isotopes in the earth sciences || isotopes in palaeoclimatology

CHAPTER 12

Isotopes in Palaeoclimatology

12.1. BACKGROUND

The importance of isotopes in helping to establish the climates of the past is evident from previous chapters and is based upon the original work of H. C. Urey and his associates on 180/160 in the 1950s.1 The approach involved recent and fossil calcium carbonate shells and was extended later into studying the atmosphere and hydrosphere where the fractionation of both oxygen and hydrogen as different isotopic molecules of water is significant. Compared with average oceanic water, fresh water and glacial ice as well as snow on continents are enriched in the light stable isotopes 160 and IH, and thus possess negative 8 180 and 8D values. The enrichment in these isotopes increases with falling air temperature, therefore showing seasonal variation. In addition, there is both a latitudinal and an elevational variation. A positive correlation has been found between the isotopic composition of sea water and its salinity, and from this has come a better understanding of the source of cold water at the bottom of the oceans. The temperature variations in these bottom waters through geological time have been examined by analysing the oxygen isotopic compositions of CaC03, silica and phosphate: the results demonstrate that there has been an overall drop, with many minor fluctuations, from about 70° C 3·4 Ga ago until today. The formation of the vast continental ice sheet in Antarctica during the Miocene epoch is signalled by an increase in 8 180 of benthonic foraminifera.

A problem in such work is that assumptions are necessary concerning the isotopic composition of sea water in earlier times and regarding the existence of isotopic equilibrium. On land, climatic changes have been monitored by assessment of the oxygen isotope

508

R. Bowen, Isotopes in the Earth Sciences© Chapman & Hall 1994

ISOTOPES IN PALAEOCLIMATOLOGY 509

compositions of CaC03 in speleothems and in land snail shells as well as of calcium phosphate in mammalian and fish bones. Since oxygen is a major component of the majority of rock-forming minerals, if two minerals equilibrate its isotopes with a common reservoir, the difference in their 8 180 values decreases as the temperature increases; the phenomenon is used to determine the temperatures of final equilibration of cogenetic oxygen-bearing minerals. This subject has been discussed in Chapter 11.

Here, the emphasis is upon isotopic palaeoclimatology. Areas of particular interest have been selected for description because the subject has expanded so rapidly that space restrictions impose such limitations.

12.2. OXYGEN ISOTOPE COMPOSITION IN THE PAST

As seen earlier, atmospheric oxygen is not in isotope equilibrium with oceanic water (the Dole effect), and there is a considerable 180 enrichment (23· 5%0) relative to SMOW. Global oxygen consumption by respiratory processes may account for this. Although the effect is not completely understood quantitatively, it should give qualitative information, monitoring, particularly, major ecological changes in the biosphere of the Earth after free oxygen arrived in the atmosphere. The 180/60 ratio of atmospheric oxygen from the geological past is probably preserved by the oxide phases (mainly magnetite) of cosmic spherules; these are ablation droplets of iron meteorites oxidized in the atmosphere at high temperatures. Fossil examples have been located, the earliest being from the Early Cambrian. These strongly suggest that the oxygen isotope composition of the magnetite particles has remained substantially unaltered during the time interval involved. Adequate quantities from the later Phanerozoic (Devonian and Tertiary) have shown that there is no difference in the Dole value between the present and 30 Ma ago. This is expected, because the global processes of oxygen production and consumption are very unlikely to have changed during this time interval. In the Devonian, different circumstances operated because at that time (350 Ma ago) the oldest continental flora (mostly leafless psilophytes) had just appeared, the higher pteridophytes coming in only towards the close of the system. Consequently, the contribution of continental photosynthesis towards an 180 enrichment in the atmospheric reservoir must have been much smaller than it is today and

510 ISOTOPES IN THE EARTH SCIENCES

its5 18 value must have been much lower than was the case later. A value of ~ 17·3%0 has been proposed, this reflecting the reduced 180 release by contemporaneous land photosynthesis.

