Abstract
The organic-rich, black mudstones that were initially described as the Black Band in Lincolnshire, Humberside and Yorkshire are known to be a local representation of the Cenomanian–Turonian Boundary Event (CTBE). This world-wide event is known as Oceanic Anoxic Event ll (OAEll) and it marks a distinctive extinction event within the Cretaceous biota. Since some of the original work on the benthic foraminifera that are found in both the Black Band and coeval sedimentary rocks, there has been a significant increase in the understanding of the biology of foraminifera, and their response to both modern and fossil low-O2 environments. While the overall event is clearly global, the local response appears to be a function of both geological setting and water depth with the occurrence of organic-rich sediments as a combination of this setting, plankton productivity and preservation.
The Cenomanian–Turonian Boundary Event (CTBE) or Oceanic Anoxic Event ll (OAEll) is regarded as one of the major global bioevents, though not as dramatic as one of the ‘big five’ (Raup & Sepkoski 1982, 1984; Milne et al. 1985; Hart 2005). The biotic changes at the CTBE are at the generic or species level, rather than that of family or above, and the percentage change is not so significant on a global scale (Gale et al. 2000). Perhaps more significant is the importance of this event as one of the Cretaceous OAEs that begun in the early Cretaceous (Weissert Event, Faraoni Event, etc.) and continued into the Santonian (OAElll): see Leckie et al. (2002). The Black Band of Yorkshire, Humberside and Lincolnshire is, of course, where the study of Cretaceous OAEs began with the seminal work of Schlanger & Jenkyns (1976) and the subsequent work on stable isotopes by Schlanger et al. (1987). The quarry on the south bank of the Humber Estuary, at South Ferriby, was the location where the first description of an oceanic anoxic event was related to changes in the water column, notably anoxia. Since that time, OAEll has been recognized as a global event (Hart & Ball 1986; Koutsoukos et al. 1990; Leckie et al. 2002), being recorded in most of the world's oceans (DSDP, ODP, IODP cores), and on every continent in both deep water mudstones (Bąk et al. 2001; Bąk 2006, 2007a, b) and relatively shallow water carbonate successions (Caus et al. 1997; Parente et al. 2008; Wohlwend et al. 2015, 2016).
Cretaceous stratigraphy in NE England
The Chalk Group of the East Midlands Shelf in NE England is represented by a succession different to that in southern England. As a result, an alternative lithostratigraphy has been developed (Wood & Smith 1978; Jeans 1980; Sumbler 1999; Mortimore et al. 2001). While other authors (e.g. Hart et al. 1991, 1993) attempted to use this ‘local’ stratigraphy, some authors (e.g. Jeans et al. 1991) have mixed the nomenclature, referring to the ‘Black Band’ and overlying nodular chalks as the Plenus Marls and Melbourn Rock (see Jefferies 1962, 1963; Jarvis et al. 1988) rather than the Flixton Member, thereby complicating the understanding of the NE England succession.
In Norfolk, Lincolnshire, Humberside and Yorkshire there is a transition from the typical ‘southern’ succession of the pale green/grey mudstones of the Plenus Marls Member (at Marham) to an omission surface at Hillingdon and the appearance of the more typical ‘northern’ succession at Heacham (Fig. 1). The CTBE succession north of Heacham was described over twenty years ago by Hart et al. (1991, 1993) and more fully by Mortimore et al. (2001). Between South Thoresby and Louth (Hart et al. 1993, Fig. 2) black mudstones appear within the CTBE succession, continuing northwards through Caistor (UK National Grid Reference TA 1234 0019), Elsham (TA 038 131), South Ferriby (SE 9930 2030), Welton (SE 9668 2788) and in the cliffs of Flamborough Head (TA 195 743), SE of Speeton (Bempton Cliff).
Locality map of the Cretaceous strata in NE England and location of sections discussed in the text (modified after Hart et al. 1993, fig. 1). Line A–A marks the southern limit of the ‘Black Band’. Between lines A–A and B–B the chalks and nodular chalks of the uppermost Cenomanian are dull red in colour, while south of line B–B the chalks are the same stratigraphical level are pale green/grey in colour.
