Microfossils stable isotopes and oceanatmosphere history


The skeletons of foraminifera and other CaCO3 fossils take up chemical signals from sea water as they grow. The most important of these chemical signals are the stable isotopes of oxygen and carbon. These signals, when extracted from the CaCO3 in a mass spectrometer, may be used to reconstruct past environmental changes such as temperature and ocean fertility, and to provide a high-resolution chemostratigraphy. The oxygen isotope technique was pioneered in the 1950s by Cesare Emiliani and the oxygen isotope stages that he initiated are now widely used as the basis for Quaternary and Tertiary stratigraphy (Figs 4.1,4.2). The technique can also be used to estimate palaeotem-perature, palaeosalinity and ice volume changes. The carbon isotope technique has been explored since the 1970s for carbon isotope stratigraphy and to provide information on the history of the carbon cycle and palaeoproductivity of the oceans.

Microfossils, especially foraminifera, are ideal for stable isotope research because they are easy to identify and readily checked for good preservation (using SEM), they have occupied a wide range of habitats and they can make up the bulk of oceanic sediments with a nearly continuous geological record.


Both oxygen and carbon isotopes can be obtained during an analysis of a single sample of CaCO3 (Box 4.1). The ratio between the heavier and lighter isotopes (i.e.

18O/16O and 13C/12C) is expressed as the delta (8) value in parts per thousand (%o). A standard sample is also

Holocene interglacial Déglaciation c. 10 kyr BP

Glacial maximum c. 18 kyr BP

Holocene interglacial Déglaciation c. 10 kyr BP

Glacial maximum c. 18 kyr BP

Interstadial c. 40 kyr BP

Ice volume Increases c. 60 kyr BP c. 82 kyr BP c. 103 kyr BP

Last Interglaclal c. 122 kyr BP Full glacial c. 150 kyr BP

Fig. 4.1 Changes in oxygen isotope ratios of epibenthic foraminiferid calcite tests through the last 150,000 years, showing fluctuations related to changing ice volume. Core 12392-1 North East Atlantic. (Modified from data in Brasier 1995.)

run so that comparisons can be made between runs or between machines. In carbonates, this standard was originally the calcite guard of a belemnite from the Late Cretaceous Pee Dee Formation of South Carolina, USA. Samples are now compared with the Pee Dee Belemnite (or PDB) via a second, usually in-house, standard such as Carrara Marble. Oxygen isotopes from modern oceanic waters are more usually calibrated against SMOW (standard mean oceanic water). The terms, heavy/light, positive/negative and





D C c

Planktonic foraminifera


Benthic foraminifera


A )

5 10

10 r

Fig. 4.2 Changes in oxygen isotope ratios of both benthic and planktonic foraminiferid calcite tests through the Tertiary, showing fluctuations related to changing water temperature and/or ice volume. Temperature estimates depend on assumed values for Sw in each period. Letters A-F refer to features discussed in the text. SMOW, standard mean ocean water. (Modified from data in Hudson & Anderson 1989.)

Box 4.1 Stable isotope analysis

1 Sample with microfossils is disaggregrated (e.g. using ultrasound) and dried.

2 Species of known habitat (e.g. surface water planktonic, infaunal benthic) are picked out for ecological studies. (Bulk samples of calcareous nannofossil carbonate from the <63 mm fraction can also give stratigraphically useful results.)

3 Specimens with evidence for secondary alteration (e.g. calcite overgrowths, pyrite) are rejected.

4 Specimens of the same size range are selected. Older machines may analyse from 1 to 40 planktonic foraminifera. Newer laser machines may analyse a single chamber.

6 The CaCO3of the test is reacted with phosphoric acid:

7 Liquid nitrogen is used to freeze the water and the CO2 gas. The frozen CO2 is transferred to the mass spectrometer at -100°C to draw off water and other impurities.

8 CO2 molecules are ionized and separated into ion beams of three different masses: 44 = 12C 16O 16O, 45 = 13C 16O 16O, 46 = 12C 16O 18O.

