Palaeocene oceans and climate: An isotopic perspective

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Earth-Science Reviews 37 (1994) 225-252

Palaeocene oceans and climate: An isotopic perspective R.M. Corfield, Department of Earth Sciences, University of Oxford, Parks Road, Oxford OXI 3PR, UK Received 8 March, 1994; revised and accepted 29 August, 1994

Abstract The early Cenozoic was a time of climatic and oceanographic transition from the Cretaceous "Greenhouse" world to the "Icehouse" world of the Neogene. 6180 measurements shed light on ocean temperature and possible polar ice fluctuations during this interval, while 313C measurements monitor fluctuations in ocean productivity, deep water circulation and atmospheric CO 2. The major features from 6a80 analysis of the early Cenozoic are general cooling of surface waters, with some evidence for transient cooling across the K / T boundary. Surface water temperatures were at a Cenozoic maximum in the early Eocene, whereas deep waters cooled then warmed during the Palaeocene. The ~13C of bulk carbonates is at a minimum at the start of the Cenozoic due to the profound crisis in ocean surface water productivity associated with the extinctions of marine plankton at the Cretaceous/Tertiary boundary, thereafter 613C values increase (in bulk carbonates, as well as planktonic and benthonic foraminifera) to their Cenozoic maximum in the late Palaeocene (c. 60 Ma), after which time they again decrease over an interval of c. 4.5 m.y. to a Cenozoic low in the early Eocene. The increase in 613C values characteristic of the Palaeocene period is probably related to a combination of increasing surface water productivity and accelerated burial of organic carbon, conversely the decline in 613C into the early Eocene is probably related to a decrease in ocean productivity and a deceleration in the rate of organic carbon burial. Benthic 313C comparisons suggest that deep waters appear to have been predominantly formed in the high southern latitudes with the exception of a short lived interval near the P a l a e o c e n e / E o c e n e boundary possibly associated with a transient climatic anomaly (The "Late Palaeocene Thermal Maximum"). The thermal change of Palaeocene deep waters may be related to the changing productivity of surface waters by controlling atmospheric CO 2 flux in a similar way to that proposed for the control of the Pleistocene glacial cycles.

1. Introduction T h e P a l a e o g e n e was a time o f p r o f o u n d clim a t i c evolution from the ice-free w o r l d of the C r e t a c e o u s with its low l a t i t u d i n a l t e m p e r a t u r e g r a d i e n t s to the glacially d o m i n a t e d w o r l d o f the N e o g e n e with s t e e p l a t i t u d i n a l t e m p e r a t u r e gradients. A l t h o u g h essentially a g r a d u a l transition, it is now k n o w n that P a l a e o g e n e climatic evolution was p u n c t u a t e d by several steps ( Z a c h o s et

al., 1993) w h e n t h e o c e a n / a t m o s p h e r e a t t a i n e d a t r a n s i e n t a n d u n s t a b l e state (e.g. at the P a l a e o c e n e / E o c e n e b o u n d a r y ) . T h e activities o f the O c e a n Drilling P r o g r a m ( O D P ) a n d its p r e d e c e s sor, t h e D e e p S e a Drilling P r o j e c t ( D S D P ) over the last d e c a d e in p a r t i c u l a r has s e e n t h e recovery of s t r a t i g r a p h i c a l l y c o m p l e t e d e e p sea sedim e n t s p a r t i c u l a r l y in the h i g h e r l a t i t u d e s which, t h r o u g h t h e a p p l i c a t i o n o f stable i s o t o p e analyses has c o n t r i b u t e d g r e a t l y to an i m p r o v e d u n d e r -

0012-8252/94/$29.00 © 1994 Elsevier Science B.V. All rights reserved SSD1 0012-8252(94)00053-0

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standing of the climatic evolution of the early Cenozoic. Current studies of Palaeogene oceanography and climatology are concentrating in particular on. (1) sources of deep water during the Palaeogene, (2) the importance of atmospheric CO 2 in modulating Palaeogene climatic change, (3) reliably reconstructing sea surface temperatures, 4) quantifying the timing and amount of glacial ice particularly in the high southern latitudes and (5) identifying transient climatic events on timescales as short as 105 yr. This contribution reviews the use of the stable isotopes of oxygen (6180) and carbon (~13C) in reconstructing both long and short term trends in early Cenozoic oceanography and climatology. The focus of this contribution is on the interval between the Cretaceous/Tertiary boundary (K/T) and the termination of the "Palaeocene

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carbon isotope maximum" in the early Eocene. Fig. 1 shows the present location of all the sites discussed. The timescale of Berggren et al. (1985) has been used throughout this contribution. Before considering oceanographic and climatic changes during this period it is worthwhile to review the basis of 6]80 and ~3C stratigraphy and its application to deciphering environmental change on geological timescales.

1.1. Oxygen isotope fractionation Oxygen and carbon isotope analyses of carbonates have become widely used palaeoceanographic techniques in the last forty-five years. The method of palaeotemperature determination using the varying proportions of the light isotope of oxygen (]60) and the heavy isotope of oxygen (180) preserved in fossil carbonates was origi-

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Sites with Palaeocene sediments

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Fig. 1. Location map of sites discussed in this contribution. Symbols indicate sites which contain K / T boundary sections, Palaeocene sections, or P / E boundary sections.

R.M. CorfieM ~Earth-Science Reviews 37 (1994) 22.5-252

nated by Urey and colleagues at Chicago (e.g. Urey, 1947; Epstein et al., 1951, 1953). It was widely applied to the interpretation of Pleistocene glacial cycles by several workers in the 50's and 60's (e.g. Emiliani, 1954; 1955; Shackleton 1967) and by virtue of its ubiquitous imprint in the carbonate shells of benthic foraminifera stands today as the premier method for recognising glacial cycles in deep sea oozes and calibrating late Cenozoic time (Imbrie et al., 1984; Prell et al., 1986). The oxygen isotope method of palaeotemperature determination relies on the fact that the fractionation and hence ratio of 180/160 is dependent on temperature. An increase in temperature results in proportionally more of the light isotope of oxygen (160) being incorporated into the calcite lattice of both organically and inorganically precipitated carbonates, whereas a decrease in temperature results in proportionally more of the heavy isotope of oxygen (180) being incorporated. Stable isotope results are reported as a ratio of ~80/~60 using the conventional 6 notation to indicate deviation (in parts per thousand or %o) from the arbitrary PDB standard of 0 using this equation:

180 = ((180/160)sample- (180/160)standard)1000 ( 180/160)standard (1) The chief limitation on the use of 6180 variations to calculate palaeotemperatures in the geological past is that absolute palaeotemperature calculations rely on a knowledge of the isotopic composition of the water (6 w) from which the carbonate was precipitated. One of the most common causes of variation in the 6180 of seawater during geological time are the frequent intervals of glaciation that characterise the Earth's climatic history. Polar ice is formed from precipitation of waters which were evaporated at low latitudes. Water vapour formed by evaporation is enriched in 160 and the remaining ocean water therefore enriched in 180. Hence during glacial maxima, seawater and carbonates precipitated in it will have more positive 8180. The fact that this effect mimics the expected increase in 6~80 caused by a drop in ocean temperature during

227

glaciation means that it is not possible to discriminate the magnitude of the ice-volume effect on the 6180 signal except by the utilisation of independent evidence of ice-volume. The Palaeocene is generally considered to have been ice-free (Miller et al., 1987; Zachos et al., 1994) and hence the ice-volume contribution to the 8180 record during this interval is suggested to have been non-existent. Palaeotemperature calculations for the pre-middle Miocene interval therefore use an appropriately modified 6 w term of between - 0.9 and - 1.2%o. Additionally, fluctuations in 6 w occur on short timescales and in locally restricted areas due to variations in the ratio of evaporation to precipitation together with Rayleigh distillation and atmospheric vapour transport. Zachos et al. (1994) have noted that ocean surface waters show considerable variations even in areas far from fresh water input or away from enclosed basins. Ocean surface water 61SO can vary by as much 1.5%o (corresponding to a temperature variation of about 6°C) between high and low latitudes (according to the estimate of Broecker, 1989), while Hudson and Anderson (1989) have suggested that the areal range of 61So in the Holocene ocean is 2%o (corresponding to a temperature variation of about 8°C). These open ocean differences are mediated by the transport of isotopically light water vapour away from the sub-tropics towards the poles resulting in a poleward gradient of decreasing 61SO. It has been only recently that the importance of this effect on reconstructing palaeo sea surface temperatures (SST) has been appreciated. Zachos et al. (1994) have shown that by filtering fossil 6J80 data using an algorithm based on the Holocene distribution of 6180 they can much more accurately reproduce real palaeo-SST distributions. It is clear therefore that future Palaeocene palaeotemperature reconstructions will need to compensate for areal 6180 variability in this way. In near shore areas marine waters register more negative 6~SO clue to dilution with fresh waters (Craig and Gordon, 1965), while in enclosed basins such as the Mediterranean evaporation results in more positive 61~O (Thunell et al., 1987). It is because of these areal variations in

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the ratio of evaporation to precipitation that 61SO and salinity are correlated, and it is common in the literature to see 8180 measurements used as a proxy for salinity variations.

carbon production, but also (1) to monitor changes in oceanic PCO2 (and by extension, atmospheric PCO z) and 2) to map the movement of deeper ocean waters during geological time.

1.2. Carbon isotope fractionation

Changes in atmospheric CO 2 Carbon isotope ratio measurements in planktonic and benthic foraminifera (ideally from the same sample) can be used to monitor CO 2 variations by quantifying the intensity of the vertical carbon isotope gradient (A613C). The vertical carbon isotope gradient is a function of the biological carbon "pump" (Broecker and Peng, 1982; Kroopnick, 1985; Berger and Vincent, 1986) whereby 12C is extracted from surface waters and transferred to deeper waters by the photosynthetic fractionation of the two stable isotopes of carbon in conjunction with downward flux of dead particulate organic matter. This is subsequently returned to deeper water dissolved T C O 2 by respiration. This approach was pioneered by Shackleton et al. (1983) who showed that enhanced contrast between surface waters and deeper waters (i.e. increased A6~3C) was correlated with intervals of reduced atmospheric pCO2 in the late Pleistocene climatic cycles. The connection is (1) increased organic carbon production in surface waters (possibly caused by enhanced nutrient availability) leads to, (2) removal of carbon from the pool of surface water dissolved CO 2 as it is fixed in organic material, which leads to (3) a decrease in surface ocean pCO 2, which (4) draws CO 2 down from the atmospheric reservoir through re-equilibration (Fig. 2). Enhanced CO 2 fixation and removal is of course accompanied by a steepening in the vertical carbon isotope gradient. Hence a higher phosphorus to carbon ratio in glacial oceans as suggested by Broecker (1982) would lead to decreased atmospheric CO2 by increasing the availability of a factor which limits the rate of photosynthesis. In interglacials the mechanism is suggested to operate in reverse. Although this method was originally developed in the late Pleistocene (e.g. Shackleton et al., 1983; Toggweiler and Sarmiento, 1985) it has been applied in other intervals where benthic and planktonic foraminiferal data are available [e.g. Spicer and Corfield (1992) in the Cretaceous, and