At South Thoresby (TA 773 406) chalks have been described (Hart et al. 1993, p. 506) as including ‘pink’ levels and some ‘reddened’ horizons, but there are no black or grey mudrocks. The three most important successions were those at South Ferriby, Caistor and Elsham, although the latter has now been infilled and landscaped. When accessible, all these quarries exposed the black or grey mudstones that were separated into distinct beds by paler-coloured chalks and marlstones. In the darkest of the black mudstones at South Ferriby (Figs 2–4), Hart et al. (1993, fig. 4, p. 500) recorded almost 9000 dinoflagellate cysts per gram, though species richness (diversity) was less than comparable levels in the Plenus Marls Member of southern England (Jarvis et al. 1988). The data presented by Hart et al. (1993) was in agreement with the earlier work of Marshall & Batten (1988) who concluded that the palynological data from the South Ferriby succession represented a stressed environment with low levels of dissolved O2 extending up into the water column. Low O2 levels at the sediment/water interface had also been suggested by Hart & Bigg (1981) on the basis of foraminiferal evidence. The black mudstones within the Flixton Member contain low diversity assemblages dominated by simple agglutinated taxa such as Ammodiscus and Glomospira (Fig. 5). This restricted assemblage is also recorded at Elsham (Hart et al. 1993, figs 5 and 6). While some aspects of the distribution of planktic foraminifera are typical of the CTBE interval, it is noteworthy that Rotalipora cushmani always disappears at the base of the Flixton Member and that Thalmanninella greenhornensis is not recorded in any of the successions. The latter is probably due to the palaeolatitude, with T. greenhornensis being at its northern limit in southern England, but the absence of R. cushmani needs careful consideration. In the Plenus Marls succession of southern England R. cushmani invariably disappears at the Bed 3/4 boundary, with the succession up to the first appearance of Helvetoglobotruncana helvetica being known as the Whiteinella archaeocretacea Interval Zone. Helvetoglobotruncana helvetica is not recorded in NE England and this is probably due to palaeobiogeography as this species is widely regarded as a warm-water (Tethyan) taxon. Dicarinella imbricata and Marginotruncana marginata are often associated with H. helvetica in southern England and so indicate the general position of the Cenomanian/Turonian boundary; see Gebhardt et al. (2010). The changes in the benthic foraminifera are also supportive of this interpretation with species such as Gavelinella baltica, G. cenomanica and Plectina mariae all disappearing alongside R. cushmani and the benthic assemblage subsequently being dominated by Gavelinella berthelini (= Gavelinella dakotensis of some authors) and Lingulogavelinella globosa (Fig. 5). The disappearance of R. cushmani is noted below several CTBE ‘anoxic events’ (e.g. DSDP Site 551, Goban Spur; Leary & Hart 1989). In the Goban Spur successions R. cushmani and T. greenhornensis disappear together, which is not the normal situation in stratigraphically complete sections and indicates that the black mudstones are not recording the full ranges of the species.
The working quarry at South Ferriby. (1) Ancholme Clay Group; (2) Thin development of the ‘Carstone’; (3) Thin development of the Red Chalk; (4) Ferriby Chalk Formation; and (5) Welton Chalk Formation. Reproduced with permission from Mortimore (2014, fig. 4.15)
The location of the ‘Black Band’ at the base of the Welton Chalk Formation, which lies above the Ferriby Chalk Formation. Reproduced with permission from Mortimore (2014, fig. 4.21).
Close-up of the ‘Black Band’. Abbreviations are as follows: SKM, ‘silty khaki marl’; BB, Black band; GM, gungy marl; XBU, ‘cross-bedded unit; and AP, ‘Adrian's pair of marls’. Reproduced with permission from Mortimore (2014, fig. 4.23a).
The lithological succession at South Ferriby and the distribution of some key species of foraminifera (modified after Hart et al. 1993, fig. 3). Letter (a) denotes the ‘60 cm bed’; letter (b) denote the ‘Black Band’; and letter (c) denote the ‘Inoceramus Pebble Bed’.
