of black shales started from marginal regions of the southern North Atlantic and the Western Interior Seaway in North America, and spread to the northern North Atlantic and Tethys Sea. Strangely, the black shales whose onset corresponds to that of the δ13C excursion have not been found in many locations. Furthermore, extensive deposition of black shales in the Tethys and some sites in the North Atlantic occurred significantly after the major shift of the δ13C excursion. Sediments in the largely unexplored Pacific basin may be the missing link in the temporal relationship between the black shale deposition and the isotopic excursion.
Introduction
The mid-Cretaceous (Barremian to Santonian; 127-84 Ma; Gradstein et al., 1994, 2004) is characterized by episodic and contemporaneous deposition of dark-colored muddy sediments enriched in organic carbon (black shales) in a wide range of marine settings. The depositional events of the black shales have been referred to as “Oceanic Anoxic Events” (OAEs). Originally, two OAEs were proposed (Schlanger and Jenkyns, 1976); OAE-1 in the early Aptian (121-120 Ma) and OAE-2 at the Cenomanian-Turonian boundary (around 94 Ma). Since then, six OAEs have been recognized as semiglobal or global events in the mid-Cretaceous (Table 1), and each of them was estimated to have lasted two million years or less (e.g., Jenkyns, 1991; Handoh, 2001).
Table 1.
Timings and durations of the Cretaceous oceanic anoxic events (OAEs).
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The mid-Cretaceous is also characterized by warm climate (e.g., Barron, 1983; Wilson and Norris, 2001). Based on oxygen isotopic records, Jenkyns et al. (2004) estimated that the sea surface temperature (SST) was over 20°C even in the Arctic region during the late Cenomanian to early Coniacian. The warm climate has been attributed to elevated atmospheric pCO2 (3–5 times higher than that of the present; Freeman and Hayes, 1992; Tajika, 1998) associated with the enhanced rate of volcanic CO2 degassing due to the hyperactive production of oceanic crust (Larson, 1991; Coffin and Eldholm, 1994; Kaiho and Saito, 1994; Eldholm and Coffin, 2000). Although new insights have been proposed in recent years (e.g., Ohkouchi et al., 1999; Leckie et al., 2002; Bice and Norris, 2002; Kuypers et al., 2002; Wagner et al., 2004), the triggers and formation processes of these OAEs still remain controversial.
Of these OAEs, OAE-2 at the Cenomanian-Turonian boundary is one of the most expanded expressions, because the organic-rich sediments were laid down in a vast area including the North and South Atlantic, Western Interior Seaway (WIS) in North America, Western Tethys, and Central Pacific (Figure 1, e.g., Schlanger et al., 1987; Philip et al., 1993, 2000; Pratt, 1985; Meyers et al., 2001). The OAE-2 was potentially associated with large, short-term changes in atmospheric pCO2 (e.g., Arthur et al., 1988; Kuypers et al., 1999; Hasegawa, 2005; Kuroda, 2005), and substantial changes in ocean chemistry (e.g., Ohkouchi et al., 1999; Kuypers et al., 2002; Snow et al., 2005). Furthermore, the Cenomanian-Turonian boundary marks one of the five major mass-extinction events during the last 100 Ma (Raup and Sepkoski, 1986; Kaiho, 1994). At the boundary, up to 25% of marine invertebrates such as deep-water bivalves and nekto-benthic ammonites underwent extinction (Harries and Kauffman, 1990; Hallam and Wignal, 1999). Kaiho (1994) compiled extinction patterns of planktonic and benthic foraminifers from various sites including the North Atlantic, WIS, European epicontinental seas, and northern Japan, and concluded that intermediate to deep marine benthic species and the deep-dwelling planktonic genus Rotalipora (e.g., Corfield et al., 1990; Norris and Wilson, 1998; Price and Hart, 2002) underwent disappearance within a relatively short time period (<500 kyr) near the last occurrence level of R. cushmani.
Figure 1.
Distributions of organic-rich black shales deposited during OAE-2 (94 Ma; Gradstein et al., 1994, 2004) compiled by Schlanger et al. (1987), Philip et al. (1993 (2000) and Lüning et al. (2004). Paleogeography is modified after Scotese and Golonca (1992). Filled circles indicate dark-coloured, organic-rich sediments with TOC contents generally higher than 2 wt%, whereas open circles indicate organic-poor sediments. Grey area shows land area at 94 Ma. Numbers indicate those of DSDP (in italics) and ODP drilling sites.