In contrast to the temperature dependent fractionation of oxygen isotopes the fractionation of carbon isotopes is controlled by photosynthesis, with the result that organic material formed by carbon fixation is enriched in the light isotope of carbon (12C). In aquatic environments the surrounding water therefore becomes progressively depleted in ~2C with increasing rate of organic carbon production. Any carbonate precipitated from this water will therefore by relatively enriched in ~3C and have a more positive 6~3C. Hence analyzing carbonate sediments downcore yields a record of changing marine productivity. Interpreting the carbon isotope record at frequencies greater than the mixing time of the ocean (c. 1000 yr) will reflect the transient partitioning of carbon isotopes between different water masses (i.e. "internal" fractionation within the marine reservoir of dissolved inorganic carbon). On timescales longer than this, any change in 6t3Ccarbon~te reflects the fractionation of the two stable isotopes of carbon between carbon reservoirs (Berger and Vincent, 1986; Shackleton, 1987). This may be thought of as "external" fractionation. The major carbon reservoirs are: (1) the ocean and atmosphere, (2) the lithosphere and (3) the biosphere. Both internal fractionation of carbon isotopes and external fractionation may contribute to excursions in the 613Ccarbonate signal. Internal fractionation mechanisms, if maintained for periods longer than the mixing time of the ocean, and in the absence of negative feedback effects, may contribute to external fractionation. An example is the case of the Palaeocene carbon isotope maximum (Corfield and Cartlidge, 1992) considered in more detail below, where enhanced surface water productivity may have led ultimately to an increase in the burial rate of organic carbon. The study of 6~3C variations can be used not only to monitor variations in the rate of organic

R.M. Corfield ~Earth-Science ReL,iews 37 (1994) 225-252

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(i.e. species-specific non-equilibrium fractionation) are trivial, or otherwise understood, and appropriately adjusted for. Since deep waters become isotopically more negative with time after leaving the surface in the area of their formation due to the rain of organic material that progressively adds ]2C to the water mass, comparison of sites of similar palaeodepth in different areas can be used to suggest the location of deep water sources (e.g. Curry and Lohmann, 1982, 1983; Duplessy et al., 1984; Shackleton et al., 1984; Miller and Fairbanks, 1985; Woodruff and Savin, 1989) and the pathway taken by deep waters as they journey from their source areas. This technique has been applied recently to Palaeogene sediments because of suggestions (e.g. Brass et al., 1982; Bralower and Thierstein, 1984; Kennett and Stott, 1990, 1991; Zachos et al., 1993) based on the ideas of Chamberlin (1906) that deep waters may have been formed by evaporative processes at low latitudes inducing salinity driven (halothermal) circulation, as opposed to the temperature driven (thermohaline) circulation characteristic of the late Neogene.

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Fig. 2. Schematic of the oceanicvertical carbon isotope gradient showing the relationship between increasing productivity, intensity of the sub-surface oxygenminimum and marine and atmospheric CO2. A = low productivityocean, B = high productivity ocean. Corfield and Cartlidge (1992) in the Palaeocene, see below]. Tracing deep water circulation Carbon isotope measurements in benthic foraminifera can, in principal, be used to trace deep water circulation. The basis of this technique is that the 613C composition of benthonic foraminifera must reflect the isotopic composition of surrounding water at the time of calcification provided that the influence of vital effects

"Vital effect" is the term used in the context of organically precipitated calcites when the 613C or 6~SO ratio differs from that of the surrounding water because of the influence of the metabolism of the secreting organism. It is clearly difficult to establish the presence of vital effects without parallel measurements of the isotopic content of the carbonate and the water from which it is precipitated (Erez and Honjo, 1981; Erez and Luz, 1983). It is equally clear that this cannot be done using fossil material. Hence any suspicions of non-equilibrium fractionation by organisms in the geological past can most easily be tested verified by using a uniformitarian approach. It has been suggested that species of foraminifera coexisting with photosymbionts show higher ~ 3 C values than species without photosymbionts (Spero and Williams, 1988, 1989; Spero et al., 1991). Apparently algal commensals extract sufficient ~2C to locally enrich the surrounding dissolved bicarbonate from which the shell is precip-

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R.M. Corfield / Earth-Science Ret,iews 37 (1994) 225-252

itated in ~3C. This is discussed further below with reference to Palaeocene planktonic foraminifera.

shown that a rapid ( < 1000 yr) negative inflection in 613C of up to 2%o is characteristic of this era boundary in both unlithified marine environments, where foraminifera and bulk sediment are typically analysed, (e.g. Boersma and Shacldeton, 1981; Shackleton and Hall, 1984; Zachos and Arthur, 1986; Stott and Kennett, 1990) and in marine and terrestrial limestone sections (e.g. Scholle and Arthur, 1980; Hsfi et al., 1982; Perch-Nielsen et al., 1982; Williams et al., 1983; Renard, 1986; Corfield et al., 1991; Corfield and Cartlidge, 1992) where whole-rock samples are analysed. Figs. 3 and 4 shows a comparison of K / T boundary 6~3C records from several deep sea and marine outcrop sites. Also included is the

2. The Cretaceous/Tertiary boundary 2.1. Carbon isotope change across the Cretaceous/ Tertiary boundary The dramatic 6~3C minimum at the K / T boundary was originally noted by Brennecke and Anderson, (1977). Scholle and Arthur (1980) further noted this excursion at the K / T boundary in terrestrial outcrops from several continents in whole-rock samples. Many studies since have

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R.M. CorfieM~Earth-Science Reviews 37 (1994) 225-252

position of the K / T boundary based on nannofossil, foraminiferal and iridium abundance evidence. The carbon isotope decline at the K / T boundary has been linked (Hsii and McKenzie, 1985; Hsii, 1986; Zachos and Arthur, 1986; Zachos et al., 1989a) to a collapse of the vertical carbon isotope gradient (Fig. 5). Cessation of photosynthesis would rapidly lead to surface water carbonate recording ~3C values similar to those of deep waters since surface water dissolved CO 2 would no longer be enriched in t3c. Although this collapse in AS~3C has not been found at all sites (Stott and Kennett, 1989) the widespread occurrence of the 3~3C minimum at the K / T boundary

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implies that a radical reduction in marine primary productivity was at least a near-global phenomenon (Zachos et al., 1985, 1989a; Zachos and Arthur, 1986; D'Hondt and Lindinger, 1988; Keller and Lindinger, 1989; Barrera and Keller, 1990). With the decrease in sunlight supposedly caused by the spreading dustcloud from the agent of the K / T extinctions, be it volcanism (e.g. McLean, 1985; Officer and Drake, 1985; Courtillot et al., 1990) or bolide impact (e.g. Alvarez et al., 1980) marine food chains were severed initiating the onset of the "Strangelove" ocean (Hsii and McKenzie, 1985) characterised by low marine productivity. The modelling experiments of Kump (1991)

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R.M. Corfield ~Earth-Science Reuiews 37 (1994) 225-252

232

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suggest that the extinctions at the K / T boundary led to isotopic homogenization of the surface layer of the ocean on timescales as short as hundreds of years. If, as suggested by Zachos et al. (1989a) the "Strangelove" ocean persisted for several hundred thousand years, then the isotopic composition of the surface ocean would have approached that of the weathering input from rivers. More recently, lvany and Salawitch (1993) have pointed out that in six out of eight D S D P / O D P sites where vertical carbon isotope gradients across the K / T boundary have been measured, the gradient does not merely decline to zero but actually becomes negative with benthic values heavier than planktonic values. They argued that this can only be because of the addition of isotopically light carbon from a source outside the ocean TCO 2 reservoir because it did not affect benthic ~13C and therefore occurred on timescales shorter than the mixing time of the ocean.

This suggestion has been challenged recently by Keller and Macleod (1993) who suggested that most DSDP and ODP K / T boundary sections have missing sediment across the boundary. Their objection rests on the requirement by Ivany and

Salawitch (1993) that the collapse of the vertical ~13C gradient be shorter than one mixing time of the ocean (c. 1 kyr). Keller and Macleod (1993) argue that these sections do not permit resolution of an event of this duration because of widespread hiatuses, hence the hypothesis of addition of ~2C to the surface ocean cannot be correct. In addition D'Hondt (pers. commun.) has pointed out that the Ivany and Salawitch hypothesis may well be ruled out by the very data that it claims to interpret. Specifically, Ivany and Salawitch assumed that the negative gradient lasted for only 1 kyr. D'Hondt has calculated that about 10% of the terrestrial biomass must have been added to the surface ocean in a single mixing cycle to maintain a ~13C of only - 1%c (assuming a 13Ct~re~tri~lbi,,m~.~ of --10%c and a perfect Strangelove ocean with zero export of marine Corg to deep water). Existing isotope data show that a -2%0 negative gradient occurred over a 500 kyr interval at DSDP 356 (Zachos and Arthur, 1986), and a -0.5%~ negative gradient occurred over an interval of a million years at DSDP Site 465 (Boersma and Shackleton, 1981). Given the magnitude of these negative gradients, the implication from these simple calculations is that the terrestrial biomass reservoir would have to have been completely consumed in a much shorter time than the measured duration of the negative ~13C gradients at these sites. Hsfi and MacKenzie (1990) had previously noted the same reversed carbon isotope gradient as Ivany and Salawitch (1993) and proposed a different explanation (Fig. 5). They suggested that reversed vertical carbon isotope gradients are generally possible during "Strangelove" events (e.g. Cretaceous/Tertiary boundary, Permo/ Triassic boundary, P r e c a m b r i a n / C a m b r i a n boundary) because the surface waters of the ocean switched from a net "production" mode to a net "respiration" mode. Their scenario for the K / T boundary is that the blooms of the calcareous dinoflagellates Braarudosphaera and Thoracosphaera (Thierstein, 1981; Perch-Nielsen et al., 1982) which are characteristic of the earliest Danian were consumed by bacteria which occupied the niches formerly inhabited by grazers in the food chain. These bacteria fed on the newly

R.M. Corfield ~Earth-Science ReL,iews37 (1994) 225-252

formed photosynthetic organic matter releasing 12C enriched CO2 to surface waters. In the absence of significant photosynthesis the dominant isotope fractionation effect in surface waters would thus favour enrichment of 12C in surface dissolved bicarbonate which would register as a reversed vertical carbon isotopic gradient when measured via carbonate precipitated in planktonic and benthonic foraminifera. A further explanation for the negative 613C shift at the K / T boundary in foraminiferal and whole-rock samples as well as the negative gradient seen at the K / T boundary has been proposed by D'Hondt and Zachos (1993). These authors have shown that the smallest size fractions of the earliest Palaeocene planktonic foraminifera are generally significantly depleted in ~3C and consequently measurements of surface water 6~3C based on such species could significantly underestimate the magnitude of the vertical carbon isotope gradient. This effect may be exacerbated by the presence of the calcareous dinoflagellate blooms which (together with coccoliths) may depart from 13C/12C equilibrium by as much as 6%~,. This of course further implies, that where smaller planktonic foraminifera, calcareous dinoflagellates and coccoliths contribute to whole rock samples (i.e. in almost all cases where a measurement is technically possible) there is a risk of recording anomalously negative 6~3C valDes.

Finally, it is rather difficult to see from first principles how the cessation of the biological carbon "pump" which maintains the vertical carbon isotope gradient could lead to negative values on its own accord. It is more likely that the surface to deep contrast would become zero.