9 The ratio between different ion beams is measured: 46/44 gives the ratio 18O/16O, 45/44 gives the ratio 13C/12C.

10 These ratios are then compared with those in a reference CO2 gas.

11 The ratios are expressed as 5 values, according to equations (2) and (3) below:

(18O/16O) sample - ( 18O/16O) standard 518O = -------------- X1000

( 18O/16O) standard

5„r ( 13C/12C) sample - ( 13C/12C) standard 513C =- X1000

( 13C/12C) standard

12 To allow for global comparison, 518O results are expressed relative to the universal PDB standard or to the SMOW standard. These can be converted as follows:

518O (calcite SMOW) = 1.03086 518O (calcite PDB)

enriched/depleted indicate a relative increase/decrease in the heavy isotope (i.e. either 18O or 13C).

Oxygen isotopes

Five main factors affect the ratio between the stable isotopes 16O and 18O in CaCO3 skeletons (Box 4.2). For the influence of one of these to be calculated, the other five will need to be estimated or known. The results obtained have been applied to a wide range of geological problems, as discussed below.

The Quaternary icehouse

Microfossils from deep sea sediments have played a crucial role in the reconstruction of palaeotemperat-ure and ice volume changes over the last 100 million years. Emiliani (1955) used the S18O in planktonic foraminifera from deep sea cores to outline oxygen isotope stages for the Quaternary, believing these to reflect only surface temperature changes alone. It later became apparent that S18O can also be influenced by ice volume changes (Shackleton & Opdyke 1973). This is because expanding ice sheets lock up more of the lighter isotope 16O that falls as precipitation, and prevent it from returning to the sea (Box 4.2). In theory, the ice volume signal can be obtained from the S18O record of deep-water smaller benthic foraminifera, such as Uvigerina and Fontbotia (formerly called Cibicidoides), if it is assumed that stable temperatures prevailed in deep waters through the glacial-interglacial cycles (e.g. Shackleton 1982). It has been argued that the deep ocean has also experienced drops in temperature (Prentice & Matthews 1991), which makes assumptions about the volume of land ice, and about the Sw of ancient waters, more of a problem.

Figure 4.1 shows the S18O record obtained from deep sea benthic foraminifera in a DSDP core spanning the last 150,000 years. Isotope stages for glacial intervals take even numbers (e.g. 2, 4, 6) while those for warmer phases take odd numbers (e.g. 1, 3, 5 and

Box 4.2 Surface processes affecting oxygen isotopes

1 Isotopic composition of the water (5w, mean 518O). More 16O than 18O evaporates in H2O from the ocean, and more 16O than 18O precipitates as rain from clouds. In the standard model clouds tend to form from evaporation at low latitudes and move towards the poles, so that there is a continuous Rayleigh distillation, leading to enrichment in 16O of high latitude clouds and snow. A similar distillation takes place with altitude. H216O is therefore preferentially stored in polar icecaps. Carbonates precipitated in sea water at a time of higher ice volume therefore have a more positive 518O than found at times of lower ice volume. Salinity is similarly affected on a regional scale: fresh water has much more 16O than does sea water, for the reasons given above. Carbonates precipitated in fresh water therefore tend to incorporate more 16O and less 18O (and hence have a more negative 518O) than those precipitated in normal sea water. Carbonates precipitated in hypersaline waters generally have a more positive 518O.

2 Temperature. Carbonates precipitated in warmer water incorporate more 16O and less 18O (and hence have a more negative 518O) than those precipitated in cooler water. This results in a fractionation of about 0.22%o PDB per1°C.

3 Mineral phase. Aragonitic foraminiferid tests are enriched by 0.6% relative to calcitic benthic foraminifera, owing to differences in the vibrational frequencies of the carbonate ions. Mg calcite is also enriched in 18O relative to calcite by 0.06% per mole %MgCO3, at 25°C.

4 Vital effect. Many species do not secrete their CaCO3 in isotopic equilibrium with sea water owing to metabolic processes. This 'vital effect' varies between taxa from the same habitat. Smaller benthic and planktonic foraminifera and calcareous nannofossils are generally closer to equilibrium values than are larger benthic foraminifera, echinoderms and corals.