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The OAE-2 is characterized by a distinct positive excursion of the carbon isotopic compositions of marine carbonate and marine and terrestrial organic matter, which has been used as a reliable time slice for stratigraphic correlation (e.g., Arthur et al., 1988; Gale et al., 1993; Jenkyns et al., 1994; Hasegawa, 1997; Voigt, 2000; Tsikos et al., 2004; Bowman and Bralower, 2005; Kolonic et al., 2005). This isotopic excursion is conventionally related to the excess burial of 13C-depleted organic carbon into the sediments during the OAE-2 (Schlanger and Jenkyns, 1976; Jenkyns, 1980; Arthur et al., 1987, 1988). The enhanced burial of organic carbon was estimated to have lasted 3 to 8 × 106 years (Meyers et al., 2001; Ohkouchi et al., 1999; Kuroda et al., 2005; Sageman et al., 2006).
In this study we compiled the spatiotemporal distribution of organic matter-rich black shales during the OAE-2 and classified them based specifically on their onset timings relative to the initiation of the δ13C positive excursion. Here we define “OAE-2 black shale” to satisfy the following criteria; (1) stratigraphic intervals characterized by a significant increase in total organic carbon (TOC) content with certain thickness and continuity, and (2) intervals stratigraphically close to the positive excursion of carbon isotopic composition. Based on the temporal pattern of black shale deposition, we will discuss the genesis of black shales associated with the anoxic event.
Chronological frameworks
Our reasoning in this study strongly relies on the chronological correlation. To discuss the time lag of black shale deposition between sites, a rigid chronological framework is critically required. In this study, we chose the carbon isotopic excursion as a time-control reference, because 1) the isotopic composition in dissolved inorganic carbon (δ13CDIC) in the surface ocean quickly responds to perturbation of the carbon cycle in the atmosphere-surface ocean system, 2) variation of carbon isotopic composition can be simulated even by a simple kinetic model as described in the following section, and 3) the isotopic composition is recorded not only in sedimentary carbonate minerals but also in sedimentary organic matter, although the isotopic signatures are prone to be overprinted during diagenesis (particularly carbonate minerals). Boundaries of the biozones may also be a reliable time slice. However, since we do not know the synchronicity of the species replacement between oceans, we supplementarily use the biostratigraphic framework in this study.
A theoretical framework for variation in carbon isotopic composition
The oceanic carbon reservoir represents a balance between inputs from volcanic/metamorphic degassing and weathering, and outputs to sedimentary inorganic and organic carbons (e.g., Tajika, 1998; Kump and Arthur, 1999). Since sedimentary organic carbon is substantially depleted in 13C relative to the sedimentary carbonate, the δ13CDIC varies with respect to a ratio of organic fraction relative to global mass accumulation rate (MAR) of total carbon (hereafter this parameter is abbreviated to forg; Figure 2).
Figure 2.
Kinetic model concerning carbon cycle around the OAE modified after Tajika (1998) and Kump and Arthur (1999). For time before and after the OAE, we assumed that the organic fraction relative to the removal flux of total carbon (forg) was set at 23% to keep mass and isotope balances. During OAE-2 (from 0 to 80 ky), the forg is increased by 5, 10 and 15%, shown as dotted, solid, and dashed lines, respectively. Results of isotopic variations in sedimentary carbonate (δ13Ccarb) and organic carbons (δ13Corg) are shown.
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To understand the fundamental relationship between forg and δ13CDIC in global oceans during the OAE-2, we reconstructed a simple one-box model through the ocean-atmosphere carbon budget (Figure 2). In the model we adopted “representative Phanerozoic values” by Kump and Arthur (1999) for input parameters (fluxes and δ13C values) and carbon amount in the ocean, and assumed that they were constant across the OAE-2. Isotopic difference between DIC and organic carbon was set at 280. Carbon isotopic compositions of sedimentary carbonate and organic carbons illustrate 1.5 to 4.10 positive excursions in response to increments of forg by 5 to 15% (Figure 2). The most realistic case is a 10% increase in forg, which results in 2.70 positive excursions of δ13Corg from −26.5 to −23.80 and δ13Ccarb from 1.5 to 4.20 (Figure 2). A rapid positive sift occurs immediately after increase of forg. Therefore, the base of the positive shift of δ13C (a narrow shaded interval in Figure 2) corresponds to the theoretical “onset” of the global increase in relative accumulation rate of organic carbon, namely, a global oceanic anoxic event.