2.2. Oxygen isotope change across the Cretaceous/ Tertiary boundary Considerable attention has been focused on the potential of any systematic oxygen isotopic change at the K / T boundary to shed light on climatic change as a possible mechanism for the K / T extinctions. Because of the high potential for diagenetic alteration of 6180 ratios, unambiguous temperature changes across the K / T

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boundary are hard to verify. While Zachos and Arthur (1986) concluded that there were no significant fluctuations in either benthic or planktonic 6180 across the K / T boundary, several other authors have noted K / T boundary sections that show an 180 enrichment in the immediate aftermath of the boundary extinctions. Sections where this increase in 6180 have been found include Caravaca in Spain, Biarritz in Switzerland, Lattengebirge in Germany and Stevns Klint in Denmark (Perch-Nielsen et al., 1982). Zachos et al. (1989b) also showed that a similar increase occurred in molluscan carbonate across the boundary at Braggs (Alabama), while Corfield et al. (1991) found evidence for an increase in 6180 in whole-rock samples from the Bottaccione Gorge and Contessa Highway sections in Umbria (Italy). Keller and Lindinger (1989) inferred a bottom water cooling of 4-5°C at the outer neritic section of E1 Kef (Tunisia). Schmitz et al. (1992) have also noted that bottom water temperatures were about 1.5°C cooler above the K / T boundary at the Stevns Klint section in Denmark. In the deep sea Shackleton and Hall (1984) showed an increase in 6180 of bulk carbonate in DSDP 527 (Fig. 6), while Zachos et al. (1989a) showed an 1SO enrichment in DSDP 577. Smit (1990) documented a 31~O decrease in bulk carbonate records in the Agost and Caravaca sections in Spain and interpreted it as an 8°C warming of surface waters. Barrera and Keller (1990) showed an inferred warming (based on decreasing 6180 values) across the K / T boundary in the Brazos River section but this trend is almost certainly the restdt of diagenetic alteration of the original signal since 613C and 6180 significantly co-vary between samples. It remains a problem that the 613C and 61SO study of the many different K / T boundary sections is based on a variety of carbonate materials (e.g. whole or bulk rock, different species of planktonic and benthonic foraminifera in differing degrees of preservation and with potentially different degrees of non-equilibrium fractionation) (Stott and Kennett, 1989). This is almost certainly the reason for the differing estimates of the magnitude of the 613C and A613C changes across the boundary. However, several authors

R.M. CorfieM/ Earth-Science Reviews 37 (1994) 225-252

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have now analysed several carbonate materials across the boundary in an attempt to minimise these differences. Schmitz et al. (1992) analysed size-controlled, well-preserved monospecific foraminifera across the K / T transition at Stevns Klint and concluded that the palaeotemperature history of the K / T boundary ocean was complex and unstable, probably as a result of changing ocean circulation. At present the balance of evidence (based on the number of sections showing whole-rock 6180 increase as well as those sections where well-preserved benthic foraminifera have been analysed) appears to suggest that there is a minor cooling across the K / T boundary. This may well be consistent with scenarios of global darkness resulting in decreased solar flux to the surface of the Earth.

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foraminiferal calcite (Fig. 7) showed a profound

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Fig. 6. The K / T ,5J80 increase ("cooling") in DSDP Site 527 (data from Shackleton and Hall, 1984).

increase such that the Palaeocene period is characterised by the most positive 61-~Cvalues of the Cenozoic (Shackleton and Kennett, 1975; Shackleton et al., 1985a, b; Shackleton, 1986, 1987). However this "gradual" increase over a period of four million years is punctuated by variations in the rate of 6a3C increase and the intensity of the vertical carbon isotope gradient (A6~3C). Following the broad two million year maximum of the late Palaeocene, ~13C values declined. This "gradual" decrease is punctuated by at least one, significant short term event at the Palaeocene/ Eocene boundary between 57.5 Ma and 58 Ma. Before considering the significance of these short term superimpositions on the structure of the Palaeocene and early Eocene carbon isotope curve, it is essential to consider the implications of the broad structure of this pattern. In essence the Palaeocene carbon isotope maximum yields three avenues of oceanographic and climatic implication that require exploration: (1)productivity and carbon reservoir changes and the intensity of

R.M. Corfield / Earth-Science Reuiews 37 (1994) 225-252

235

data sets (e.g. Oberh~insli and Toumarkine, 1985; Miller et al., 1987). A recent, relatively detailed parallel compilation of planktonic and benthic 6180 and 613C data is that of Corfield and Cartlidge (1992). Carbon isotope data from Palaeocene and early Eocene surface dwelling planktonic foraminifera (Morozouella, Acarinina), thermocline dwelling planktonic foraminifera (Subbotina) and benthonic foraminifera (Nuttallides, Cibicidoides) from three areas they studied are shown in Fig. 8. Like bulk isotope data these data show that foraminiferal ~13C values increase in all three depth groups from the early Palaeocene (c. 64 Ma) until they peak in the late Palaeocene at about 60 Ma. Values then decrease reaching a minimum between 55 and 56 Ma. The amplitude

the dissolved sub-surface oxygen minimum zone, (2) sources of deep water and (3) potential atmospheric carbon dioxide flux and consequent climatic change.

3.1. Productiuity and carbon reseruoir change in the Palaeocene At present few outcrop sections or deep sea drilling sites have sufficiently high resolution parallel benthonic and planktonic data to provide a detailed picture of the palaeoceanography of the Palaeocene period. Bulk carbonate ~13C stratigraphies through the Palaeocene are far more common (Shackleton et al., 1985a; Shackleton, 1986) while several authors have presented relatively coarse resolution planktonic and benthonic

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.... BENTHICS Fig. 8. ~13C change in surface dwelling planktonic, thermocline dwelling planktonic and benthonic foraminifera during Palaeocene time at DSDP 577, a composite succession of the Leg 74 sites (DSDP Sites 525 and 527) and a composite of Leg 113 sites (ODP sites 689 and 690). Data from Shackleton et al. (1984), Stott et al. (1990) and Corfield and Cartlidge (1992). Redrawn from Corfield and Cartlidge (1992).

236

R.M. Corfield ~Earth-Science Rez,iews 37 (1994) 225-252

of the @~3C maximum varies between the different depth groups, the data indicate a 1.4%o increase in 613C in benthonic foraminifera in the Leg 86 succession and a 1.7%o increase in the Leg 74 succession. The @13C increase in thermocline dwellers (Subbotina) was between 1.5 and 2.00%o in the Leg 86 and the Leg 74 successions. The increase in surface water (Morozovella/Acarinina) g13C was close to 3.0%o at these sites. The long-term decline in benthonic 613C values in the latest Palaeocene/early Eocene was 2.5%o in the Leg 86 succession, 2.4%o in the Leg 74 succession and 3.2%o in the Leg 113 succession. Thermocline values decrease by 2.8%o in the Leg 86 succession, 3.6%o in the Leg 74 succession and 3.4%o in the Leg 113 succession. The decline in surface water ~3C was 3.4%o in the Leg 86 succession, 3.0%0 in the Leg 74 succession and 3.6%o in the Leg 113 succession. Three features of the Palaeocene carbon isotope maximum are highlighted by comparison of these data: (1) Benthonic @~3C values broadly mirror the increase and subsequent decrease in intermediate and surface water @13C values (although benthic @~3C change appears to lag planktonic 6L~C change), (2) the differences between the three depth monitors increase as @13Cvalues become more positive, (3) the 613C decline from the late Palaeocene into the early Eocene was of greater amplitude than the @~3C increase from the early Palaeocene into the late Palaeocene. The increase in benthonic @~3C values, together with differences in the amplitude of the @J3C signal between the different depth groups, suggest that the Palaeocene carbon isotope maximum was composed of both an internal carbonisotope fractionation effect (i.e. fractionation within the oceanic reservoir of total dissolved carbon) and an external carbon-isotope fractionation effect (i.e. exchange between carbon reservoirs, for example, the ocean and the sedimentary carbon reservoir). Changes in global benthonic ~3C best monitor external fractionation effects since the average chemistry of the ocean is better represented close to the seafloor rather than at the surface (Shackleton, 1985). Accordingly, the increase and subsequent decrease in benthonic 61~C during the Palaeocene suggests that isotopi-

cally light carbon moved from one reservoir (the ocean-atmosphere system) to another, (probably the lithosphere) reaching a maximum transfer rate during the late Palaeocene, and that this light carbon was subsequently either returned to the ocean in the early Eocene, or that the rate of ~2C output from the ocean relative to 12C input decreased markedly. Further evidence for the ubiquity of large scale carbon isotope fractionation during the Palaeocene is that the long term 313C decline from the late Palaeocene into the early Eocene, as well as the transient @~3C decrease at the Palaeocene/Eocene boundary (see below) have been located in terrestrial carbonates (soil nodules and herbivore tooth enamel) from the Big Horn Basin in the continental USA (Zachos et al., 1991; Koch et al., 1992). The evidence for internal fractionation during the Palaeocene ~3C maximum is the difference in the amplitude of the surface ~3C and benthonic iSt3C marine isotope curves. The relative increase in surface water 6J3C is due to enhanced photosynthesis in surface waters which depletes surface water TCO 2 in ~2C and is therefore expressed in the shells of surface dwelling planktonic foraminifera such as Morozouella (most common in the low latitudes) and Acarinma (the most abundant surface dweller in the high latitudes). Enhanced photosynthesis in surface waters implies an intensification of the sub-surface oxygen minimum zone as the excess organic material is oxidised e.g. Shackleton et al. (1985a). A more sensitive proxy for the intensity of the subsurface oxygen minimum is the vertical carbon isotope difference or gradient (A613C) between surface and d e e p e r dwelling planktonic foraminifera or between surface dwelling planktonics and benthonic foraminifera. The presence of an expanded oxygen minimum in the late Palaeocene of DSDP Site 577 was originally inferred (Shackleton et al., 1985a) because A6~3C between Morozot~ella and Subbotina increased as surface water 6J3C values became more positive during Palaeocene time. This finding was challenged (Miller et al., 1987) on the basis of supposedly unchanging A613C differences when measured between Morozo~,ella and benthonic foraminifera which, it was suggested,

237

R.M. CorfieM~Earth-Science Ret,iews37 (1994) 225-252

indicated that the oxygen minimum was not expanded during the period of heavy ~13Cvalues in the late Palaeocene. The more recent estimates of Corfield and Cartlidge (1992) based on a larger data set from more sites and reproduced here in Fig. 9, do after all, indicate systematic changes in vertical A~]3C in DSDP Site 577, the Leg 74 and the Leg 113 sites from the beginning of the Palaeocene to the 313C minimum in the early Eocene. In these three regions there is a positive correlation between A/~3C and surface 613C values, whether the vertical carbon isotope gradient is monitored between surface and thermocline dwelling planktonic foraminifera, or between surface dwelling planktonic foraminifera and benthonic foraminifera. This increase in A~13Cwith increase in surface water ~13C supports the original hypothesis of Shackleton et al. (1985a) of a global intensification of the sub-surface oxygen minimum. As Fig. 8 shows, maximum A613C values are contemporaneous with the peak in surface water 6~3C (a composite monitor of ocean surface-water

Shatsky Rise, North Pacific Ocean

productivity as well as whole ocean 13C) in the three locations. Thereafter these steep gradients decline. In DSDP Site 577 the A~13C gradient between Morozouella and Subbotina has declined to less than 0.2%o after about 55 Ma while in the Leg 74 sites it was reduced to about 1%o. In the Leg 113 sites the vertical carbon isotope gradient also diminished as surface 613C values declined in the early Eocene but thereafter a substantial A~13C gradient (about 1.5%o) was maintained. Thus data from Leg 74 sites and Leg 113 sites suggest that after a relatively short-lived productivity decline in the early Eocene productivity re-established itself. The interval of low productivity was apparently longer in the low latitude site 577. R e a e t al. (1990) have suggested that the longterm 613C decline across the P a l a e o c e n e / E o c e n e boundary was caused by input of ~2C from the land due to enhanced run off induced by elevated rainfall. Additionally, Thomas (1991) has implicated volcanism in contributing to the transient P / E boundary negative 313C excursion. How-