12.3. CRETACEOUS EXTINCTIONS

In Chapter 8, reference was made to the iridium anomaly at the Cretaceous-Tertiary boundary; palaeontological evidence shows that there was a Middle to Late Maastrichtian crisis for tropical organisms living in shallow waters, implying that the factors of environmental decline were already operative at that time. Concomitantly, a large-scale global biotic extinction manifested itself in the latest Cretaceous, together with a terminal regression in many parts of the world. By the early part of the Late Maastrichtian, all these processes broadly affected shallow shelf and platform habitats, diminishing the ecospace and increasing the competition as well as producing more variable turbidity, salinity and water movement in the surviving shallow-water Tethys. From Pacific foraminifera, oxygen isotope data indicate a significant decrease in marine surface-water temperatures through the earliest to middle Late Maastrichtian which coincided with the demise of the carbonate platform assemblage in Tethys.

Additional isotopic information from diverse benthic and pelagic organisms, especially from Atlantic foraminifera and Danish pelagic carbonate analyses, suggest highly variable Late Maastrichtian and boundary-zone temperature patterns.2-4 Various organisms with different habitats demonstrate distinct trends. Many benthic and midwater groups show small upward temperature movements of a few degrees Centigrade in the earliest Late Maastrichtian, but some planktonic foraminifera indicate a simultaneous decline in temperature which implies a decrease in the temperature gradient through the oceanic water column.5 It has been proposed that there was an overall broadening of the vertical Atlantic temperature gradients across the Cretaceous-Tertiary boundary which was expressed variably at different sites by a temperature decline of I-3°C (the larger figure referring to benthic waters) or, at one site, a 3°C rise in surface temperatures as opposed to a 2° C drop in benthic ones. B. Buckardt and N. O. Jorgensen in 1979 recorded major excursions in oxygen and carbon isotopic data across the C-T boundary in Denmark.4 If these are interpreted as temperatures, there was a l20 C fluctuation with a net loss of 4° C within


a few hundred thousand years and diminished oceanic oxygen levels. Probably, therefore, there were rapid and important temperature excursions through the Late Maastrichtian and across the boundary of certain tropical and north temperate sites, such events most likely having stressed stenothermal Tethyan taxa well before and at the end of the Maastrichtian.

Data from pelagic sequences imply spreading oxygen depletion in the terrestrial tropical oceans and the demise of the stenotopic benthic and mid-water biotas, commencing at roughly the Middle-Late Maastrichtian boundary. Of course, other environmental factors were involved too; acting in concert with those discussed, they may have stressed the tropical shelf biota at the end of the Middle Maastrichtian and begun the terminal Cretaceous biota crisis about 2 Ma prior to the end of that period.

Interestingly, a major short-term extinction occurring during the Cretaceous itself, actually defining the Turonian-Coniacian boundary in the Western Interior of North America, took place in conjunction with the onset of an oxygen crisis in the basin of which the end labelled maximum transgression, as E. G. Kauffman indicated in 1977.6

Versions of greater oxygen restriction perhaps played a role in the planetary Cretaceous extinction episode.7

12.4. CAMPANIAN-TO-PALAEOCENE PALAEOTEMPERATURE AND CARBON ISOTOPE SEQUENCE

These have been studied by A. Boersma and results were reported in 1984.8 They may be categorized with reference to surface and bottom temperatures.

12.4.1. Surface Temperatures The following planktonic foraminifera were used.

Maastrichtian: Rogoglobigerina or Pseudoguembelina. Palaeocene: Guembelitria, Acarinina, Morozovella.

An unusual situation was noted among the species of the very earliest Tertiary, namely that the surface zone seems to be dominated by the heterohelicids, to judge by their oxygen isotopic composition. Also, several species gave anomalously negative carbon isotope values. Later in the Palaeocene, the heterohelicids were apparently deep-dwelling,


TABLE 12.1 Pa1aeco1ogic isotope data

Foraminifera species

Pseudoguembelina excolata Globotruncana contusa Globotruncana mayaroensis Hedbergella monmouthensis Heterohelix pulchra

[PO (%0)

-0·S7 -o·n -0·43 -0·17 +0·27

[PO (%0)

+2·21 +2·21 +1·35 +2·43 +1-68

while other surface-dwelling species yielded more positive carbon isotope values.