In many of the successions investigated by Hart et al. (1993, fig. 2) and Dodsworth (1996) in Lincolnshire and Humberside the general distribution of foraminifera varies little from that shown in Figure 5. The erosive boundary at the top of the Ferriby Chalk Formation appears to be co-extensive with the sub-Plenus Marls erosion surface of southern England. The overlying Flixton Member, and particularly the ‘Black Band’, have often been taken as the Northern Province equivalent of the Plenus Marls Member. This view was initially supported by the occurrence of the belemnite Actinocamax plenus (Hill 1888), although this record is unsubstantiated (see Wood in Gaunt et al. 1992; Wood & Mortimore 1995; Wood et al. 1997). Wood (in Gaunt et al. 1992, p. 88) listed the fossils typical of the Plenus Marls Member (Beds 4‒6 and Bed 7) and noted (op. cit., pp. 88–89) the extinction of R. cushmani at the top of the Ferriby Chalk Formation while its range in southern England invariably extends to the top of Bed 3 in the Plenus Marls Member. While this is all true, one must be careful at making biostratigraphical conclusions using a species at the very northern limit of its distribution.
At Melton Ross, which is located immediately SE of Elsham, Wood & Mortimore (1995) and Wood et al. (1997) described an anomalously thick succession of the Flixton Member with – as shown in Wood & Mortimore (1995, fig. 2) and Wood et al. (1997, fig. 3) – a number of dark-coloured mudstones that appear to be stratigraphically older than the successions at South Ferriby, Caistor and Elsham. Though studied for palynomorphs and geochemistry these ‘older’ mudstones were not investigated for foraminifera. The presence of characteristic benthic foraminifera of the Plenus Marls Member (Beds 1–3) and/or the presence of R. cushmani (and T. greenhornensis) would have confirmed the placing of the R.cushmani/W. archaeocretacea boundary and facilitated an accurate correlation of these mudstones with Beds 1‒3 of the Plenus Marls succession in southern England. Should this anomalous succession at Melton Ross ever be exposed again, then it is imperative that a foraminiferal analysis is undertaken.
Wood et al. (1997, p. 342) equated their ‘Bed 3’ with Bed 3 of the Plenus Marls Member succession using the distribution of the dinoflagellates cyst Lithosphaeridium siphoniphorum. This correlation would confirm that the main peaks in the stable isotope excursion (CIE) occur in Beds 4–6 of the southern England succession, just as in the Rheine succession of Germany where black shales appear to equate with the higher parts of the Plenus Marls succession (Ernst et al. 1983, 1984).
The Black Band ‘proper’ is, therefore, coeval with Bed 4 of the southern England succession (Wood et al. 1997). It contains virtually no macrofossils aside from fish debris that is often concentrated into slightly lighter coloured mudstones within the overall black mudstones. Fish debris (scales and ichthyoliths) are common in the microfossil residues (Fig. 5), together with an unusual assemblage of simple, agglutinated foraminifera (e.g. Ammodiscus, Glomospira).
Δ13C stable isotope excursion at the CTBE
Stable isotope data now provide valuable evidence on the global carbon cycle and the complex interplay between the biosphere–atmosphere–hydrosphere–lithosphere (Jenkyns 2010; Jarvis et al. 2015 and references therein). Major episodes of organic-rich, black shale deposition of Cretaceous age, corresponding to the ‘Oceanic Anoxic Events’ (OAE) such as OAEll at the Cenomanian/Turonian boundary have been described by many authors (e.g. Leckie et al. 2002). As organic carbon is preferentially enriched in the lighter isotope 12C, its removal from the oceanic reservoir renders global sea waters relatively enriched in 13C (Marshall 1992). Hence, positive carbon-isotope excursions (e.g. at the CTBE) have been interpreted in terms of increased burial of organic carbon attributed to enhanced preservation under reduced O2 conditions (e.g. Bralower & Thierstein 1984) or driven by changes in surface water productivity (delivering more organic carbon to the sea floor, e.g. Erba 1994). Positive δ13C excursions indicate higher levels of organic carbon burial and imply a global change in the ocean system that can be stratigraphically important for correlation. In recent years the use of such data in precise correlation has become an important part of stratigraphy, often giving more precision than biostratigraphical data (Jarvis et al. 2006, 2015; Jenkyns 2010). Associated with the positive δ13C excursion at the CTBE are a number of locations, again world-wide, that record the presence of anoxic, black mudstones often with a high TOC content. One of the first described was that in South Ferriby (Schlanger & Jenkyns 1976). At the time, and subsequently, an expansion of the oxygen minimum zone (OMZ) has been postulated as the causal mechanism (e.g. Jarvis et al. 1988).