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238

R.M. Corfield~Earth-Science Reviews37 (1994) 225-252

ever, the decline in A613C in either event would not be predicted directly by simple influx of isotopically light carbon from either the continental lithosphere or from mantle outgassing, although both would probably contribute to the negative shifts noted in benthonic ~13C. Nevertheless, weathering of organic carbon from the continents a n d / o r mantle degassing could affect marine carbon isotope gradients by altering nutrient flux and therefore productivity. Another possible cause of marine productivity change is that proposed by Rea et al. (1990) who suggested that the marked reduction in zonal wind intensities noted by Janacek and Rea (1983) and Rea et al. (1985) caused a decrease in surface water productivity through reduced vigour of surface water circulation. Other evidence for an expanded oxygen minimum zone in the late Palaeocene is the presence of black or grey-green shales in the Atlantic at Sites 385 and 387 (Bermuda Rise) which have been interpreted as evidence of sluggish deep water circulation (Tucholke and Vogt, 1979). Pak and Miller (1992) have suggested that these sediments indicate a reduction of deep water ventilation in the early Palaeocene and dysaerobism on the Bermuda Rise between Planktonic foram. Zone P3 (latest early Palaeocene) and Nannofossil Zone NP10 (latest P a l a e o c e n e / e a r l i e s t Eocene). Furthermore, the North Sea basin is considered to have been dysaerobic at this time (Knox, pers. commun., 1991), although this may only have been because of tectonic isolation (B. Schmitz, pers commun., 1994). Kaiho (1991) has monitored deep-ocean dysaerobism using benthonic foraminiferal morphology. His data suggest (Kaiho, 1991; Fig. 10) that there may have been a decrease in deep ocean dissolved oxygen during the late Palaeocene, with higher levels of dissolved oxygen in the early Palaeocene and early Eocene. This would agree with the suggestion of late Palaeocene oceanic dysaerobism inferred on the basis of 613C gradients (Shackleton et al., 1985b; Corfield and Cartlidge, 1992). Another lithospheric sink for organic carbon is the continental lithosphere. There is Palaeocene coal in the Powder River basin of the continental United States (Flores et al., 1989) as well as

Palaeocene oil in the Pool Creek Shale of Alaska (T. Dill, pers commun., 1988). These reserves of t2C would also contribute to high 6 13Cbenthonic values. Hence there is both marine and terrestrial sedimentary evidence that might account for the elevated ~3C of late Palaeocene benthonic foraminifera. 3.2. Intermediate~deep water circulation in the Palaeocene Ocean

Water masses are defined by temperature and salinity and their depth therefore varies depending upon distance from their source area. As an example, in the Holocene, deep water is nearly at the surface in the North Atlantic and Southern Ocean whereas the base of Antarctic Intermediate Water shoals from about 2000 m in the Southern Ocean to about 1500 m in the North Atlantic. There can be substantial differences in CO2, alkalinity, carbonate ion saturation state, and nutrient content between water masses depending upon their relative age, and partly because of this, distinctions between water masses can vary between authors. For example, Emery and Meincke (1986) define vertical water masses as follows: Upper waters (including thermocline waters): 0-500 m; Intermediate waters: 500-1500 m; Deep and Abyssal Waters: 1500 m-Bottom. Others distinguish Upper and Lower deep water with the interface at about 2000-2500 m. This convention is followed in this contribution with the result that most of the sites discussed herein can be considered upper deep water to lower intermediate water. The sites discussed are limited to this depth range because of the necessity of analysing the ~3C of foraminifera that are as well preserved as possible [i.e. have never been substantially below the calcite compensation depth (CCD)]. It is of course difficult to make specific assertions about Palaeocene water mass structure in the absence of direct measurement of the characteristics that are used to define Holocene water masses. Hence we do not know with any confidence where the precise intermedia t e / deep water transition was in the past. In the present day and for much of the Neogene deep ocean waters have been formed at

R.M. Corfield / Earth-Science Reviews 37 (1994) 225-252

high latitudes (Woodruff and Savin, 1989; Wright et al., 1991, 1992). Northern Component Waters (NCW) are formed in the Norwegian Sea while Southern Ocean Waters (SOW) are formed in the Weddell Sea by the overturn of surface waters as they cool during the winter, becoming denser and subsequently sinking. This mode of deeper water formation powers the thermohaline system of deep water circulation. Carbon isotopic analyses and faunal analyses of benthic foraminifera assemblages suggest that deep waters have been formed in the Southern Ocean throughout the Cenozoic (e.g. Barrera et al., 1987; Miller et al., 1987; Wright et al., 1991, 1992; Thomas, 1992). During the Holocene, SOW flowed into the South Atlantic, Indian and Pacific Oceans where it either aged as in the Pacific or mixed with NCW in the Atlantic and Indian Oceans. During Cenozoic time deep waters have also been formed in both the North Atlantic and the Mediterranean. Significant NCW probably did not form until about 19 Ma when deep Atlantic S~3C first became enriched relative to 613C in the deep Pacific (Wright et al., 1991). Most authors however consider the first NADW to form in the M. Miocene at about 12.5 Ma since the Greenland-Scotland ridge was too shallow before 12.5 Ma to permit outflow of Norwegian Sea water. Hence the developing Atlantic-Pacific contrast between 19 Ma and 12.5 Ma may be due to Tethyan outflow or some other factor (Norris, pers. commun.). Prior to the middle Miocene, the only significant source of deep or intermediate waters other than SOW was probably the Mediterranean and shallow epicontinental seas of Tethys (Brass et al., 1982; Kennett and Stott, 1990). These waters would have formed deep waters by a evaporation driven density increase. This type of salinity related deep water circulation is termed halothermal. There are precedents for this mode of deep water circulation albeit on a smaller scale. For example, at present some intermediate depth waters in the Indian Ocean are formed in the Arabian Sea and the Red Sea by evaporative processes (Wyrtki, 1973; Zachos et al., 1992). Although evidence for the sources of Paleocene intermediate deep waters is incomplete,

239

Miller et al. (1987) on the basis of oxygen and carbon isotope differences between benthonic foraminifera in DSDP Sites 577 and 524, suggested that the Cape Basin was supplied by deeper waters formed in the high southern latitudes in the late Palaeocene with reduced southern deep water production in the early Eocene. Katz and Miller (1991) had previously suggested that the Southern Ocean was the source of deep waters during the late Palaeocene (60-58 Ma) and the early Eocene (57-52 Ma). Two more recent studies of benthic $13C distributions are in broad agreement with these findings. Corfield and Cartlidge (1992) suggested on the basis of a comparison of ~13Cbenthic data from DSDP 577 (equatorial Pacific), a composite section based on the Leg 74 sites (south Atlantic), and ODP 690 (Weddell Sea) that intermediate deep waters were formed predominantly in the high (southern) latitudes during the Palaeocene with the possible exception of the Palaeocene/Eocene boundary interval when intermediate deep water production slowed or halted in this region, and a contribution from the low latitudes became more important. Pak and Miller (1992) also suggested that intermediate deep waters were formed in the high latitudes during the Palaeocene, and that at the P / E boundary Southern Ocean intermediate and upper deep water production shut-down and was replaced by production in the low latitudes, thereby giving rise to a halothermal circulation system. They further suggested that following the P / E boundary, the early Eocene was characterised by production of intermediate deep waters both in the high southern latitudes as well as in the low latitudes, differing in this latter point with Corfield and Cartlidge (1992). Fig. 10 summarises the benthic 6~3C data of several authors including the most recent studies. Additional data is from Barrera and Huber (1991) for the high southern latitude site ODP 738C and from Seto et al. (1991) for the high southern latitude site ODP 752. The benthic 6a3C data plotted in Fig. 10 suggest that the history of Palaeocene intermediate deep water circulation may be summarised as follows:

R.M. Corfield / Earth-Science Reviews 37 (1994) 225-252

240

(1) During the early Palaeocene (65.8-63.4 Ma) the heaviest values are from the high southern latitudes (ODP 738C, Prydz Bay; ODP 752 Kerguelen Plateau) and the mid-latitudes (Walvis Ridge, Southeast Atlantic). North Pacific DSDP Site 577 exhibits the lightest (hence oldest) ~3C values. This, together with a maximum interbasinal ~ ~3C gradient of c. 1%o suggests that the high southern latitudes were the predominant source of intermediate deep waters during this interval. (2) Between 63.4 and 61.6 Ma the 313C of benthonic foraminifera in DSDP Site 577 increased and became more positive than coeval values from the other sites. This may signify a decrease in high southern latitude outflow since Pacific sites are most likely to register the mean 313C content of the deep ocean. It may also suggest that a low latitude source became predominant during this interval although at present we have no site close to the potential outflows

from putative low latitudes sources (such as Tethys) to test this hypothesis. The interbasinal isotopic gradient also decreases to c. 0.5%0 during this interval. (3) Between 61.6 and 58.4 Ma interbasinal gradients are at their Palaeocene maximum (c. 1.5%o). The high southern latitudes (Leg 113 sites, Weddell Sea) are most positive, with Leg 74 values only about 0.2%o more negative. The lightest ~t3c values occur in the equatorial Pacific (DSDP Site 577) although the high southern latitude site ODP Site 752 registers even more negative values. However, it has been suggested (Zachos et al., 1992) that values from ODP Site 752 may be affected by diagenesis which tends to result in a negative shift in ~13C. (4) Between 58.4 and 58.0 Ma ~3C values from the Leg 113 and Leg 84 composite sections and DSDP Site 577 converge. Maximum interbasinal isotopic gradients are about 0.5%o during this

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R.M. CorfieM~Earth-Science Rev&ws37 (1994) 225-252

interval. This may be due to a progressive decrease in the importance of the high southern latitudes as a source area for intermediate deep waters, and may be a prelude to the Palaeocene/ Eocene boundary event. (5) Between 58.0 and 57.6 Ma 613C values in the Leg 113 composite section are lighter than values in the other sections. This may indicate complete shutdown of high southern latitude intermediate deep water formation possibly accompanied by production of water by evaporative processes in the low latitudes (Kennett and Stott, 1990; Thomas, 1991; Corfield and Cartlidge, 1992; Pak and Miller, 1992; Zachos et al., 1993). The implications of the Palaeocene/Eocene boundary event are discussed in more detail below. (6) Between 57.6 and 56.3 Ma Leg 113 values are once again the most positive while Leg 74 and DSDP Site 577 values are similar. This again suggests overall high latitude intermediate deep water production. (7) Between 56.3 and 52.4 Ma the record from DSDP Site 577 is isotopically lightest but Leg 74 data become more positive than data from the Leg 113 sections. This appears to agree with the suggestion by Pak and Miller (1992), that the early Eocene interval was characterised by the production of intermediate deep waters both in the high southern latitudes as well as in the low latitudes. Zachos et al. (1992) compiled benthic 613C data from the Indian Ocean and suggested (again by comparisons with DSDP Site 577 and ODP Site 690B) that Indian Ocean waters were older than those of the Antarctic but similar or younger than the intermediate/upper deep waters of the central Pacific. They interpreted these data as suggesting a Southern Ocean deep water source during the late Palaeocene and early Eocene. In the early part of the late Palaeocene (stage 2 on Fig. 10) this finding is at variance with the new compilation presented here, where 613Cbenthonic data from DSDP Site 577 become significantly more positive than data from ODP 738C and ODP 752 suggesting the possibility of deeper water production in the low latitudes or reduced deeper water production in the high latitudes. However, from the middle part of stage 3 their

241

interpretation for the latest Palaeocene and early Eocene is in broad agreement with those of Katz and Miller (1991), Corfield and Cartlidge (1992), Pak and Miller (1992) and the scenario discussed above. The compilation of benthic 6L80 data by Zachos et al. (1992) for the early Eocene suggest that values were very similar to those of high latitude planktonic foraminifera (Stott et al., 1990; Barrera and Huber, 1991). This reinforces their interpretation that, during this interval when WSDW production is suggested to have been at its most intense, a large component of the deeper waters in the Indian Ocean were still formed in association with the cooler surface waters surrounding Antarctica. 3.3. Palaeocene climate change and potential CO 2

flux The above discussions of isotopic change in the early Cenozoic have indicated that surface water productivity probably underwent substantial change during Palaeocene time. Productivity change has been implicated (Shackleton et al., 1983; Toggweiler and Sarmiento, 1985) in the control of atmospheric CO 2 during the last deglaciation. Shackleton et al. (1983) found that a 1%o change in A613C was approximately equivalent to a 100 ppm change in atmospheric CO 2 concentration. Increasing A6~3C accompanies decreasing atmospheric CO 2 and vice versa, via the mechanism previously described. The large changes in 613C found during the Palaeocene beg the question of whether Palaeocene climate changes (monitored using 3180 change) are related to changing atmospheric CO 2 concentrations. As Zachos et al. (1993) have noted, the oceanic mixed layer and the atmosphere equilibrate on timescales of 10 2 years i.e. much faster than the equilibration time of the surface and deep ocean. Hence any productivity driven changes in atmospheric CO 2 are effectively instantaneous on geological timescales and in the absence of negative feedback effects could hypothetically maintain atmospheric p C O z at new levels assuming that productivity is the only limit-