5 Diagenesis. 518O is easily reset by meteoric and burial diagenesis. Fluids tend to carry lighter isotopes and therefore make the ratios more negative. Specimens selected for study must therefore be free from diagenetic overgrowths (see Marshall 1992; Corfield 1995). Microfossils from ODP and DSDP cores are often but not invariably better preserved than those from exposed cratonic successions.

so on, back through time). Both glacial maxima and low sea level are inferred at c. 150,000 years BP (stage 6), and 20,000 years BP (stage 2). Rapid changes to full interglacial conditions with high sea levels took place at c. 122,000 years (stage 5e) and again at 10,000 years BP (stage 1). The increase in ice volume appears to have been prolonged, with episodic improvements to interstadial conditions at c. 103,000 years (stage 5c), 82,000 years (stage 5a) and 40,000 years (stage 3). A wide range of evidence, from reef terraces in New Guinea to ice cores in Antarctica, has supported this story.

Hays et al. (1976) showed that the regularity of Quaternary climatic oscillations was driven by changes in solar insolation brought about by the 'Milankovitch' orbital parameters of precession (~19 kyr), obliquity (~41 kyr) and orbital eccentricity (~100 kyr). Similar oscillations have been convincingly demonstrated back into the Mesozoic, and have been calibrated against the magnetostratigraphical scale for the Tertiary.

Shackleton & Kennett (1975) have inferred a lack of ice prior to this time, and that Sw was about —1%o. This assumption, however, gives cool tropical sea surface temperatures contrary to evidence from fossil distributions. Prentice & Matthews (1988, 1991) have argued that there is little evidence on which to base icevolume estimates and suggest that the benthic S18O values mainly record changes in the temperature of bottom waters. This problem is not yet resolved.

Oxygen isotope records have also been obtained from well-preserved microfossil materials in the Late Cretaceous (Jenkyns et al. 1994) when bottom waters appear to have been much warmer than at present. Diagenesis and burial metamorphism have generally reset the S18O values in older rocks exposed in cratonic successions. The emphasis therefore shifts away from microfossils towards robust and well-preserved macro-fossils and cements in sedimentary rocks.

The Tertiary oxygen isotope record

The Tertiary oxygen isotope record of benthic and plankonic foraminifera (Fig. 4.2) reveals general 18O enrichment through time. Low S18O values in benthic foraminifera from the Palaeocene (Fig. 4.2A) suggest that bottom waters were relatively warm, with a marked 'climatic optimum' in the Early Eocene (Fig. 4.2B). The fall in S18O of both surface and bottom waters through the Middle to Late Eocene, and the rapid fall in temperature at the Eocene-Oligocene boundary (Fig. 4.2C), has been attributed to falling temperatures. Part of the fall, however, may have been due to initial growth of the Antarctic ice cap (e.g. Zachos et al. 1992).

Both bottom-water and surface temperatures remained relatively cool though the Oligocene (Fig. 4.2D). The divergence between bottom and surface S18O values in the Middle Miocene (Fig. 4.2E) implies warmer surface waters and/or an expansion of ice sheets such as those in Antarctica (e.g. Prentice & Matthews 1988). The steep fall in bottom-water S18O in the Pliocene (Fig. 4.2F) has been taken to indicate the build-up of Northern Hemisphere land ice.

A major problem concerns assumptions about the Sw of sea water prior to the Middle Miocene.


In rivers and lakes, Sw depends on the altitude and temperature of the precipitate plus the effects of humidity and evaporation (Box 4.2, Fig. 4.3). Ostracod carapaces from glacial lakes, for example, show strongly negative S18O that can be used to reconstruct climate change through the Late Quaternary (e.g. Hammerlund & Keen 1994).

In brackish water estuaries and deltas, the Sw (mean S18O) of sea water is diluted by isotopically light 16O from rivers, so that S18O values of CaCO3 skeletons generally become more negative than in coeval sea water. Glacial meltwater, for example, brought in negative S18O values to the Gulf of Mexico via the Mississippi delta during the Pleistocene (e.g. Williams et al. 1989).