Carbon isotopic stratigraphy and foraminiferal biostratigraphy
The OAE-2 spans two planktonic foraminiferal biozones (Figure 3; see also Leckie, 1985; Premoli Silva et al., 1999; Coccioni and Luciani, 2005; Keller and Pardo, 2004): the Rotalipora cushmani Zone in the late Cenomanian (96.6 to 94.0 Ma; Bralower et al., 1997) and the Whiteinella archeocretacea Zone in the latest Cenomanian to early Turonian (94.0 to 93.0 Ma). The Cenomanian-Turonian boundary, defined as an ammonite biostratigraphic boundary between the Neocardioceras juddii and Watinoceras devonense Zones (e.g., Bengtson et al., 1996), falls within the W. archeocretacea Zone. Gale et al. (1993) compiled carbon isotopic records of OAE-2 sediments from South England and WIS to correlate them with the foraminiferal biostratigraphic zones. They concluded that the initiation of the positive isotopic shift corresponds to the uppermost part of the R. cushmani Zone, and that the δ13C values reach a maximum at the boundary between the R. cushmani and W. archeocretacea Zones (Figure 3). A maximum plateau phase in the carbon isotopic record is observed in the lower part of the W. archeocretacea Zone. The record then gradually returns to preexcursion values in the early Turonian.
Figure 3.
A schematic diagram representing the positive excursion of δ13C in sedimentary records with biostratigraphic zones of planctonic foraminifers across the OAE-2 interval (modified after Gale et al. (1993), Kuypers et al. (2002) and Tsikos et al. (2004)). Also shown are three types of black shales classified based on their onset timings (see text).
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Spatiotemporal distribution of OAE-2 black shales
In this section, we compiled the evidence of OAE-2 black shales reported previously to categorize them into three types on the basis of the onset timing. They are those which started to deposit 1) long before the δ13C positive excursion started (Type I), 2) simultaneously with the onset of the isotopic excursion (Type II), and 3) long after the isotopic excursion started (Type III).
Type I black shales
In some regions of the North Atlantic, it has been reported that the deposition of organic-rich sediments started significantly earlier than the onset of the δ13C excursion associated with the OAE-2. The regions include the Demerara Rise (DSDP/ODP Sites 144, 1258, 1260 and 1261; Kuypers et al., 2002; Erbacher et al., 2004), western Venezuela coast (Perenz-Infante et al., 1996), and proximal part of the Tarfaya coastal basin in Morocco (S75; Luderer, 1999; Lüning et al., 2004; Kolonic et al., 2005) (Figure 4).
Figure 4.
Stratigraphic variations in concentration of total organic carbon (TOC) and isotopic composition of organic carbon (δ13Corg) across Type I black shales. From right, Colorado section in the Western Interior Seaway in USA (after Meyers et al., 2001; Sageman et al., 1997; Sageman and Lyons, 2004), southwestern parts of North Atlantic (DSDP Site 144 after Kuypers et al. (2002), and ODP Site 1261 after Erbacher et al. (2004)), and proximal part of Tarfaya coastal basin in southeastern North Atlantic (Site S75; after Luderer, 1999; Lüning et al., 2004). Lithological columns indicate occurrences of dark and/or light intervals. Plausible ranges of the onset of the δ13C positive excursion are shaded. Dashed line indicates the biostratigraphic boundary between the R. cushmani and W. archaeocretacea Zones.
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At DSDP Site 144 from the Demerara Rise, in the OAE-2 interval the TOC content (2 to 7%) was somewhat lower than that before the onset of the positive isotopic shift (5 to 10%) (Figure 4; Kuypers et al., 2002). At ODP Sites 1258, 1260 and 1261 from Demerara Rise (Erbacher et al., 2004), clear positive δ13C excursions were also observed around the Cenomanian-Turonian boundary. In the OAE-2 interval at Site 1261, the TOC content ranges mostly from 5 to 19%, which is only slightly elevated relative to the pre-OAE-2 interval (3 to 13%). Therefore, we can say that the MAR of organic carbon in the Demerara Rise could have remained high and not changed significantly across the onset of the δ13C excursion.
It has also been reported that the pre-OAE-2 sediments (e.g., Hartland Shale) from the WIS contain abundant organic carbon (Sageman et al., 1997, 1998; Meyers et al., 2001; Simons and Kenig, 2001; Sageman and Lyons, 2004). The TOC content is continuously high in intervals beneath the base of the carbon isotopic excursion, whereas it decreased down to <1% simultaneously with the onset of the isotopic excursion at the basal part of the Bridge Creek Limestone (Figure 4). The temporal decrease in TOC content is followed by a rapid increase in TOC at the maximum plateau phase of the isotopic excursion.