R.M.CorfieM~Earth-ScienceRet4ews37(1994)225-252

242

ing factor. If the calibration established for the Pleistocene can be applied to the Palaeogene then the increase in Palaeocene A~13Csurface_benthic of up to 3%o could in theory decrease atmospheric CO 2 concentration by up to 300 ppm. Shackleton et al. (1983) found that their estimates of pCO 2 change during the last deglaciation agreed well with those derived by Neftel et al. (1982) from direct measurements of Pleistocene palaeo-CO 2 from bubbles preserved in ice-cores. Corfield and Cartlidge (1992) compiled oxygen isotope data for the Palaeocene from the same samples as illustrated in Fig. 8. These data are shown in Fig. 11. A palaeotemperature scale for these 6lSO measurements is also shown. To correct for an ice-free planet they used an estimated 6 w of - 1.2%~ for planktonic species of the gen-

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era Morozovella, Acarinina and Subbotina as well as benthonic foraminifera and minimised potential artifacts in the data set by not correcting for hypothesised local area ~180 enrichment effects arising from evaporative processes. Because of the presumed lack of polar ice the interpretation below of 6180 change in the Palaeocene assumes that 8180 fluctuations solely reflect temperature change. Fig. 11 clearly shows that latitudinal seasurface temperature gradients were very shallow during the Palaeocene compared to the present day. Corfield and Cartlidges' (1992) estimate of 0.25°C/1 ° of latitude between Site 577 and the Leg 74 sites for the late Palaeocene is similar to the estimate by Shackleton and Boersma (1981) of 0.29°C/1 ° of latitude during the early Eocene. ~180 values in surface dwelling planktonics

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Fig. 11. 61So change in surface dwelling planktonic, thermocline dwelling planktonic and benthonic foraminifera during Palaeocene time at DSDP 577, a composite succession of the Leg 74 sites (DSDP Sites 525 and 527) and a composite of Leg 113 sites (ODP sites 689 and 690). Data from Shackleton et al. (1984), Stott et al. (1990) and Corfield and Cartlidge, (1992). Redrawn from Corfield and Cartlidge (1992). Also plotted is a palaeotemperature scale assuming a 6,,, of - 1.2.

243

R.M. Corfield~Earth-Science Reviews37 (1994) 225-252

the thermocline except that the 613C data for Morozovella and Subbotina also diverge (Fig. 8) supporting the suggestion of habitat separation. Contemporaneous data from the Leg 74 sites support this hypothesis. This habitat separation may be associated with the acquisition of photosymbionts as suggested by Norris et al. (pers. commun.) The potential acquisition of photosymbionts by the Palaeocene planktonic foraminifera may also affect isotopic ratios in an analogous way to Holocene species with photosymbionts. In the Holocene such species tend to exhibit higher 613C values than species without photosymbionts (Spero and Williams, 1988, 1989; McConnaughey, 1989; Spero et al., 1991). This enrichment in ~3C is proportional to light intensity so that surface dwellers tend to show the heaviest values. Hence

remain relatively invariant during the Palaeocene in the Leg 86 succession implying little variation in sea-surface temperatures. In the early Eocene 8180 values increase slightly suggesting a cooling in sea surface temperature from about 20°C to approximately 18.5°C. The record from the deeper dwelling planktonic foraminiferal genus Subbotina shows a greater degree of high-frequency variation which may be due to seasonal variations in their position in the water column as suggested by Stott et al. (1990) for high southern latitude Subbotina. The 8180 records of surface dwelling planktonics and Subbotina diverge between 64 and 65 Ma. This may be due to the colonisation of surface waters by keeled Morozovella as suggested by Corfield and Cartlidge (1992). This divergence could be explained by changes in the intensity of

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Fig. 12. Comparison of benthonic 61SO from the late Maastrichtian to the early Eocene. Leg 86 data from Corfield and Cartlidge (1992) and from Pak and Miller (1992), Leg 74 data from Corfield and Cartlidge (1992), Leg 113 data from Stott et al. (1990) and Corfield and Cartlidge (1992), ODP 738C data from Barrera and Huber (1991), ODP 752 data from Seto et al. (1991). The early and early late Palaeocene 180 enrichment, and the subsequent late Palaeocene and early Eocene 180 depletion have been highlighted by the solid trend line.

244

R.M. CorfieM~Earth-Science Ret,iews 37 (1994) 225-252

if this relationship holds true for Palaeogene species then at least part of the reason for the heavy values of 6~3C in surface dwelling planktonic foraminifera (Morozovella and Acarinina) may be because of this variety of "vital effect". It should however, be emphasised that the Palaeocene 613C maximum is not an artifact of this phenomenon because benthic species show a parallel but lesser amplitude enrichment to planktonic species during Palaeocene time. The 6~SO difference between surface dwellers and Subbotina species for the remainder of the Palaeocene and the early Eocene is greatest in DSDP Site 577 which, as expected, suggests the presence of well-developed thermocline in the low-latitude Pacific at this time. The benthonic 6~SO record at DSDP Site 577 implies a decline in deep water temperatures from about 65 Ma until 62 Ma. This was followed by a dramatic increase in deep water temperatures (of 6°C) between 62 Ma and 55 Ma from the early Late Palaeocene into the early Eocene. The benthonic 61SO record in the composite Leg 74 succession does not suggest early Palaeocene and early late Palaeocene cooling, although a temperature increase of about 3.5°C between 61 Ma and 55 Ma (i.e. synchronous with the 61SO decrease in DSDP 577) is observed. Due to the presence of a hiatus in the early part of the Late Palaeocene in the Leg 113 sites it is not possible to assess any early-late Palaeocene benthonic temperature change. However, a warming in deep water temperatures between 61.5 and 57.5 Ma can be clearly discerned. Unlike the lower latitude sites there was also a warming in surface (monitored using Acarinina) and near-surface (monitored using Subbotina) waters over the same interval. The same pattern of deep water 61SO increase and subsequent decrease is found in other sites, notably in the Indian Ocean (Zachos et al., 1993). Benthic 61SO data through the Palaeocene are summarised in Fig. 12. To assist the eye a trend line highlighting the early and early late Palaeocene 61SO increase and the subsequent late Palaeocene and early Eocene 6~SO decrease is drawn through the data. Corfield and Cartlidge (1992) suggested that the deep water cooling (in the early and the

early/late Palaeocene) and the deep water warming (in the late Palaeocene and the early Eocene) were linked to the changes in the carbon budget implied by the 613C record. In DSDP Site 577 the 6180 increase of c. 0.8%o corresponds to a potential cooling of deep waters by about 33.5°C between 65.5 Ma and 62 Ma, and is broadly contemporaneous with the increase in surface water 613C of about 2.8%o. After a thermal minimum at 61.9 Ma deep water 8180 decreased by about 1.5%o (implying a c. 6°C warming) contemporaneous with a surface water 6~3C decrease of about 3%o. Early and early/late Palaeocene sediment is missing from the Leg 113 succession so it is not possible to correlate 6~80 and 613C change over this interval. However, a 6180 decrease of c. 1.5%o between 61 Ma and 55.5 Ma (late Palaeocene to early Eocene interval) correlates with the decrease in surface water 6t3C of 3%o over the same interval. In the Leg 74 composite section benthic 6|~O changes through the entire Palaeocene interval are of smaller amplitude than those found in the DSDP Site 577 and composite Leg 113 successions. The Leg 74 composite succession does not show an increase in benthonic 61SO in the early and early/late Palaeocene interval (during the Palaeocene 6~3C increase) but does show a (albeit comparatively depressed) benthonic 6~SO decrease of c. 0.8%o in the late Palaeocene and early Eocene interval between 62 and 55 Ma contemporary with the 6 ~3C decline of c. 2.5%c in this succession. Palaeocene 61SO data from ODP 738C show a c. 1%o increase in the early and early/late Palaeocene followed by a c. 1%c decrease in the late Palaeocene which corresponds well with the timing and direction of the benthic 61SO trend from the other sites. Data from ODP Site 752 similarly show a 1.3%o increase followed by a 0.5%c decrease, again broadly in step with the other sites. The anomalously light values from ODP 752 may reflect either fractionation out of isotopic equilibrium or diagenetic resetting of the 8180 signal. Note that the latter process need not necessarily alter the trajectory of 6~SO change. Furthermore, Katz and Miller (1991) generated data from the Palaeocene of ODP Hole 702B which shows an increase in 6180benthonic in the subantarctic South

R.M. Corfield / Earth -Science Reciews 37 (1994) 225-252

Atlantic which was contemporaneous with the increase in ~13Cbenthonic in the early and early late Palaeocene. Following this, their data clearly show the marked 3180 decline in the late Palaeocene to early Eocene interval. Corfield and Cartlidge (1992) suggested that the apparent broad correlation between ~]3C and 6180 change in the Palaeocene is the result of productivity driven changes in atmospheric CO 2. They hypothesised that the implied increase in photosynthesising biomass in the e a r l y - e a r l y / l a t e Palaeocene could have had the effect of decreasing the p C O 2 of surface waters. This in turn could have lead to a draw-down of atmospheric

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246

R.M. CorfieM~Earth-Science Reviews"37 (1994) 225-252

least four independent data sets. The correlation is particularly strong in the Leg 113 data (where data only exists in significant quantities for the late Palaeocene to early Eocene interval) where there is also substantial covariance of surface dwelling, thermocline dwelling and benthic foraminiferal ~5x3C and /5180 (see Fig. 11). This latter observation is consistent with the suggestion that the southern high latitudes were the major locus of deep water production through much of the Palaeocene and early Eocene. Hence it may be that the effect of the implied atmospheric CO 2 change was felt most strongly at the locus of bottom water production affecting all levels of the water column and manifesting itself by correlation of benthonic 613C and 6~80 in lower latitude sites. Such a hypothesis is consistent with results of modelling experiments (e.g. Barron and Peterson, 1991) which suggest that CO 2 change has the most impact at high latitudes. To rigorously test this idea it is critical to obtain high-resolution, parallel planktonic and benthonic records for a high latitude site through the early and early/late Palaeocene interval i.e. when ~3C was increasing and the Corfield and Cartlidge (1992) hypothesis predicts that cooling of all levels of the water column should have occurred. It should be noted that the correlation is good between benthonic 613C and benthonic ~180 and much poorer between surface water 6~3C and benthonic ~Iso. This may imply that if there was a relationship between temperature change and CO~ flux, then it was the burial of organic carbon (monitored using benthonic 613C change) rather than simply productivity change (monitored using surface 313C change)which was the major control on the system. It has already been noted that benthic 6~3C change apparently lagged surface water ~ J3C change during the Palaeocene carbon isotope maximum. This feature is clearly worthy of further study. The amplitude of the benthonic temperature increase from the Palaeocene into the Eocene was generally twice as great as the benthonic temperature decrease in the e a r l y / l a t e Palaeocene. As suggested by Corfield and Cartlidge (1992) this may be partly related to the