In hypersaline lakes, lagoons and restricted seas such as the Mediterranean Sea, the increased ratio of evaporation to precipitation means that 16O is removed, leaving both waters and CaCO3 enriched in 18O (e.g. Thunell et al. 1987). Seasonal evaporation of fresh water can produce a similar trend, as seen in larger benthic foraminifera across modern Florida Bay (Brasier & Green 1993). Primary productivity in marginal settings can be high, with considerable nutrient

Oxygen Stable Isotope Diagram

Fig. 4.3 Diagram illustrating how the stable isotopes of oxygen and carbon in microfossil skeletons will tend to vary with depth and salinity. Some typical genera are shown: 1, coccolithophorid in surface waters; 2, Globigerinoides in surface waters; 3, Globorotalia in intermediate waters; Fontbotia is epibenthic; Uvigerina is endobenthic. Black arrows indicate the most typical isotopic trends seen as environments become more extreme.

Fig. 4.3 Diagram illustrating how the stable isotopes of oxygen and carbon in microfossil skeletons will tend to vary with depth and salinity. Some typical genera are shown: 1, coccolithophorid in surface waters; 2, Globigerinoides in surface waters; 3, Globorotalia in intermediate waters; Fontbotia is epibenthic; Uvigerina is endobenthic. Black arrows indicate the most typical isotopic trends seen as environments become more extreme.

inputs from the land, so that bottom sediments tend to be organic rich and S13C values also tend to become more negative and highly variable, though exceptions to this rule are known.


Calculation of the palaeotemperature from skeletal carbonate (palaeothermometry) can be determined from the following equations:

Calcite: t(°C) = 16.9 - 4.4(Sc - Sw) + 0.10(Sc - Sw)2 (after Grossman & Ku 1986)

where Sc and SAr are the mean S18O of CO2 produced from calcite or aragonite respectively, by the reaction of phosphoric acid at 25°C, and Sw is the S18O of CO2 in equilibrium with water at 25°C, both versus PDB.

These equations assume that the value of Sw is known (i.e. that the salinity and ice volume are known) and that the vital effect is zero. Since temperature may vary seasonally, and some organisms may vary their water depth with growth (e.g. planktonic foraminifera), it is clear that bulk samples provide crude estimates of palaeotemperature but microsamples can give great precision.

Carbon isotopes

Carbon is not only an essential building block for life, it also modulates the climate of the planet (through CO2) and allows for oxygenation of the atmosphere (through photosynthesis and carbon burial). At the Earth's surface, carbon is mainly found in either the oxidized reservoir (as CO2, HCO- and carbonate minerals) or in the reduced reservoir (as organic matter, fossil fuels and native C). In the oxidized reservoir, the amount of dissolved CO2 and HCO- in the oceans is vastly greater (%) than that of CO2 in the atmosphere. The 'mixing time' for CO2 to pass through the atmosphere and ocean is about 1000 years. Carbon isotopic studies are beginning to reveal that the interchange between these reservoirs has seldom achieved a stable balance (Box 4.3).

There are two stable isotopes of carbon: 12C (98.9%) and 13C (1.1%). The 13C/12C ratio in the atmospheric CO2 gas (currently -7% PDB) is isotopically lighter than that of dissolved CO2 and HCO- in the oceans (currently -1% PDB) but an isotopic equilibrium is maintained between them because of the mixing effect of wind and waves. In an inert world, the ratio of 13C/12C in marine HCO- would closely reflect that of primordial mantle carbon, which still escapes in the form of volcanic CO and CO2 (-5% S13C PDB). In the living world, however, the 13C/12C ratio inclines towards a heavier value because autotrophs preferentially select the lighter isotope 12C during photosynthesis. Living organic matter therefore has an average S13C value of -26% PDB (i.e. strongly negative), and the S13C of the ocean and atmosphere are correspondingly depleted in 12C (i.e. positive).

Calcareous nannoplankton and some foraminifera living in surface waters secrete CaCO3 tests in which the S13C value is more or less in equilibrium with surface water HCO- (c. +2% PDB). A number of factors cause S13C values to vary, as shown in Box 4.3 and Figure 4.3. Beneath the photic zone, both the degradation of phytoplankton (especially by heterotrophic bacteria) and the release of respiratory CO2 result in the return of 12C to the water column. This can be seen in the more negative S13C of deeper water benthic foraminifera in the Atlantic (+1 to +0.5% PDB). In the modern Pacific, where the ocean width is large and the bottom waters are comparatively old, much oxygen has been removed during respiration so that the apparent oxygen utilization (AOU) index and the S13C values are correspondingly lower (-0.5 to +0.0% PDB). On the deep sea floor of modern oceans, the bottom waters, which have originated from shallow polar regions, help ventilation and also bring in more 13C. Beneath the sediment-water interface, bacterial decay of organic matter releases 12C-enriched CO2 back into the pore waters.