Warming/cooling at the CTBE
In the early descriptions of the Plenus Marls Member in south-east England, Jefferies (1962, 1963) drew attention to warm/cool alternations between Beds 1–8 of the succession, especially at Merstham (his type locality). These warming/cooling events were related to faunal changes and, especially, the influx of the cool-water belemnites that give their name to the member. Jeans et al. (1991) suggested that the Plenus Marls Member was recording a ‘glacial’ event in high latitudes, despite the prevailing view of the mid-Cretaceous being that of a ‘greenhouse world’ (Wilson et al. 2002; Bice et al. 2003; Schouten et al. 2003; Moriya et al. 2007; Gallagher et al. 2008; Wang et al. 2014; Wendler & Wendler 2016). Over the last 20+ years numerous authors have taken up the ‘Plenus Cold Event’ theme (e.g. Voigt et al. 2003, 2006, 2015) using both palaeontological and geochemical information.
A number of ‘cold’ – or cool – events taking place within the 500 kyrs – 750 kyrs of the Plenus Marls Member implies either palaeoceanographic change (disturbing water masses) or volcanic activity. The latter, in the short term, can cause cooling as seen in northern Europe in 1783, after the famous eruption of the Laki Fissure in Iceland. Europe suffered three exceptionally harsh winters and failed harvests in the summer months. In Iceland, 75% of the livestock died, as did c. 25% of the population. Benjamin Franklin, who at the time was the Commissioner for the United States in Paris, wrote one of the first papers that tried to link volcanic eruptions to changing levels of CO2 in the atmosphere and climate change in 1784. Though not a scientist, his work is often forgotten, but his ideas were extended by Chamberlin (1897) who must be regarded as the first geoscientist to attempt the link between atmospheric pCO2 and climate, a debate that continues today. In the longer term (climate rather than weather), the CO2 erupted normally causes warming, as evidenced by the Deccan Traps in the latest Cretaceous and earliest Paleogene (Courtillot & Renne 2003; Chenet et al. 2007, 2009; Adatte et al. 2014; Keller 2014; Punekar et al. 2014; Schoene et al. 2015).
The CTBE has been linked to volcanic activity in the Caribbean Large Igneous Province (Sinton & Duncan 1997; Kerr 1998; Wignall 2001; Turgeon & Creaser 2008), though the dating of this is uncertain. More recently Bond & Grasby (2017, fig. 8) have shown the close proximity of volcanic activity on the Ontong Java Plateau (2), the Caribbean LIP and on Madagascar. While such volcanic activity has been linked to extinctions (Courtillot 1999; Wignall 2001, 2007; Courtillot & Renne 2003; Bond & Wignall 2014) cooling is not – apparently – part of the processes involved. Recent work on strontium, osmium, molybdenum, neodinium, sulphur and iodine geochemistry around the CTBE suggests that the intrusion and/or weathering of a mafic, Large Igneous Province (LIP) – either Caribbean or High Arctic – may have been occurring at this time (Owens et al. 2013; Jenkyns et al. 2017). The changes in water geochemistry may have been responsible for increased productivity (generating the OAE) and the draw-down of CO2 produced the Plenus Cold Event(s), which can now be correlated quite widely.
Oxygen minimum zone in modern oceans
In the modern ocean, surface water productivity often creates an oxygen-deficient layer that is (approximately) located between 200–700 m water depth (see, for example, Hart & Koutsoukos 2015, fig. 4). Below the OMZ, oxygen levels recover and by a depth of c. 1250 m can be back to surface values. These deeper, oxygenated, waters are created by the present-day, cold water driven ‘global oceanographic conveyor’ (Broecker 1969, 1991).
In the Cretaceous ocean the modern ‘global oceanographic conveyor’ (Hay et al. 2005, fig. 8) did not exist and the oceanographic circulation was more fragmented: this was the ‘Eddy Ocean’ of Hay et al. (2005, fig. 11) and Hay (2008). If O2 levels did not recover (as today) below the OMZ, what effect would that have on both the sedimentary succession and the enclosed benthic biota?