Sediment successions described above, in which sedimentary TOC content exhibits no substantial increase across the onset of OAE-2, imply that the MAR of organic carbon was not substantially changed at these sites. In this sense, we can categorize these sediments as “Type I” black shales (Figure 3). It should be noted that the Type I black shales are distributed only in marginal regions of the southern North Atlantic and WIS (Figure 5). These regions may be called “refugia” of the black shale. Although the MAR of organic carbon should have been high at these sites, the Type I black shales should have not contributed significantly to the global carbon cycle, because they lack temporal correlation with the major positive shift of δ13CDIC.
Figure 5.
Distribution of black shales categorized as Type I (filled circle), II (filled square) and III (filled triangle). Paleogeography is after Scotese and Golonca (1992).
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Type II black shales
The OAE-2 black shales deposited simultaneously with the initiation of the δ13C excursion are categorized as “Type II” (Figure 3). These black shales must be responsible for the positive δ13C shift in the atmosphere-ocean system during the OAE-2. Strangely, we did not find many Type II black shales. Based on the previous results, they appear to be observed in the southern North Atlantic (Figure 5). A typical black shale of this type is the one observed in the distal part of the Tarfaya coastal basin (site S13; Kuhnt et al., 1990; Kuypers, 2001). Although organic-rich sediments (TOC > 5%) are sporadically intercalated in an interval below the δ13C excursion at this site, the TOC profile illustrates a substantial increase at the base of the positive δ13C excursion (Figure 6). It suggests that MAR of organic carbon significantly increased at the onset of OAE-2.
Figure 6.
Stratigraphic variations in TOC content and δ13Corg value in the vicinity of Type II black shales deposited in distal part of Tarfaya coastal basin (Site S13; after Kuhnt et al., 1990 and Kuypers, 2001) and southeastern North Atlantic (DSDP Site 367 after Herbin et al., 1986 and Kuypers, 2001). Plausible ranges of the onset of the δ13C positive excursion are shaded. Dashed line indicates the biostratigraphic boundary between the R. cushmani and W. archeocretacea Zones.
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Another example is the OAE-2 interval recovered from DSDP Site 367, offshore Senegal (Herbin et al., 1986; Kuypers, 2001). At this site, a significant increase in TOC content was observed at the base of a δ13C positive shift (Figure 6). However, thin organic-rich layers are also observed several meters below the OAE-2 horizons (Figure 6). Therefore, this black shale may not be typical Type II, but could have had an origin somewhat similar to those of Type I. The deposition of black shales may have expanded from the marginal regions of the southern North Atlantic such as Venezuela, the Demerara Rise, and the Tarfaya coastal sea to wider parts of the southern North Atlantic including Site 367 and S13 (Figure 5).
Type III black shales
“Type III” black shale is defined as those deposited after the initial isotopic shift of the major excursion (Figure 3). In most cases, the timing corresponds roughly to the base of the maximum plateau of the positive δ13C excursion, and to the biostratigraphic boundary between the R. cushmani and W. archeocretacea Zones. A representative of Type III black shale is the Livello Bonarelli cropping out in the Apennines, central Italy, which was deposited on the pelagic marginal shelf in the western Tethys Sea (Arthur and Premoli Silva, 1982; Ohkouchi et al., 1997, 1999; Tsikos et al., 2004; Kuroda et al., 2005). The base of the Bonarelli black shale is located immediately above the last occurrence of R. cushmani (Tsikos et al., 2004; Coccioni and Luciani, 2005). Variation of the δ13C record of organic matter in the Bonarelli black shale and adjacent limestone samples indicates that the initiation of black shale deposition can be correlated with the maximum plateau phase of the δ13C excursion, just above a sharp negative δ13C spike that punctuates the positive excursion (Figure 7; Kuroda, 2005; N. Ohkouchi, unpublished data). Sediments deposited at the Oued Smara section in Tunisia around the Cenomanian-Turonian boundary also document the onset of an organic-rich interval coincident with the base of the maximum plateau phase of the δ13C excursion, near the base of the W. archeocretacea Zone (Figure 7; Accarie et al., 1996; Lüning et al., 2004). A similar sequence was also observed in the equivalent succession in the Nebour section in Tunisia (Nederbragt and Fiorentino, 1999; Lüning et al., 2004), indicating an initial deposition of the black shale (TOC ~6%) at the base of the δ13C excursion close to the base of the W. archeocretacea Zone. Based on the paleogeographic reconstruction by Philip et al. (2000), both Italian and Tunisian sections were located in the western part of the Tethys in the latest Cenomanian.