Gradients are therefore found in S13C through both the water column and the sediment (Fig. 4.3). These gradients tend to show an inverse relationship with oxygen and phosphorus concentrations, which is because organic degradation removes oxygen from the

Box 4.3 Surface processes affecting carbon isotopes

Surface water productivity. Where primary productivity is high, 12C is preferentially removed from the ocean and atmosphere. Raised productivity produces an increase in the A513C gradient between benthos and plankton and can result in temporal shifts towards more positive 513C. Biological oxidation. The respiration of organic matter in mid-waters and on the sea floor results in a return of 12C to the water column. Increased rates of biological oxidation will produce a decrease in the D513C gradient and can result in temporal shifts towards more negative 513C. Upwelling and mixing. Where 13C-depleted waters are brought up to the surface, as by upwelling, then 513C values of surface waters are correspondingly lowered (e.g. off Peru). A similar reduction in 513C can be brought about seasonally by the influence of summer stagnation on the open shelf (e.g. east coast USA) and by the influence of humic-rich fluvial or swamp waters in coastal regions (e.g. north of Florida Bay, USA). Microhabitat effect. A further gradient is found in sediments. Here, the 513C of pore waters becomes increasingly out of equilibrium with sea water values as depth below the sediment-water interface increases. This is because of the build up of respiratory CO2 and HCO-in pore waters.

5 Carbon burial. Factors which tend to raise the global proportion of organic matter buried in sediment are liable to raise the 513C of the whole ocean-atmosphere system. Such factors include raised primary productivity, increased mid- to bottom-water stagnation, and raised rates of sediment accumulation.

6 Vital effect. Taxa are known to differ in the proportion of HCO- taken in from sea water (c. 0%o PDB) and from cytoplasm (c. -28% PDB). Although foraminifera show much less vital fractionation than seen in echinoderms and corals, many species show a vital effect (e.g. due to photosymbiosis). Fractionation may even change during growth (e.g. larger rotaliids become more positive; larger miliolids become more negative; Murray 1991). Where possible, a single species and a single size fraction should be used for studies of trends through time.

7 Diagenesis. 513C is much less easily reset by meteoric and burial diagenesis than is 518O. Even so, most dia-genetic fluids tend to carry 12C and can therefore make the ratios more negative. Specimens selected for study must therefore be free from diagenetic overgrowths and cements (see Marshall 1992).

water column but returns both 12C and P. Such environmental gradients can be measured by calculating the difference in S13C (AS13C; called the 'delta del 13C') between surface water microfossils (e.g. Globigerin-oides spp. or calcareous nannoplankton) and a coexisting epibenthic species (e.g. Fontbotia wuellerstorfi), or between the latter and an infaunal taxon (e.g. Uvigerina sp.; Fig. 4.3).

The Quaternary carbon pump

The role of CO2 in climate change has been suspected since the nineteenth century. W.S. Broecker first suggested that carbon isotopes could provide a proxy for changing CO2 through the ice ages and Shackleton et al. (1983) were able to reveal the nature of this record. They found that the AS13C of planktonic-benthic foraminifera has varied markedly over the last 130,000 years in a way that can be tied to changes in ice volume revealed by the S18O record. AS13C proves to be greatest during glacial phases and least during interglacial phases (Fig. 4.4). This may be taken to infer that the partial pressure of atmospheric CO2 was least during glacial phases and greatest during interglacials, which has since been confirmed by direct measurements from ice cores. The changes in AS13C may also indicate major changes in primary productivity through the climatic cycle.