If the benthic foraminiferal assemblage changes that we see in the mid-Cretaceous are the result of the interaction with a mid-Cretaceous OMZ, then what can we learn about this process from modern oceans? The present-day OMZ intersects with many continental shelves, with good examples being offshore Pakistan and Oman in the Indian Ocean. In these areas, the sediments deposited within the OMZ are rarely – if ever – organic-rich and there are certainly no ‘black shales’ (= anoxic events) recorded. While changes in foraminiferal distributions are recorded (Gooday et al. 2000, 2009; Larkin & Gooday 2009), there is no evidence of significant, oxygen-depleted assemblages. Unlike these modern occurrences, there is evidence from the Tarfaya Atlantic Coastal Basin (Morocco) that, in the mid-Cretaceous, organic-rich sediments were being deposited on comparable shelves (Kuhnt et al. 1997, 2005; Tsikos et al. 2004; Kolonic et al. 2005; Poulton et al. 2015; Jenkyns et al. 2017).
In southern England, the Plenus Marls Member records changes in both the planktic and benthic assemblages of foraminifera (Jefferies 1962, 1963; Carter & Hart 1977; Jarvis et al. 1988; Hart & Leary 1989; Leary et al. 1989; Paul et al. 1999; Dodsworth 2000; Keller et al. 2001; Tsikos et al. 2004), but no confirmed oxygen-depleted assemblage. In the models provided by Jarvis et al. (1988, fig. 34) the distribution of foraminifera shows a reduction to a limited number of taxa, including two characteristic species (Gavelinella berthelini and Lingulogavelinella globosa), well-known in the chalk facies of the Anglo-Paris Basin. These two species are also ‘survivors’ in many other areas, though often known by different names (e.g. Gavelinella dakotensis).
The primary controls on benthic foraminifera in modern seas and oceans are the presence of O2 and nutrients (Gooday 1988; Moodley et al. 1998; Van der Zwaan et al. 1999; Friedrich 2010), with temperature and salinity of – possibly – lesser importance (Murray 1973; Van der Zwaan et al. 1999). In many of these accounts, sediment characteristics are not considered. In an area such as Plymouth Sound (Oxford et al. 2004, figs 6 and 7), clean, mobile sand-waves support hardly any in-situ living foraminifera while the more muddy, organic-rich sediments record the highest living population. This is, of course, the result of nutrient levels, but it is the sedimentary regime that is the ultimate control.
Food availability, O2 content and benthic foraminifera were included in the TROX-model of Jorissen et al. (1992). This indicated that, in shallower-water environments O2 was the important factor while nutrients were more critical in deeper-water settings.
A measure of benthic foraminiferal abundance (BFN – Benthic Foraminiferal Number) has become an important part of environmental assessment in both modern, and fossil, environments (Bernhard 1986; Erbacher et al. 1998; Holbourn et al. 1999; Van der Zwaan et al. 1999; Gebhardt 2006; Friedrich & Hemleben 2007). In both modern high carbon flux environments (Gooday 1994, 2003) and Cretaceous black shales (e.g. OAElb, Holbourn et al. 2001; Herrle et al. 2003a, b) BFN values increase with organic matter flux to the sea floor. However, this increase is reduced as the organic matter consumes O2 on the sea floor (Murray 1991). The reduced O2 levels on the sea floor have led to the suggestion that some species are able to ‘cope’ with the environment better; the so-called opportunistic species such as modern Stainforthia and Bulimina (Jorissen et al. 1992; Alve 1995) or mid-Cretaceous Osangularia and Gavelinella (Herrle et al. 2003a, b). In the OAEll record described by Jarvis et al. (1988, fig. 34), Gavelinella berthelini and Lingulogavelinella globosa were two of the ‘survivors’. In many Cretaceous OAE successions, a number of authors have identified Neobulimina, Tappanina and other praebuliminids as being capable of surviving in low-O2 conditions (Koutsoukos et al. 1990; Koutsoukos & Hart 1990a; Erbacher et al. 1998; Holbourn et al. 1999; Gebhardt 2006; Friedrich et al. 2009).
A similar assemblage (Gavelinella, Lingulogavelinella, Praebulimina, etc.) is reported within, and adjacent to OAEll in the Upper Eagle Ford Formation of Texas (Lowery & Leckie 2017, fig. 3). The associated radiolarians illustrated by Lowery & Leckie (2017, fig. 14) closely resemble, in terms of species and preservation, taxa reported by Koutsoukos & Hart (1990b) from the Brazilian Margin and their model for the distribution of the foraminifera (Koutsoukos & Hart 1990a, fig. 5) might well be applicable to the situation in Texas.