At site number 384, there was little change in estimated surface temperature, early Maastrichtian values lying near 14° C, increasing to 15-16° C in the later Maastrichtian. Some representative data may be quoted (Table 12.1).8 Site number 151/152 gave Maastrichtian temperatures from three samples: they were near to 21 ° C. As will be seen below, these are cooler than in site 356 and occur in what was the Caribbean arm of the Tethys Ocean.

At site 356, Campanian values demonstrate ocean surface temperatures close to 30° C. The earliest Palaeocene sample was collected above a visible C-T discontinuity; it comes from the early part of the Globigerina eugubina zone and a temperature of almost 24 ° C was estimated from it. During this time interval, there seems to have been a hiatus in sedimentation. A temperature near 19° C in Zone PIa is registered by an unusual species, Tubitextularia cretacea, but Guembelitria lived below it, at temperatures near 16° C.

Some measurements were made on the Agulhas Plateau and gave temperatures of 15 to 16°C during the Late Maastrichtian. Here, there were insufficient Tubitextularia cretacea to measure. In the latest Maastrichtian of the Falkland Plateau, a single surface temperature measurement is near 15° C.

12.4.2. Bottom Temperatures These were made as far as possible on unispecific samples, although multi specific ones often had to be employed.

At site 384, the bottom temperature record for an estimated palaeodepth of 3500 m or so parallels the surface temperature record through the Maastrichtian and Palaeocene. In the Maastrichtian, the tempera-


tures were 10-Il ° C, slightly cooler than those in the Palaeocene (almost 12° C). At site 356, the estimated palaeodepth was almost 1000 m and bottom temperatures of the Late Maastrichtian reached 14°C near the boundary. On the Agulhas Plateau, estimates were near 9° C and on the Falkland Plateau almost 12° C at a similar palaeodepth.

12.4.3. Surface Carbon Isotopic Values In the Palaeocene Globigenna eugubina Zone and early Zone PI, there is a problem in interpreting the surface carbon isotopic values. Throughout the time interval, the planktonic foraminiferal species registering the warmest temperatures record more negative carbon values than cooler species. In the Recent and through most of the Tertiary, many of the warmest species record the most positive carbon isotope values, deeper (cooler)-dwelling ones recording more negative carbon isotope ratios. The carbon isotope difference noted between Recent Globigennoides sacculifer and average benthonic foraminifera in the Recent is approximately equivalent to the observed gradient in the water column (see W. S. Broecker in 1971 9). At present, the explanations for the anomalous carbon isotope profiles during the Early Palaeocene are sub judice; for this reason, A. Boersma used the highest carbon value in the samples involved.

The carbon isotope records at the surface at all the sites were very similar, those from the latest Campanian through the Early and Middle Maastrichtian clustering near + 2%0 with the values increasing through the Late Maastrichtian to the range +2, 3 to +2·6%0 in the Abathomphalus mayaroensis Zone immediately below the C-T boundary. Lowest values were found on the Agulhas and Falkland Plateaux and the Rio Grande Rise.

There is an important excursion in the planktonic carbon isotope ratios across the C-T boundary everywhere: they become markedly more negative by up to 1·5%0 in the earliest Palaeocene Globigenna eugubina Zone.

12.4.4. Surface-to-Bottom Carbon Isotope Differences The bottom temperature estimates are not as reliable when based upon samples with varying species composition, and this problem is complicated in bottom carbon isotope measurements because of the large carbon isotope variations between species. Differences from 0·7 to 1%0 between species are frequently found.8 All this has to be remembered in attempting an interpretation of the estimates of the carbon isotopic


gradient through the water column. It was demonstrated that through the Maastrichtian the average difference was ca 1·2%0; surface values range from +2%0 in the Early Maastrichtian to almost +2·5%0 in the Late Maastrichtian.