Figure 7.
Stratigraphic variations in TOC contents and isotopic compositions of organic carbon (δ13Corg) and carbonate carbon (δ13Ccarb) across Type III black shales. From right, eastern North Atlantic (DSDP Site 105 after Herbin et al. (1987) and Kuypers et al. 2004a), western North Atlantic (ODP Site 641 after Thurow et al., 1988), western Tethys (Oued Smara in southwestern Tunisia after Accarie et al. (1996) and Lüning et al. (2004), and Apennines in central Italy after Kuroda (2005) and N. Ohkouchi (unpublished data)). Lithological columns indicate occurrences of dark and/or light intervals. Plausible ranges of the onset of the δ13C positive excursion are shaded. Dashed line indicates the biostratigraphic boundary between the R. cushmani and W. archeocretacea Zones.
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At ODP Site 763 in the Exmouth Plateau, eastern Indian Ocean, Rullkötter et al. (1992) observed a distinct black shale layer with elevated TOC content of up to 26%. At ODP site 762 in the Exmouth Plateau, a 20-cm-thick dark layer has also been recovered (Shipboard Science Party, 1990). Although carbon isotopic records are as yet unavailable for these sites, these black shales may be categorized as Type III because they span the W. archeocretacea Zone (Wonders, 1992). But it should be noted that the foraminiferal assemblages in this site might reflect local dissolution effects that made the biostratigraphic boundaries unclear. In the western margin of the Tethys (Apennines and Tunisia), as well as the Exmouth Plateau, the sufficient elevation in accumulation rates of organic carbon occurred after the major shift of the δ13C excursion (Figure 5). Thus, it is noteworthy that the deposition of organic-rich sediments in at least several areas of the Tethys Sea should have occurred significantly after the onset of the δ13C excursion. In the central Tethys such as Tibet, the OAE-2 sediments exhibited slightly elevated TOC content (~1.7%, Wang et al., 2001), although the precise timing of the onset of the organic matter-rich interval is unknown.
At ODP Site 641 in the eastern North Atlantic (Galicia Margin, offshore Spain), the sedimentary succession across the Cenomanian-Turonian boundary also contains black shales with elevated TOC of up to 13% (Figure 7; Thurow et al., 1988). Because the onset of the organic-rich sediment at this site clearly corresponds to the end of the major positive shift during the δ13C positive excursion (Figure 7), it should be categorized as Type III. An equivalent interval from DSDP Site 105, off Cape Hatteras (western North Atlantic), would also be categorized as this type, whose onset corresponds to the base of the maximum plateau phase of the δ13C excursion (Figure 7; Herbin et al., 1987; Kuypers et al., 2004a). These evidences suggest that the black shale deposition also spread at this time to the northern parts of the North Atlantic, at least off Cape Hatteras and the Galicia Margin (Figure 5).
Pacific Basin: the missing link?
Unlike the North Atlantic, WIS, and western Tethys, spatiotemporal distribution of organic-rich sediments in the proto-Pacific Ocean is less well understood. So far, only three black shales of OAE-2 have been reported from the Pacific pelagic basin; Shatsky Rise (Figure 1, DSDP Site 305), Hess Rise (DSDP Site 310) and Mariana Basin (ODP Site 585) (Schlanger et al., 1987 and references therein). Since the black shales from Shatsky and Hess Rises were sampled only as drilling breccia, we cannot precisely reconstruct the onset timing of these black shales. At Mariana Basin, a layer of 2 cm-thick black shale with the maximum TOC of 9.9% was recovered (Shipboard Scientific Party, 1986; Schlanger et al., 1987). This black shale layer was postulated to be a reworked bed, because of the mixture of planktonic foraminiferal species including W. archeocretacea, R. cushmani, and R. greenhornensis (Premoli Silva and Sliter, 1986), and lithological features of the adjacent successions showing frequent intercalation of reworked deposits like turbidites (Whitman et al., 1986). Drilling at Shatsky Rise during ODP Leg 198 also failed to recover the black shales corresponding to OAE-2 (Shipboard Scientific Party, 2002).