greater amplitude of the ~13Cbenthonic decrease from the Palaeocene into Eocene compared to the ~13Cbenthonic increase in the early to late Palaeocene. Furthermore, this ~13C decline may be due to the addition, to the atmosphere and surface ocean, of CO2 depleted in 12C from mantle a n d / o r terrestrial sources as discussed above (Owen and Rea, 1985; Rea et al., 1990). An increase in atmospheric CO 2 from these reservoirs together with outgassing of marine CO 2 due to reduced surface water productivity could account for the pronounced apparent temperature rise from the late Palaeocene into the early Eocene. The hypothesis of productivity or carbon-burial controlled CO2 flux and consequent "icehouse/ greenhouse" climate change in the Palaeocene and early Eocene is similar to models suggested for other times of very positive ~13Ccarbonate values. Vincent and Berger (1985) have suggested that during the middle Miocene the deposition of the circum-Pacific Monterey Shale led to a draw down of atmospheric CO 2 and thereby conditioned the planet for glacially dominated climates. Arthur et al. (1988) have suggested that the deposition of the black shales at the Cenomanian/Turonian boundary led to CO 2 draw-down and possible climatic cooling. Finally it should be noted that another explanation has been proposed for the cause of the Palaeocene carbon isotope maximum that is completely unrelated to marine productivity change. Oberh~insli and Perch-Nielsen (1990) have suggested that the early to late Palaeocene cooling led to an extension of high latitude terrestrial floras into the lower latitudes. They speculate that this resulted in increased burial rates of organic carbon in the terrestrial realm. A problem with this approach, which they acknowledge, is that the current size of the terrestrial carbon reservoir is over forty times smaller than the combined marine sedimentary and oceanic carbon reservoirs. One or both of these reservoirs have been implicated by the majority of authorities (Shackleton et al., 1985; Shackleton, 1986; Miller et al., 1987; Corfield and Cartlidge, 1992 and in this contribution) who have addressed the problem of the Palaeocene carbon isotope maxi-

R.M. Corfield / Earth-Science Reuiews 37 (1994) 225-252

mum. The involvement of reservoirs of this very large size is in fact required to explain the magnitude of the observed 313C excursion. To overcome this objection, Oberh~insli and Perch-Nielsen suggest that the terrestrial carbon reservoir was very much larger during Palaeocene time. At present there is insufficient palaeobotanical evidence to verify this suggestion. In contrast the assertion of an increase in marine productivity is supported by accumulation rate data on vertebrate remains (fish teeth) as well as bulk carbonate (Shackleton et al., 1984) in addition to accumulation rates of planktonic foraminifera (Corfield and Norris, unpubl, data).

4. The late Palaeocene thermal maximum ("Palaeocene / Eocene boundary event") Superimposed on the long-term ~ 3 C decline from the late Palaeocene into the early Eocene is a substantial transient excursion in planktonic and benthonic 613C and 6180 originally discovered in ODP 690B (Kennett and Stott, 1990; Kennett and Stott, 1991) and now referred to as the Late Palaeocene Thermal Maximum (LPTM) (Zachos et al., 1993). In ODP Site 690 planktonic 6~80 decreased markedly and rapidly implying that surface water temperatures increased by 5 6°C, benthic 8~So also increased suggesting a deep water warming of about 4°C (Zachos et al., 1993). This warming apparently occurred over an interval as short as 10 kyr but was immediately followed by a cooling lasting 100 kyr (Zachos et al., 1993). Benthic 613C values decreased synchronously with the 31So decrease by over 3%o, planktonic values similarly decreased but by an even larger amount which effectively reduced the vertical carbon isotope gradient to zero. ,~13C values thereafter recovered but to values 1%o lower than before the event reflecting the fact that this short term excursion was superimposed on the long term decline in ~13C into the early Eocene. This Late Palaeocene Thermal Maximum (Zachos et al., 1993) close to the P a l a e o c e n e / E o c e n e boundary is also associated with the most profound episode of benthic foraminiferal extinction in the last 100 million

247

years. It has been suggested (Thomas, 1990; 1991) that this event was associated with a significant reorganisation of deep water circulation. This excursion has now also been noted in many other deep sea and terrestrial sections (Barrera and Huber, 1991; Thomas, 1991; Corfield and Cartlidge, 1992; Koch et al., 1992; Pak and Miller, 1992; Stott, 1992) although with varying amplitudes and durations which may be artifacts of the timescales and sedimentation history of the various sites. Thomas (1991) has suggested that the 6~80 and 613C minimum at the P a l a e o c e n e / E o c e n e boundary may have been caused by exhalation of mantle CO 2 from North Atlantic volcanism, such a hypothesis is compatible with that of Owen and Rea (1985) who suggested a link between hydrothermal activity, an increase in atmospheric CO 2 and probable Eocene " G r e e n h o u s e " climate. However, Stott (1992) has inferred on the basis of 613C measurements of the organic component of the foraminiferal calcite matrix that the apparent warming near the P a l a e o c e n e / E o c e n e boundary was not associated with an increase in CO 2 but a decrease of up to 500 ppm. He further suggests that this CO 2 change near the P / E boundary is a consequence rather than a cause of the LPTM.

5. Closing remarks It has been generally assumed that the broad Palaeocene carbon isotope maximum is in some way a unique perturbation to the carbon budget of the Cenozoic. This of course it is, because the division between the Cenozoic and the Mesozoic is not based on palaeoceanographic criteria but is arbitrarily placed at a interval of substantial mass extinction (the K / T boundary). Comparisons of Palaeocene and Cretaceous fil3c values (e.g. data in Corfield et al., 1991) suggest that the Palaeocene may have been very similar in terms of carbon budget to the Cretaceous. It is therefore worth considering the possibility that there is nothing unusual in the amplitude of Palaeocene ~ 3 C and the rapidity with which ~13C values increased in the aftermath of

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R.M. Corfield ~Earth-Science Reviews 37 (1994) 225-252

the Cretaceous/Tertiary boundary, after all, conditions were merely recovering to those that characterised at least the preceding 30 million years of Earth history. If, as implied by this reasoning, the Palaeocene is merely an outpost of the Cretaceous, then it was the broad ~3C minimum in the early Eocene (not, incidentally, to be confused with the "LPTM") that was significant in terms of change to the overall carbon budget of the Earths oceans and atmosphere, and the K / T boundary was merely an externally imposed anomaly.

Acknowledgements Thanks always to Julie Cartlidge for her expert operation of the mass spectrometers in the Stable Isotope Laboratory at Oxford and for compiling some of the data illustrated in this contribution. I am grateful to Birger Schmitz, Steve D'Hondt, Dick Norris and Larry Jones for careful reviews of the completed manuscript. My thanks to Dick Norris, Birger Schmitz, Jim Zachos, Marie-Pierre Aubry and Bill Berggren for many stimulating discussions on Palaeocene isotopic change. Thanks also to Stella Charisi for providing me with a copy of an obscure reference at short notice and Leon Clarke who commented on an early version of this manuscript. The stable isotope laboratory in Oxford is supported by University funding.

References AIvarez, L.W., Alvarez, W., Asaro, F. and Michel, H.V., 1980. Extraterrestrial cause for the Cretaceous-Tertiary extinction. Science, 208(4448): 1095-1107. Arthur, M.A., Dean, W.E. and Pratt, L.M., 1988. Geochemical and climatic effects of increased marine organic carbon burial at the Cenomanian/Turonian boundary. Nature, 335: 714-717. Barrera, E,, Huber, B.T., Savin, S.M. and Webb. P.-N., 1987. Antarctic marine temperatures: Late Campanian through early Paleocene. Paleoceanography, 2: 21-47. Barrera, E. and Huber, B.T., 1991. Paleogene and early Neogene oceanography of the southern Indian Ocean: Leg 119 foraminifer stable isotope results. In: J. Barron, B.

Larsen et al. (Editors), Proc. ODP, Sci. Results. College Station, TX (Ocean Drilling Program), 119: 693-717. Barrera, E. and Keller, G., 1990. Stable isotope evidence for gradual environmental changes and species survivorship across the Cretaceous/Tertiary boundary. Paleoceanography, 5: 867-890. Barton, E. and Peterson, W.H., 1991. The Cenozoic ocean circulation based on ocean General Circulation Model results. Palaeogeogr., Palaeoclimatol., Palaeoecol., 83: 128. Berger, W.H. and Vincent, E., 1986. Deep-sea carbonates: Reading the carbon isotope signal. Geol. Rundsch., 75(1): 249-269. Berggren, W.A., Kent, D,V. and F[ynn, J.J., 1985, Paleogene geochronology and chronostratigraphy. In: N.D. Snelling (Editor), The Chronology of the Geological Record. Geol. Soc. Mere., 10: 141-195. Boersma, A. and Shackleton, N.J., 1981. Oxygen and carbon isotope variations and planktonic foraminifera depth habitats, late Cretaceous to Paleocene, Central Pacific, Deep Sea Drilling Project Sites 463 and 465. In: J. Thiede, T.L. Vallier et al., Initial Reports of the Deep Sea Drlling Project, 62. U.S. Government Printing Office, Washington, pp. 513-526. Bralower, T.J. and Thierstein, H.R., 1984. Low productivity and slow deep-water circulation in mid-Cretaceous oceans. Geology, 12: 614-618. Brass, G.W., Southam J.R. and Peterson, W.H., 1982. Warm saline bottom water in the ancient ocean. Nature, 296: 620-623. Brennecke, J.C. and Anderson, T.F., 1977. Carbon-isotope variation in pelagic sediments (abs.). EOS Trans. Am. Geophys. Union, 58: 415. Broecker, W., 1982. Ocean chemistry during glacial time. Geochim. Cosmochim. Acta, 46: 1689-1705. Broecker, W.S., 1989. The salinity contrast between the Atlantic and Pacific oceans during glacial time. Paleoceanography, 4: 207-212. Broecker, W.S. and Peng, T-H., 1982. Tracers in the Sea. Lamont- Doherty Geological Observatory, Palisades, N.Y., 690 pp. Chamberlin, T.C., 1906. On a possible reversal of deep sea circulation and its influence on geologic climates. J. Geol., 14: 363-373. Chappell, J. and Shackleton, N.J., 1986. Oxygen isotopes and sea level. Nature, 324(6093): 137-140. Corfield, R.M., Cartlidge, J.E., Premoli-Silva I. and Housley, R.A., 1991. Oxygen and carbon isotope stratigraphy of the Palaeogene and Cretaceous limestones in the Bottaccione Gorge and the Contessa Highway sections, Umbria, Italy. Terra Nova, 3(4): 414-422. Corfield, R.M. and Cartlidge, J.E, 1992. Oceanographic and climatic implications of the Palaeocene carbon isotope maximum. Terra Nova, 4: 443-455. Corfield, R.M. and Cartlidge, J.E., 1993. Whole-rock oxygen and carbon isotope stratigraphy of the Paleogene and Cretaceous/Tertiary boundary in ODP Hole 807C. In: L.