Many benthic foraminifera in low-O2 environments are often much smaller in size (e.g. Bernhard 1986). While this could be an adaption strategy to low-O2 conditions, it might also be a function of more rapid reproduction rates leading to an ‘apparent’ size reduction. It is noticeable that, in many areas, such events are also characterized by an increase in agglutinated taxa (Bernhard 1986; Koutsoukos & Hart 1990a; Coccioni & Galeotti 1993). In the Black Band there are a small number of black mudstones within the equivalent of the Plenus Marls Member, with a record of a limited, or agglutinated, assemblage (Hart & Bigg 1981). In the Carpathian Orogenic Belt, deeper-water assemblages record much higher numbers of, often simple (Ammodiscus, Glomospira, etc.), agglutinated foraminifera (Bąk et al. 2001, 2014; Bąk 2006, 2007a, b). Between these ‘extremes’ are numerous examples of mixed black, organic-rich, mudstones, inter-bedded with grey mudstones or limestones, with a mixture of benthic foraminiferal assemblages (Fisher et al. 2005; Gebhardt et al. 2010) some of which also include occurrences of radiolaria. Hart & Koutsoukos (2015, fig. 8) attempted to explain this by means of water depth, using evidence from the Brazilian margin (Mello et al. 1989; Koutsoukos & Hart 1990a, b; Koutsoukos et al. 1990, 1991, 1993) as a guide. In the offshore areas east of Brazil, the CTBE record of enhanced TOC simply adds to the continuing ‘anoxia’ of that silled basin that developed in the latest Aptian and Albian. In other areas, such as in the Black Band of Humberside, the Crimea (Fisher et al. 2005) and the Rehkogelgraben succession of central Austria (Gebhardt et al. 2010), there are black, organic-rich mudstones inter-bedded with the normal ‘oxygenated’ succession.
In other locations (e.g. Contessa Highway, Gubbio, Italy: see Hart & Koutsoukos 2015, fig. 2) the black, organic-rich mudstones occupy the whole of the positive stable isotope excursion and, in these locations, there are no intervening pale-coloured, carbonate-rich, horizons. There are, however, a number of locations where black, TOC-rich mudstones are known outside the carbon isotope excursion at the CTBE. In northern Germany, black mudstones are recorded inter-bedded with chalks in the Lower Turonian (Hilbrecht & Hoefs 1986) while in Texas Lowery & Leckie (2017) reported pre-OAEll organic-rich shales (of economic importance to the hydrocarbons industry). Lowery & Leckie (2017) attributed these organic-rich sedimentary rocks to upwelling generated by a significant transgressive event in the Cenomanian. In Oman, Wohlwend et al. (2016, fig. 9) recorded the presence of organic-rich mudstones (TOC >4%) in the Natih B Member of Jabal Qusaybah and other localities in the area immediately inland of the Oman Mountains. The locations where this occurs are in Central Oman (Wohlwend et al. 2016, fig. 2) in the middle of a carbonate platform/shelf, well away from a shelf edge where upwelling might occur. These occurrences of pre- and post-OAEll organic-rich sediments are, currently, little understood. If they are not related to local upwelling, can they be generated by sea-level rise ‘lifting’ the OMZ into the depositional areas, although the carbonate sediments in Oman do not appear to have been deposited in 200+ m water depth?
The occurrence of benthic foraminifera within what appear to be anoxic mudstones has created what is known as the ‘anoxic benthic foraminiferal paradox’.
Anoxic benthic foraminiferal paradox
The occurrence of benthic foraminifera in what are perceived as anoxic, organic-rich mudstones is problematic (Friedrich 2010). In the Black Band of Humberside, the occurrence of a small-sized, agglutinated assemblage of foraminifera is an example of the problem that has been reported quite widely (Koutsoukos et al. 1990; Erbacher et al. 1998; Holbourn et al. 1999; Gebhardt 2006; Friedrich et al. 2009). In attempting to explain some of the problems, Friedrich (2010) drew attention to an issue raised by Oxford et al. (2004, fig. 8) some years earlier. Mudstone layers, after deposition, are subject to compaction by <70%–80%. The time represented by a typical micropalaeontological sample is, therefore, quite considerable and may contain numerous oxic-dysoxic-anoxic ‘mini-cycles’. Such samples could yield a reduced assemblage, with representatives from a range of harsh to less harsh environments. This problem arose at Christian Malford (Wiltshire, UK) where, in a series of borehole cores, every sample of the organic-rich mudstones of the Phaeinum Subzone (Upper Callovian) of the Oxford Clay Formation yielded diverse, well-preserved assemblages of foraminifera, statoliths and otoliths (Hart et al. 2016) despite exquisite, soft-bodied, preservation of squid-like cephalopods that would imply sea-floor anoxia (Wilby et al. 2004, 2008). In both the Jurassic and Cretaceous examples, having mudstones containing benthic foraminifera finely inter-bedded with organic-rich, anoxic mudstones implies that the foraminifera are able to re-colonize the favourable environments almost immediately. If this requires sea floor migration of the benthic species then this could not be instantaneous within a large basin. Water depths during deposition of the Black Band would also be deeper than storm wave base and so the re-deposition of living individuals from suspension following storms would be impossible (see Hart et al. 2017).