There was a major excursion of the carbon isotope gradient at the C-T boundary that caused a negligible gradient in and above the Globigerina eugubina Zone; at that time, the highest value for a planktonic foraminifer is equal to or less than the carbon values at the bottom. In spite of the uncertainty in bottom values, the extremely reduced gradient is still probably real. The warmest, i.e. the shallowest, foraminifera record very negative carbon values at the ocean surface. The anomalous values must reflect a major alteration in the carbon isotope distribution through the water column at the C-T boundary, according to A. Boersma in 1984.8

Carbon isotope gradients through the Early Palaeocene gradually increase as the surface-dwelling species show gradually more positive I3C values. However, deep-dwelling planktonics and the benthics do not demonstrate this same increase. Later, both the deep planktonics and the benthic foraminifera become almost 1%0 more positive and this results in the Late Palaeocene carbon excursion which was recorded by P. M. Kroopnick et al. in 1977.10 Subsequently, carbon values drop off slightly, but intermediate and deep values stay at the higher levels. The small surface drop accounts for the slightly reduced carbon gradients at this time.

The following inferences may be drawn from the researches described. Firstly, the main palaeotemperature change of the Maastrichtian took place between the Late Campanian and the Early Maastrichtian Guembelitria trican'nata Zone, not at the C-T boundary. Secondly, a major carbon isotope event occurred exactly at this boundary; the negative excursion in the carbon isotope values registered in the earliest Palaeocene Globigerina eugubina Zone did not occur in underlying Cretaceous sediments, but took place only at the C-T boundary. Thirdly, the surface temperature of the North Atlantic Ocean rose in the very Early Palaeocene, bottom temperatures decreasing to promote strong vertical temperature gradients through the water column. Finally, a marked increase in surface /5!3C values took place during the Palaeocene and, in its later stages, the bottom and intermediate waters also registered increased carbon isotope values resulting in a Late Palaeocene carbon excursion. Then both the absolute /5!3C


values and the range of values within the oceans were exceptionally great.

12.5. OXYGEN ISOTOPES IN NEOGENE MOLLUSCAN FOSSILS AND QUATERNARY FORAMINIFERA

Temporal changes in the oxygen isotopic composition of specimens of the bivalves Macoma and Pecten from four sections of Neogene rocks exposed in the Eel River Basin of northern California resulted from temperature and global glacier volume changes and were discussed by J. R. Dodd et al. in 1984. 11 They stated that the temperature declined approximately 10-15° C during the geological history of the basin in question and was probably 5-10° C during much of the Plio-Pleistocene. However, there are no obvious trends despite the fact that the depth of deposition decreased appreciably through that time interval. The reason for this uniformity may have been a strong upwelling of deep water which maintained low shallow-water temperatures. Samples collected from shallow water and locations further inland possess lower 180/60 ratios than do samples from other sections, either because these waters were warmer or because of dilution of sea water by isotopically light fresh water. A difference in fractionation of oxygen isotopes between calcite (Pecten) and aragonite (Macoma) fossils may well provide a method for determination of oxygen isotope-based palaeotemperatures capable of eliminating the factor arising from the isotopic composition of the water in which these animals lived. T.he carbon isotopic composition showed a vital effect and proved to be more variable in the nearshore samples. Related work was carried out on palaeotemperatures and glacially induced changes in the oxygen isotope composition of sea water during the Late Pleistocene and Holocene in the Tanner Basin, California; reference was made to this in Section 9.5. The researchers involved were M. I. Kahn et al. in 1981. They analysed the benthic foraminiferal species Uvigerina peregrina, of which specimens were collected from a piston core taken at a depth of 1207 m in the shallower seaward part of the basin. 12 The data obtained demonstrate that glacially induced changes in the isotopic composition of sea water attained a maximum of I '5%0 in the Late Pleistocene. Additional isotopic data from the planktonic foraminiferal species Globigerina quinqueloba, corrected for glacial and regional water-mass


effects, show an increase in mean surface water temperatures of ca 5° C from close to the end of the Wisconsin glaciation to the Holocene thermal maximum at ca 7500 BP. Actually, the isotopic minimum which was recorded by Uvigerina peregrina at some 40 000 BP is considered unusually low compared with that recorded from open-ocean cores.