The Tertiary carbon isotopic record

Figure 4.5 shows the carbon isotopic record for much of the Tertiary. Note that this pattern is very different from the oxygen isotope record (Fig. 4.2) and shows no long-term trend. High S13C values of about 3%o PDB are found in the Late Cretaceous. A rapid fall to 2% took place across the K-T boundary and into the basal Palaeocene. At the K-T boundary, the AS13C fell

Vostok ice core

V19-30 E Pacific

12392-1 NE Atlantic

Fig. 4.4 Changes in the difference (A) between S13C values of planktonic and benthic foraminiferid tests ((a) core V19—30) and epibenthic and endobenthic foramimferid tests ((b) core 12392-1) through the last 150,000 years, showing fluctuations are related to changes in atmospheric CO2 ((a) Vostock ice core). (Modified from data in Brasier 1995.)

AS C%0PDB (planktonic/benthic)

A5 C%0 PDB (epibenthic/endobenthic)

Fig. 4.4 Changes in the difference (A) between S13C values of planktonic and benthic foraminiferid tests ((a) core V19—30) and epibenthic and endobenthic foramimferid tests ((b) core 12392-1) through the last 150,000 years, showing fluctuations are related to changes in atmospheric CO2 ((a) Vostock ice core). (Modified from data in Brasier 1995.)

to about 1%o (Fig. 4.5A), which has been taken to indicate the devastating effect of an extraterrestrial cometary impact on primary productivity (e.g. Hsu et al. 1982).

The Tertiary carbon isotope record shows evidence for two long-term cycles (Shackleton & Kennett 1975) with peaks in the Late Palaeocene and Middle Miocene. The Palaeocene provides a maximum S13C for the Tertiary of c. +3.5% (Fig. 4.5B) accompanied by a rise in the AS13C between planktonic and benthic foraminifera. This may have been due to high rates of productivity and carbon burial under greenhouse conditions.

A sharp fall in S13C occurred across the Palaeocene-Eocene transition (Fig. 4.5C) which was even greater than that across the K-T boundary. A mass extinction of about 50% of deep sea benthos took place at this time. The mid-late Eocene boundary interval was accompanied by a marked divergence in planktonic and benthic values and an increase in diatom abundance and diatom and dinoflagellate diversity. Together, these may be taken to indicate the effects of an increased thermal gradient on surface water productivity and carbon burial. Values were moderate during the Oligocene (Fig. 4.5D).

DSDP sites S28 and S29 South Atlantic

Fig. 4.5 The carbon isotopic record of the Late Cretaceous (K) to Tertiary, obtained mainly from planktonic microfossil carbonates in DSDP sites S28 and S29 of the South Atlantic. Letters A-E refer to features discussed in the text. P, Pleistocene; PLI, Pliocene. (Modified from Shackleton 1987.)

The S13C peak in the early-mid Miocene (Fig. 4.5E) coincided with widespread diatomites around the Pacific (the Monterey Event). The subsequent fall of c. 2.5%o to Recent values may owe much to the greater oxidation of organic matter brought about by cooler glacial oceans.

Further background information on isotopes can be found in the books on isotope geology by Faure (1986) and Hoefs (1988). Tucker & Wright (1990) and Marshall (1992) provide overviews from a sedimento-logical perspective. Williams et al. (1989) give an expanded discussion of Cenozoic isotope stratigraphy. Hudson & Anderson (1989) and Corfield (1995) review some of the achievements of oxygen isotope studies, while Murray (1991) reviews oxygen and carbon isotope data from benthic foraminifera. Brasier (1995) brings together stable isotope and other data used to interpret palaeoclimates and nutrient levels, while Purton & Brasier (1999) show how stable isotopes can be used to estimate changes in seasonality, ocean stratification, growth rate and life span.


Brasier, M.D. 1995. Fossil indicators of nutrient levels. 1: Eutrophication and climate change. Geological Society Special Publication 83, 113-132. Brasier, M.D. & Green, O.R. 1993. Winners and losers: stable isotopes and microhabitats of living Archaiadae and Eocene Nummulites (larger foraminifera). Marine Micro-palaeontology 20, 267-276. Corfield, R.M. 1995. An introduction to the techniques, limitations and landmarks of carbonate oxygen isotope palaeothermometry. Geological Society Special Publication 83, 27-42.

Emiliani, C. 1955. Pleistocene temperatures. Journal of

Geology 63, 538-575. Faure, G. 1986. Principles of Isotope Geology. John Wiley, New York.