R.M. Corfield ~Earth-Science Reviews 37 (1994) 225-252 Kroenke, W. Berger, T. Janacek et al. (Editors), Proc. ODP, Sci. Results. College Station, TX (Ocean Drilling Program), 130: 259-268. Courtillot, V., Vandamme, D., Besse, J., Jaeger, J.J. and Javoy, J., 1990. Deccan volcanism at the Cretaceous/ Tertiary boundary: Data and inferences. In: V.L. Sharpton and P.D. Ward (Editors), Global Catastrophes in Earth History. An interdisciplinary conference on impacts, volcanism and mass mortality. Geol. Soc. Am. Spec. Pap., 274: 401-410. Craig, H. and Gordon, L.I., 1965. Deuterium and oxygen-18 variations in the ocean and the marine atmosphere. In: E. Tongiorgi (Editor), Stable Isotopes in Oceanographic Studies and Palaeotemperatures. Consiglio Nazionale delle Richerche, Lab. di Geologia Nucleate, Pisa. Cresta, S., Monechi, S. and Parisi, G., 1989. Mesozoic-Cenozoic stratigraphy in the Umbria-March Area. Memoire descrittive della Carta Geologica D'Italia, XXXIX. lstituto Poligrafico E Zecca Dello Stato, 185 pp. Curry, W.B. and Lohmann, G.P., 1983. Reduced advection into Atlantic Ocean deep eastern basins during last glaciation maximum. Nature, 306: 577-580. Curry, W.B. and Lohmann, G.P., 1982. Carbon isotopic changes in benthic foraminifera from the western South Atlantic: Reconstruction of glacial abyssal circulation patterns. Quat. Res., 18: 218-235. D'Hondt, S. and Lindinger, M., 1988. An extended Cretaceous/Tertiary boundary stable isotope record: Implications for paleoclimate and the nature of the K / T boundary event. Global catastrophes in Earth history. Luna. Planet. Inst; pp. 40, 41. D'Hondt, S. and Zachos, J.C., 1993. On stable isotopic variation and earliest Paleocene planktonic foraminifera. Paleoceanography, 8: 4: 527-547. Duplessy, J-C., Shackleton, N.J., Matthews, R.K., Prell, W., Ruddiman, W.F., Caralp, M. and Hendy, C.H., 1984. C 13 record of benthic foraminifera in the last interglacial ocean: implication for the carbon cycle and the global deep water circulation. Quat. Res., 21: 225-243. Emery, W.J. and Meincke, J., 1986. Global water masses: summary and review. Oceanol. Acta, 9(4): 383-391. Emiliani, C., 1954. Depth habitats of some species of pelagic foraminifera as indicated by oxygen isotope ratios. Am. J. Sci., 252: 149-158. Emiliani, C., 1955. Pleistocene temperatures. J. Geol., 63(6): 538-575. Epstein, S., Buchsbaum, R., Lowenstam, H.A. and Urey, H.C., 1951. Carbonate-water isotopic temperature scale. Geol. Soc. Am. Bull., 62: 417-426. Epstein, S., Buchsbaum, R., Lowenstam, H.A. and Urey, H.C., 1953. Revised carbonate-water isotopic temperature scale. Geol. Soc. Am. Bull., 64: 1315-1326. Erez, J. and Honjo, S., 1981. Comparison of isotopic composition of planktonic foraminifera in plankton tows, sediment traps and sediments. Palaeogeogr., Palaeoclimatol., Palaeoecol., 33: 129-156. Erez, J. and Luz. B., 1983. Experimental paleotemperature

249

equation for planktonic foraminifera. Geochim. Cosmochim. Acta., 47: 1025-1031. Flores, R.M., Warwick, P.D. and Moore, T.A., 1989. A review of depositional aspects and a guide to Palaeocene coalbearing sequences, Powder River basin. In: Coal and Hydrocarbon Resources of North America, 1 - - Tertiary and Cretaceous Coals in the Rocky Mountain region. USGS: 1-10. Hsfi, K.J., 1986. Cretaceous/Tertiary boundary event. In: K.J. Hsfi (Editor), Mesozoic and Cenozoic Oceans. Am. Geophys. Union, Geodymanics Series, 15: 75-84. Hsii, K.J., He, Q., McKenzie, J.A., Weissert, H., Perch-Nielsen, K., Oberh~insli, H., et al., 1982. Mass mortality and its environmental and evolutionary consequences. Science, 216(4543): 249-256. Hsii, K.J. and McKenzie, J.A., 1985. A "strangelove"ocean in the earliest Tertiary, In: E.T. Sundquist and W.S. Broecker (Editors), The Carbon Cycle and Atmospheric CO,: Natural variations Archaen to Present. Geophys. Monogr. Am. Geophys. Union, 32: 487-492. Hsii, K.J. and McKenzie, J.A., 1990. Carbon-isotope anomalies at era boundaries; Global catastrophes and their ultimate cause. In: V.L. Sharpton and P.D. Ward (Editors), Global Catastrophes in Earth History. An interdisciplinary conference on impacts, volcanism and mass mortality. Geol. Soc. Am. Spec. Pap., 274: 61-70. Hudson, J.D. and Anderson, T.F., 1989. Ocean temperatures and isotopic compositions through time. Trans. R. Soc. Edinburgh, (Earth Sciences) 80: 183-192. Imbrie, J., Hays, J.D., Martinson, D.G., McIntyre, A., Mix, A., Morley, J.J., Pisias, N.G., Prell, W. and Shackleton, N.J., 1984. The orbital theory of Pleistocene climate: support from a revised chronology of the marine 8~SO record. In: A. Berger, J. Imbrie, J.D. Hays, G. Kukla and B. Saltzman (Editors), Milankovitch and Climate. Reidel, Hingham, Mass., pp. 269-305. Ivany, L.C. and Salawitch, R.J., 1993. Carbon isotopic evidence for biomass burning at the K / T boundary. Geology, 21: 487-490. Janacek, T.R. and Rea, D.K., 1983. Eolian deposition in the northeast Pacific Ocean: Cenozoic history of atmospheric circulation. Geol. Soc. Am. Bull., 94: 730-738. Kaiho, K., 1991. Global changes of Paleogene aerobic/ anaerobic benthic foraminifera and deep sea circulation. Palaeogeogr., Palaeoclimatol., Palaeoecol., 83: 65-86. Katz, M.E. and Miller, K.G., 1991. Early Palaeogene benthic foraminiferal assemblages and stable isotopes in the Southern ocean. In: P.F. Ciesielski, Y. Kristoffersen et al. (Editors), Proc. ODP, Sci. Results. College Station, TX (Ocean Drilling Program), 114: 481-512. Keigwin, L.D. Jr., 1980. Palaeoceanographic change in the Pacific at the Eocene-Oligocene boundary. Nature, 287: 722-725. Keller, G. and Lindinger, M., 1989. Stable isotope, TOC, and CaCO 3 record across the Cretaceous/Tertiary boundary at E1 Kef, Tunisia. Palaeogeogr., Palaeoclimatol., Palaeoecol., 73: 243-265.

250

R.M. Corfield ~Earth-Science Rec'iews37 (1994) 225-252

Keller, G. and Macleod, N., 1993. Comment on "Carbon isotopic evidence for biomass burning at the K / T boundary". Geology, 21: 1149-1151. Kennett, J.P. and L.D. Stott, 1990. Proteus and Proto-Oceanus: Ancestral paleogene oceans as revealed from Antarctic stable isotope results. In: P.F. Barker, J.P. Kennett et al. (Editors), Proc. ODP, Sci. Results. College Station, TX (Ocean Drilling Program), 113: 829-848. Kennett, J.P. and Stott, L.D., 1991. Abrupt deep sea warming, paleoceanographic changes and benthic extinctions at the end of the Palaeocene. Nature, 353: 225-229. Koch, P.L., Zachos, J. and Gingerich, P., 1992. Correlation between isotope records in marine and continental carbon reservoirs near the Palaeocene/Eocene boundary. Nature, 358: 319-322. Kroopnick, P., 1985. The distribution of 13C of TCO 2 in the world oceans, Deep-Sea Res., 32: 57-84. Kump, L., 1991. Interpreting carbon-isotope excursions: Strangelove oceans. Geology, 19: 299-302. McConnaughey, T., 1989. ISC and 180 isotopic disequilibrium in biological carbonates, II. In vitro simulation of kinetic isotope effects. Geochim. Cosmochim. Acta, 53: 163-171. McLean, D.M., 1985. Mantle degassing-induced dead ocean in the Cretaceous/Tertiary transition. In: E.T. Sundquist and W.S. Broecker (Editors), The Carbon Cycle and Atmospheric C02: Natural Variations Archaen to Present. Geophys. Monogr. Ser., AGU, Washington D.C., 32: 493503. Miller, K.G. and Fairbanks, R.G., 1985. Oligocene to Miocene benthic foraminifera and abyssal circulation changes. In: E.T. Sundquist and W.S. Broecker (Editors), The Carbon Cycle and Atmospheric CO2: Natural Variations Arcbaen to Present. Geophys. Monogr. Ser., AGU, Washington D.C., 32: 469-486. Miller, K.G., Janacek, T.R., Katz, M.E. and Keil, D.J., 1987. Abyssal circulation and benthic foraminiferal changes near the Palaeocene/Eocene boundary. Paleoceanography, 2(6): 741-761. Neftek A., Oeschger, H., Schwander, J., Stauffer, B. and Zumbrunn, R. 1982. Ice core sample measurements give atmospheric CO 2 contents during the past 40,000 years. Nature, 295: 220-223. Oberh/insli. H and Toumarkine, M., 1985. The Paleogene oxygen and carbon isotope history of DSDP Sites 522, 523 and 524 from the central south Atlantic Ocean. In: K.J. Hsii and H. Weissert (Editors), South Atlantic Paleoceanography. Cambridge University Press, pp. 125-147. Oberh~nsli, H and Perch-Nielsen, K. von Salis, 1990. The Paleocene 3C-event Was it due to changes in the storage rate of terrestrial biomass? Veroff. Ubersee-Mus., A10: S.99-112, Officer, C.B. and Drake, C.L., 1985. Terminal Cretaceous environmental events. Science, 227: 1161-1167. Owen, R.M. and Rea, D.K., 1985. Sea-floor hydrothermal activity links climate to tectonics: the Eocene carbon dioxide greenhouse. Science, 227: 166-169. Pak, D.K. and Miller, K.G., 1992. Paleocene to Eocene ben-

thic foraminiferal isotopes and assemblages: Implications for deep-water circulation. Paleoceanography, 7: 405-422. Perch-Nielsen, K., McKenzie, J. and He, Q., 1982. Biostratigraphy and isotope stratigraphy and the "catastrophic" extinction of calcareous nannoplankton at the Cretaceous/Tertiary boundary. GSA Spec. Pap., 190: 353-371. Prell, W., lmbrie, J., Martinson, D.G., Morley, J., Pisias, N.G., Shackleton, N.J. and Streeter, H., 1986. Graphic correlation of oxygen isotope stratigraphy: Application to the late Quaternary. Paleoceanography, 1: 137-162. Rea, D.K., Leinen, M. and Janacek, T.R., 1985. Geologic approach to the long-term history of atmospheric circulation. Science, 227: 721-725. Rea, D.K., Zachos, J.C., Owen, R.M. and Gingerich, P.D., 1990. Global change at the Palaeocene-Eocene boundary: climatic and evolutionary consequences of tectonic events. Palaeogeogr., Palaeoclimatol., Palaeoecol., 79: 117-128. Renard, M., 1986. Pelagic carbonate chemostratigraphy (St, Mg, 180, 13C). Mar. Micropalaeontol., 10: 117-164. SchmitL B., Keller, G. and Stenvall, O., 1992. Stable isotope and foraminiferal changes across the Cretaceous-Tertiary boundary at Stevns Klint, Denmark: Arguments for longterm oceanic instability before and after bolide-impact event. Palaeogeogr., Palaeoclimatol., Palaeoecol., 96: 233260. Scholle, P.A. and Arthur, M.A., 1980. Carbon isotope fluctuations in Cretaceous pelagic limestones: Potential stratigraphic and petroleum exploration tool. Am. Assoc. Petrol. Geol. Bull., 64: 67-87. Seto, K., Nomura, R. and Niitsuma, N., 1991. Data Report: Oxygen and carbon isotope records of the upper Maastrichtian to lower Eocene benthic foraminifers at Site 752 in the eastern Indian Ocean. In: J. Weissel, J. Peirce et al. (Editors), Proc. ODP, Sci. Results. College Station, TX (Ocean Drilling Program), 121: 885-890. Shackleton, N.J., 1967. Oxygen isotope analyses and Pleistocene temperatures re-assessed. Nature, 215: 15-17. Shackleton, N.J., 1982. The deep-sea sediment record of climate variability. Prog. Oceanogr., 11: 199-218. Shackleton, N.J., 1985. Oceanic carbon isotope constraints on oxygen and carbon dioxide in the Cenozoic atmosphere. In: E.T. Sundquist and W.S. Broecker (Editors), The Carbon Cycle and Atmospheric CO2: Natural Variations Archaen to Present. Geophys. Monogr. Ser.. AGU, Washington D.C., 32: 455-468. Shackleton, N.J. and members of the shipboard scientific party, 1984. Accumulation rates in Leg 74 sediments. In: T.C. Moore Jr., P.D. Rabinowitz et al. (Editors), Initial Reports Deep Sea Drilling Project. U.S. Government Printing Office, Washington D.C., 74: 621-644. Shackleton, N.J., 1986. Palaeogene stable isotope events. Palaeogeogr., Palaeoclimatol., Palaeoecol., 57: 91-102. Shackleton, N.J., 1987. The carbon isotope record of the Cenozoic: history of organic carbon burial and of oxygen in the ocean and atmosphere. In: J. Brooks and A.J. Fleet (Editors), Marine Petroleum Source Rocks. Geol. Soc. London Spec. Publ., 26: 423-434.