12.6. COMPARISON OF ISOTOPES AND PLANKTON IN A LATE QUATERNARY CORE

Interesting information on this theme was provided by T. J. Crowley and R. K Matthews in 1983; their core was obtained from the stable central waters of the North Atlantic gyre and its uppermost 200 cm sampled at 5 cm intervals, with additional samples being taken every 2· 5 cm for some sectionsY These researchers set up a preliminary stratigraphy by measuring percentages of CaC03; the results showed three minima which coincided with 180 stages 2,4 and 6 of C. Emiliani in 1955.14 Only minor dissolution effects on the foraminifera were observed. CLIMAP procedures were utilized in processing samples. IS, 16

Foraminiferal taxa were identified following the system ofN. G. Kipp of 1976.17 Isotopic analyses were carried out using a VG Micromass 602D mass spectrometer: the relevant procedures were given by W. B. Curry and R. K Matthews in 1981.18 All involved the 212-250,um fraction of Globigerinoides rnber in order to reduce to a minimum noise arising from isotopic variations between different-size fractions.

This Mid-Atlantic Ridge core yielded some extremely important results. Its site on the western flank of the ridge under the central waters of the relevant gyre was selected because earlier palaeo-ecological studies showed that water-mass variations in gyre centres are less than in any other part of the global ocean and also because the degree of shell preservation in North Atlantic sediments is generally better than in other oceans. The most significant findings include a record of low sedimentation rate (1'2 cm/103 a) which covers 180 stages 2-6, and the results show a 1· 75%0 variation over the glacial cycle. The general shape of the 180 record agrees well with patterns from cores having higher rates of sedimentation. There is very little record modification by bioturbation, which probably reflects the location of the core in the lowproductivity region of the sub-tropical gyre (low delivery rates of organic matter to the floor of the sea ought to support smaller populations of burrowers). This attestation of the standard 180 record in a province


possessing a low rate of sedimentation demonstrates that geographical regions in which the planktonic 180 record is dominated by the ice volume effect can be extended from the equatorial to the sub-tropical oceans. Both regions are characterized by temperature stability, but the annual range experienced is of course greater in the sub-tropical. The conclusion was drawn that co-variant benthonic-planktonic 180 changes in the Tertiary period can be interpreted primarily in terms of ice-volume changes over a broader range of conditions than previously proposed.

With regard to Emiliani's isotopic stages, odd-numbered ones represented warmer time intervals, and even-numbered stages were cooler; transitions between them were termed 'terminations' by W. S. Broecker and J. van Donk in 1970.14. 19 Whilst Emiliani assumed that the isotope record mainly reflects variations in sea-surface temperatures, it was convincingly shown by W. Dansgaard and H. Tauber in 1969 as well as byN. J. Shackleton and N. Opdyke in 1973 that past variations in the 180ro ratio in planktonic and benthonic foraminifera primarily refer to changes in the isotopic composition of sea water owing to the growth and decay of the great continental ice sheets.14.20.21 In fact, the relationship between the isotope record and the global ice volume was used later as a stratigraphical tool, for example by CLIMAP members.22 This tool involves assumptions, however: namely that isotopic changes are synchronous worldwide, within the mixing time of the oceans and within the resolution of the deep-sea sediments involved. Consequently, benthic bioturbation will reduce the amplitude of such isotopic variations and diminish the record of climatic variability in the past, possibly obscuring short-term events. Another effect of bioturbation, Le. that its occurrence may offset the timing of isotopic events thereby entailing a decrease in the stratigraphical resolution, is discussed below.