Grossman, E.L. & Ku, T.L. 1986. Oxygen and carbon isotope fractionation in biogenic aragonite: temperature effects. Chemical Geology 59, 59-74. Hammerlund, D. & Keen, D.H. 1994. A Late Weichselian stable isotope and molluscan stratigraphy from southern Sweden. GFF116, 235-248. Hays, J.D., Imbrie, J. & Shackleton, N.J. 1976. Variations in the Earth's orbit: pacemaker of the ice ages. Science 194, 1121-1132.

Hoefs, J. 1988. Stable Isotope Geochemistry. Springer-Verlag, Berlin.

Hsu, K.J., McKenzie, J.A. & He, Q.X. 1982. Terminal Cretaceous environmental and evolutionary changes. Geological Society of America 190, special paper, 317328.

Fig. 4.5 The carbon isotopic record of the Late Cretaceous (K) to Tertiary, obtained mainly from planktonic microfossil carbonates in DSDP sites S28 and S29 of the South Atlantic. Letters A-E refer to features discussed in the text. P, Pleistocene; PLI, Pliocene. (Modified from Shackleton 1987.)

Hudson, J.D. & Anderson, T.F. 1989. Ocean temperatures and isotopic composition through time. Transactions of the Royal Society of Edinburgh: Earth Sciences 80, 183-192.

Jenkyns, H.C., Gales, A.S. & Corfield, R.M. 1994. Carbon and oxygen-isotope stratigraphy of the English chalk and Italian Scaglia and its palaeoclimatic significance. Geological Magazine 131, 1-34.

Marshall, J.D. 1992. Climatic and oceanographic signals from the carbonate rock record and their preservation. Geological Magazine 129, 143-160.

Murray, J.W. 1991. Ecology and Palaeoecology of Benthic Foraminifera. Longman, Harlow.

Prentice, M.L. & Matthews, R.K. 1988. Cenozoic ice volume history: development of a composite oxygen isotope record. Geology 17, 963-966.

Prentice, M.L. & Matthews, R.K. 1991. Tertiary ice sheet dynamic: the snow gun hypothesis. Journal of Geophysical Research 96(B4), 6811-6827.

Purton, L.M.A. & Brasier, M.D. 1999. Giant protist Num-mulites and its Eocene environment: life span and habitat insights from S18O and S13C data from Nummulites and Venericardia, Hampshire Basin, UK. Geology 27, 711-714.

Shackleton, N.J. 1982. The deep sea sediment record of climate variability. Progress in Oceanography 11, 199-218.

Shackleton, N.J. 1987. The carbon isotope record of the Cenozoic history of organic carbon burial and oxygen in the ocean and atmosphere. In: Brooks J.R.V. & Fleet A.J.

(eds) Marine Petroleum Source Rocks. Published for the Geological Society by Blackwell Scientific Publications, Oxford, pp. 423-435.

Shackleton, N.J. & Kennett, J.P. 1975. Paleotemperature history of the Cenozoic and the initiation of Antarctic glaciation: oxygen and carbon isotope analyses of DSDP sites 277, 279 and 281. Initial Reports Deep Sea Drilling Project 29, 743-755.

Shackleton, N.J. & Opdyke, N.D. 1973. Oxygen isotope and paleomagnetic stratigraphy of Equatorial Pacific core V28-238. Oxygen isotope temperatures and ice volumes on a 105and 106year scale. Quaternary Research 3, 39-55.

Shackleton, N.J., Hall, M.A., Line, J. & Shuxi, C. 1983. Carbon isotope data in core V19-30 confirm reduced carbon dioxide in the ice age atmosphere. Nature 306, 319-322.

Thunell, R.C., Willims, D.F. & Howell, M. 1987. Atlantic-Mediterranean water exchange during the Late Neogene. Paleoceanography2, 661-678.

Tucker, M.E. & Wright, V.P. 1990. Carbonate Sedimentology. Blackwell Scientific Publications, Oxford.

Williams, D.G., Lerche, I. & Full, W.E. 1989. Isotope Chronostratigraphy: theory and methods. Academic Press Geology Series, San Diego.

Zachos, J.C., Breza, J. & Wise, S.W. 1992. Early Oligocene ice-sheet expansion on Antarctica: sedimentological and isotopic evidence from Kerguelen Plateau. Geology 20, 569-573.

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