Sedimentary successions including the Cenomanian-Turonian boundary have been well investigated in the Yezo Group, distributed in northeastern Japan (e.g., Hirano, 1995; Kawabe et al., 1996; Hasegawa, 1997; Nishi et al., 2003; Takashima et al., 2004). A distinct positive excursion with a short-term negative spike was observed in carbon isotopic composition of organic matter in a pyrite-rich mudstone interval within the W. archeocretacea Zone (Hasegawa, 1997). However, in this region organic-rich black shales were not observed across the Cenomanian-Turonian transition (Hasegawa, 1997; Takashima et al., 2004).
General discussion and conclusions
In this study, we classified the OAE-2 black shales into three types based on their onset timing. Although the amount of data is rather limited, we found apparent spatial distribution of the onset timing of black shale deposition. Black shale deposition started in the marginal regions of the southern North Atlantic and WIS before the major isotopic excursion, then appears to have spread to the wider regions of the southern North Atlantic at the onset of the δ13C excursion, and then further spread to the Tethys Sea and northern parts of the North Atlantic during the plateau phase of the isotopic excursion. We consider that there were “refugia” of OAE-2 in coastal regions of the southern North Atlantic and WIS, and that deposition of organic-rich sediments spread from the refugia to the other oceans (Figure 5). However, it should be noted that it is not necessary that the spreading pattern of the depositional sites of black shale be continuous laterally like “blankets” of black shale, because organic-poor sediments were also deposited in some sites such as the western North Atlantic (e.g., DSDP Site 391; Figure 1).
A puzzling phenomenon is that we have not observed many black shales whose depositions temporally correspond to the isotopic excursion (i.e., Type II black shale). Furthermore, to our knowledge, the Type II black shales have been found only in the southern North Atlantic (Figure 5). As discussed previously, the 30 positive shift of δ13C in the oceanic carbon reservoir can be explained by a ca. 10% increase in forg (Figure 2). Therefore, the black shale deposition must have been substantially enhanced at that time, probably in wide areas of the ocean. However, currently available geological records from many areas including the Tethys Sea indicated that the black shales did not begin to deposit simultaneously with the onset of the δ13C positive excursion. We think that the Type II black shales were widely deposited somewhere we have not investigated yet.
The spreading pattern of black shale suggests an important implication for the paleoceanography during the OAE-2 interval. A potential scenario is that the pattern reflects an expansion of the anoxic water masses from the refugia to the global ocean, although the anoxic water masses did not spread over a wide area of the ocean. The expansion of the anoxic water masses could be attributed to enhanced oceanic stratification. Kaiho (1994) proposed that habitat-specific extinction of foraminifers around the top of the R. cushmani Zone resulted from the global expansion of an anoxic water mass in the intermediate to deep waters.
Alternatively, the spreading pattern of black shale deposition could reflect changes in surface ocean ecology. Recent organic geochemical and nitrogen isotopic evidence indicate that organic matter in the Cretaceous black shales is commonly characterized by a large contribution of diazotrophic cyanobacteria (e.g., Ohkouchi et al., 1997, 2003, 2006; Kuypers et al., 2004b; Dumitrescu and Brassell, 2005). If this were the case, spreading of a cyanobacterial bloom could have been directly responsible for the expansion of black shale deposition during OAE-2. The extensive blooming of diazotrophic cyanobacteria has frequently been observed in the euphotic zone of the stratified water column as well as in warm, nutrient-poor open ocean regions such as the tropical Atlantic, Red Sea, and South China Sea (e.g., Capone et al., 1997; Montoya et al., 2004). Ecology of the planktonic cyanobacteria remains largely unknown and factors controlling the cyanobacterial blooming in the surface ocean are still a matter of debate (e.g., Pearl, 1996; Bianchi et al., 2000). However, a nutrient other than nitrogen, either phosphorus, molybdenum, or iron, is potentially a limiting factor for the formation of cyanobacterial blooms (Falkowski, 1997; Fuhrman and Capone, 2000; Dyhrman et al., 2006). Therefore, the global spreading pattern of black shale may have reflected the nutrient distribution of surface water. More evidence is required for critically evaluating and extending our view described in this study. Extensive exploration for black shales in the Pacific Ocean may provide the solution, although much of the Mesozoic oceanic sediments in the Pacific have been lost due to plate subduction.
Acknowledgments
We thank H. Kitazato, A. Taira, M.F. Coffin, T. Sakamoto, H. Nishi, S. Kiyokawa, M. Okada, Y. Kashiyama, E. Tajika, R. Tada, and H. Tokuyama for helpful comments and discussion. Comments from K. Moriya and an anonymous reviewer helped to improve the manuscript. This work was financially supported by the Japan Society for the Promotion of Science (Grant number 16-10902).