In many other European successions there are several instances where, within the overall CTBE isotope excursion, there is a number of discrete black, mudstone, TOC-rich, reportedly anoxic, horizons. It is this aspect of OAEll that is further explored and related back to the successions on Humberside. Successions with black, organic-rich, beds within the carbon isotope excursion of OAEll are reported in the Crimea (Fisher et al. 2005; Dodsworth 2004) and Central Austria (Gebhardt et al. 2010). In both of these examples the black shales are almost devoid of foraminifera and there is no comparable assemblage of agglutinated taxa to that reported from the Black Band.
In the deeper waters of the Carpathian Mountains (Skole Nappe, Outer Carpathians, southern Poland) Bąk (2007a) and Bąk et al. (2001, 2014) recorded assemblages of radiolarian associated with deep-water agglutinated foraminifera (Rhizammina sp., Psammosiphonella sp., Ammodiscus sp., Recurvoides sp., Trochammina sp. and Gerochammina sp.). This setting is mid-upper bathyal and a much deeper-water setting than the successions in Austria (Gebhardt et al. 2010) and the Crimea (Fisher et al. 2005). When recorded against water depths (Fig. 6), the locations discussed in the text appear to indicate that the microfossil assemblages recorded are a function of depth and, presumably, dissolved oxygen. This also suggests that, in the Cretaceous Ocean, there is no evidence for oxygen recovery with depth below an expanded minimum zone.
Schematic interpretation of the potential water depths represented by the various successions discussed in the text. The model is based on the profiles developed by Koutsoukos & Hart (1990a) for the Atlantic Margin of Brazil. S.L., sea level; O.M.Z., postulated position of the oxygen minimum zone.
Summary
The Black Band that we see outcropping in Lincolnshire, Humberside and Yorkshire forms a distinctive horizon within the overall chalk succession. The high concentrations of dinoflagellates cysts recorded in the black mudstones appear to have been stimulated by enhanced nutrient supply from the intrusion and/or weathering of Large Igneous Provinces (in the Caribbean or High Arctic). This enhanced productivity removed dissolved oxygen from the sediments, creating a biological response and the occurrence of assemblages dominated by agglutinated foraminifera. The water depths of the various locations discussed here created the particular response that is recorded in the geological record and point to a general depletion of O2 in the mid-Cretaceous water column.
Acknowledgements
The author wishes to thank a number of colleagues for fruitful discussions on the problems associated with the Black Band and coeval strata in various parts of the world, including Haydon Bailey (Network Stratigraphic, Potters Bar, UK), Holger Gebhardt (Geological Survey of Austria, Vienna), Ian Jarvis (Kingston University, Kingston upon Thames, UK), Hugh Jenkyns (Oxford University, Oxford, UK), Eduardo Koutsoukos (Faro, Portugal), Mark Leckie (University of Massachussetts, Amherst, USA), Christopher Lowery (University of Austin, Texas, USA), Bruce Tocher (Stavanger, Norway), Michael Wagreich (University of Vienna, Austria), Helmut Weissert (ETH Zurich, Switzerland) and the late Christopher Wood (to whom this paper is dedicated). The two reviewers (Rory Mortimore and Christopher Paul) and editor are thanked for their invaluable assistance, as are the technical staff (Tim Absalom and James Quinn) at Plymouth University who provided some of the final figures.
Funding
This research received no specific grant from any funding agency in the public, commercial, or not-for-profit sectors.
Scientific editing by Stephen K. Donovan
- © 2018 The Author(s). Published by The Geological Society of London for the Yorkshire Geological Society. All rights reserved