R.M. Corfield / Earth-Science Ret,iews 37 (1994) 225-252 Shackleton, N.J. and Kennett, J.P., 1975. Paleotemperature history of the Cenozoic and the initiation of Antarctic glaciation: Oxygen and carbon isotope analyses in DSDP Sites 277, 279, and 281. In: J.P. Kennett, R.E. Houtz et al. (Editors), Initial Reports. Deep Sea Drilling Project. U.S. Government Printing Office, Washington, 29: 743-755. Shackleton, N.J. and Boersma, A., 1981. The climate of the Eocene Ocean. J. Geol. Soc. London, 138: 153-157. Shackleton, N.J. and Hall, M.A., 1984. Carbon isotope data from Leg 74 sediments, In: T.C. Moore Jr., P.D. Rabinowitz et al. (Editors), Initial Reports Deep Sea Drilling Project. U.S. Government Printing Office, Washington D.C., 74: 613-619. Shackleton, N.J., Hall, M.A., Line, J. and Shuxi, C., 1983. Carbon isotope data in core V19-30 confirm reduced carbon dioxide in the ice age atmosphere. Nature, 306(5941): 319-322. Shackleton, N.J., Imbrie, J. and Hall, M.A., 1984. Oxygen and carbon isotope record of East Pacific core V19-30: Implications for the formation of deep water in the late Pleistocene North Atlantic. Earth Planet. Sci. Lett., 65: 233244. Shackleton, N.J., Hall, M.A. and Bleil, U., 1985a. Carbon isotope stratigraphy, Site 577. Initial Reports Deep Sea Drilling Project, 86: 5//3-512. Shackleton, N.J., Corfield, R.M. and Hall, M.A., 1985b. Stable isotope data and the ontogeny of Palaeocene planktonic foraminifera. J. Foraminiferal Res., 15: 321-336. Smit, J., 1990. Meteorite impact, extinctions and the Cretaceous/Tertiary boundary. Geol. Mijnb., 69: 187-204. Spero, H.J. and Williams, D.F., 1988. Extracting environmental from planktonic foraminiferal delta ~3C data. Nature, 335: 717-719. Spero, H.J. and Williams, D.F., 1989. Opening the carbon isotope "vital effect" black box, 1. Seasonal temperatures in the euphotic zone. Paleoceanography, 4: 593-601. Spero, H.J., Lerche, I. and Williams, D.F., 1991. Opening the carbon isotope "vital effect" black box, 2. Quantitative model for interpreting foraminiferal carbon isotope data. Paleoceanography, 6: 639-655. Spicer, R.A. and Corfield, R.M., 1992. A review of terrestrial and marine climates in the Cretaceous with implications for modelling the "Greenhouse" Earth. Geol. Mag., 129: 169-180. Stott, LD. and Kennett, J.P., 1989. New constraints in early Tertiary palaeoproductivity from carbon isotopes in foraminifera. Nature, 342: 6249: 526-529. Stott, L.D. and Kennett, J.P., 1990. The paleoceanographic and paleoclimatic signature of the Cretaceous/Paleogene boundary in the Antarctic: stable isotopic results from ODP Leg 113. In: P.F. Barker, J.P. Kennett et al. (Editors), Proc. ODP, Sci. Results. College Station, TX (Ocean Drilling Program), 113: 829-848. Stott, L.D., Kennett, J.P. Shackleton, N.J. and Corfield, R.M., 1990. The evolution of Antarctic surface waters during the Paleogene: Inferences from the stable isotopic composition of planktonic foraminifers. In: P.F. Barker, J.P. Ken-

251

nett et al. (Editors), Proc. ODP, Sci. Results. College Station, TX (Ocean Drilling Program), 113: 849-864. Stott, L.D., 1992. Higher temperatures and lower oceanic pCO2: a climate enigma at the end of the Paleocene epoch. Paleoceanography, 7: 395-404. Thierstein, H.R., 1981. Late Cretaceous nannoplankton and the change at the Cretaceous/Tertiary boundary. Soc. Econ. Paleontol. Mineral. Spec. Publ., 32: 355-394. Thomas, E., 1990. Late Cretaceous through Neogene deep-sea benthic foraminifers (Maud Rise, Weddell Sea, Antarctica). In: P.F. Barker, J.P. Kennett et al. (Editors), Proc. ODP, Sci. Results. College Station, TX (Ocean Drilling Program), 113: 571-594. Thomas, E., 1991. The latest Palaeocene mass extinction of deep sea benthic foraminifera: Result of global climate change, Geol. Soc. Am., Abstr. Programs, 23: A141. Thomas, E., 1992. Middle Eocene-Late Oligocene bathyal benthic foraminifera (Weddell Sea): faunal changes and implications for ocean circulation. In: D.A. Prothero and W.A. Berggren (Editors), Eocene-Oligocene Climatic and Biotic Evolution. Princeton University Press, Princeton, pp. 245-271. Thunell, R.C., Williams, D.F. and Howell, M., 1987. AtlanticMediterranean water exchange during the late Neogene. Paleoceanography, 2: 661-678. Toggweiler, J.R. and Sarmiento, J.L., 1985. Glacial to interglacial changes in atmospheric carbon dioxide: The critical role of ocean surface water at high latitudes. In: E.T. Sundquist and W.S. Broecker (Editors), The Carbon Cycle and Atmospheric CO2: Natural Variations Archaen to Present. Am. Geophys. Union Geophys. Monogr., 32: 163-184. Tucholke, B.E. and Vogt, P.R., 1979. Western North Atlantic: sedimentary evolution and aspects of tectonic history. Initial Reports Deep Sea Drilling Project, 43: 791-825. Urey, H.C., 1947. The thermodynamic properties of isotopic substances J. Chem. Soc., 562-581. Vincent, E. and Berger, W.H., 1985. Carbon dioxide and polar cooling in the Miocene: The Monterey Hypothesis. In: E.T. Sundquist and W.S. Broecker (Editors), The Carbon Cycle and Atmospheric CO2: Natural Variations Archaen to Present. Am. Geophys. Union Geophys. Monogr., 32: 455-468. Williams, D.F., Healy-Williams, N., Thunell, R.C., Leventer, A., 1983. Detailed stable isotope and carbonate records from the upper Maestrichtian-Lower Paleocene section of Hole 516F (Leg 72) including the Cretaceous/Tertiary boundary. In: P.F. Barker, R.L. Carlson, D.A. Johnson et al. (Editors), Initial Reports Deep Sea Drilling Project. U.S. Government Printing Office, Washington, 72: 921929. Woodruff, F. and Savin S.M., 1989. Miocene deepwater oceanography. Paleoceanography, 4(1): 87-140. Wright, J.D., Miller, K.G. and Fairbanks, R.G., 1991. Evolution of modern deepwater circulation: evidence from the late Miocene southern ocean. Paleoceanography, 6(2): 275-290.

252

R.M. Corfield ~Earth-Science ReL'iews37 (1994) 225-252

Wright, J.D., Miller, K.G. and Fairbanks, R.G., 1992. Early and middle Miocene stable isotopes: implications for deepwater circulation and climate. Paleoceanography, 7(3): 357-389. Wyrtki, K., 1973. Physical Oceanography of the Indian Ocean. In: B. Zeitzschel (Editor), The Biology of the Indian Ocean. Springer, Berlin, pp. 18-36. Zachos, J.C., Arthur, M.A., Thunell, R.C., Williams, D,F. and E.J. Tappa, 1985. Stable isotope and trace element geochemistry of carbonate sediments across the Cretaceous/ Tertiary boundary at Deep Sea Drilling Project Hole 577, Leg 86. In: G.R. Heath, L.H. Burckle et al. (Editors), Initial Reports Deep Sea Drilling Project, U.S. Government Printing Office, Washington, 86: 513-532. Zachos, J.C. and Arthur, M.A., 1986. Paleoceanography of the Cretaceous/Tertiary boundary event: inferences from stable isotopic and other data. Paleoceanography, 1: 5-26. Zachos, J.C., Arthur, M.A. and Dean, W.E., 1989a. Geochemical evidence for suppression of pelagic marine productivity at the Cretaceous/Tertiary boundary. Nature, 337: 61-64. Zachos, J.C., Arthur, M.A. and Dean, W.E., 1989b. Geochemical and paleoenvironmental variations across the Cretaceous/Tertiary boundary at Braggs, Alabama. Palaeogeogr., Palaeoclimatol., Palaeoecol., 69: 245-266. Zachos, J., Gingerich, P.D. and Koch, P.L., 1991. Marine and terrestrial carbon isotope records across the Paleocene/ Eocene boundary: Application to reconstructing paleoclimates. Geol. Soc. Am., Abstr. Programs, 23: A142. Zachos, J.C., Rea, D.K., Seto, K., Nomura, R. and Niitsuma, N., 1992. Paleogene and early Neogene deep water paleoceanography of the Indian Ocean as determined from benthic foraminifer stable carbon and oxygen isotope records. In: R.A. Duncan, D.K. Rea, R.B. Kidd, U. von Rad and J.K. Weissel (Editors), The Indian Ocean: A synthesis of results from the Ocean Drilling Program. Am. Geophys. Union Geophys. Monogr., 70: 351-385.

Zachos, J,C., Lohmann, K.C., Walker, J.C.G. and Wise, S.W., 1993. Abrupt climate change and transient climates during the Paleogene: A marine perspective. J. Geol., 101: 191213. Zachos, J.C., Stott, L.D. and Lohmann, K.C., 1994. Evolution of early Cenozoic marine temperatures. Paleoceanography, 9(2): 353-387. Richard Corfield was born in London in 1962. He graduated with honours from the University of Bristol in 1983 with a degree in Zoology and received his PhD from the University of Cambridge in 1987 where he researched the evolution and stable isotope systematics of the Palaeogene planktonic foraminifera. After he completed his graduate studies Dr. Corfield was awarded a Research Fellowship from the Natural Environment Research Council (a position he held concurrently with a Junior Research Fellowship at Jesus College, Oxford). He established and now runs the Stable Isotope Laboratory in the Department of Earth Sciences of the University of Oxford, where he has been appointed to the permanent research staff. He is assisted in the laboratory by his wife, Julie. His research interests centre on the palaeoceanography and stable isotope stratigraphy of the Palaeogene, the Cretaceous and the Silurian. He is consultant to the mass spectrometer manufacturer VG lsotech (a division of Fisons Instruments) for whom he has developed an individual acid-bath carbonate preparation and multiport system (the lsocarb II). He is currently UK national correspondent for IGCP Project 31)8 "Events around the Palaeocene/Eocene boundary" to which this paper is a contribution.