12.7. STRATIGRAPICAL UNCERTAINTY ARISING FROM BIOTURBATION

W. H. Hutson of CLIMAP in 1980 effected oxygen isotope analyses of two species of planktonic foraminifera, namely Globorotalia infiata and Globigerinoides sacculifera, from a deep-sea core taken off the coast of Durban, South Africa.23 These revealed a discrepancy exceeding 8 cm in the depth of Termination II within it, showing a stratigraphical


discrepancy of ca 4500 years. A number of research workers studied the effect of bioturbation on deep-sea sediments by examining the distribution of tracers of virtually instantaneous events ('impulses') in deep-sea cores. The vertical distribution of the tracer within a core can be alluded to as an impulse response function, which represents the response of the original impulse after it has been vertically mixed by benthic organisms. It is significant because it represents a numerical weighting function. Hutson presented a bioturbation model and chose for this a simple impulse response function similar in general shape to those empirically derived from analysis of ash layers in deep-sea cores.23 This produced an offset in isotopic curves similar to that observed in his core. It was concluded that variations in species abundance contributes to stratigraphical uncertainty and should be considered in stratigraphical interpretation of the oxygen isotope record.

12.8. FURTHER FORAMINIFERAL WORK

The Tertiary 8 180 record of planktonic and benthic foraminifera has been interpreted by R. K. Matthews and R. Z. Poore in 1980 in terms of a comparison of the relevant data with those from the average Late Pleistocene and assuming a constant tropical sea-surface temperature.24

From this, they estimated global ice volume; the approach suggested that the Earth had a significant ice budget (hence glacio-eustatic sealevel fluctuations), at least since the Eocene and maybe even throughout much of Cretaceous time as well.

In 1986, M. L. Delaney and E. A. Boyle studied the implications of lithium in foraminiferal tests as regards high-temperature hydrothermal circulation fluxes and oceanic crustal generation rates.25 The lithium content in planktonic foraminiferal calcite was analysed in order to evaluate the temporal variability of sea-water lithium concentrations over the past 116 Ma. The mean foraminiferal calcite Li/Ca in each time interval is no more than 16% greater nor 25% less than the mean Li/Ca ratio of all samples. Minima which are 25% lower than in adjacent time intervals occur between 50 and 60 Ma as well as between 80 and 90 Ma ago, with maxima at other times. At no time in the past 40 Ma does mean Li/Ca seem to be higher than at present. Coupled with an oceanic mass balance model for lithium, this implies that oceanic lithium concentrations and therefore high-temperature hydrothermal circulation fluxes during the past 40 and perhaps 100 Ma did not exceed 30 to 40%


more than the present levels for time intervals any larger than a maximum of I Ma. It also implies that the fluxes in question were not a factor of two higher 100 Ma ago.

12.9. OSTRACODS

J. T. Durazzi in 1977 presented interesting data regarding the stable isotopes in Recent ostracod shells from six localities off Florida?6 It was demonstrated that the (j 180 results accord with precipitation of the shells in isotopic equilibrium with sea water, and no strong correlation with carbon isotope ratios, temperature or salinity was observed.

12.10. SOME CARBON-13/CARBON-12 DATA

There appears to have been a major oceanographic change on Earth during the Miocene epoch, the oxygen isotope ratios of benthic foraminifera increasing markedly between 15 and 14 Ma (Middle Miocene) and indicating a significant cooling of deep ocean waters and a growth of the Antarctic continental ice sheet. Maximal (jISO values were attained during the Late Miocene when, some 8 Ma ago, deep-water temperatures and ice volume were probably similar to those now. The (jl3C records show a negative shift at about 6 Ma ago.

The average value of 0 BC of oceanic HCO; is controlled by the fluxes of 13C-depleted (reduced) carbon and l3C-rich (oxidized) carbon into and out of the ocean, and also by redox reactions in organic materiills within the ocean. Major controls of the regional distribution of (jl3C values of deep waters are the Ol3C values of HC03 formed prior to isolation of waters from the surface, and the progressive addition to them of bicarbonate formed by the oxidation of l3C-depleted organic matter. In the course of marine transgressions, huge quantities of l3C-depleted organic carbon are deposited and preserved on the continental shelves and in lOW-lying areas of the coast: this results in the l3C enrichment of dissolved marine HCO;. On the other hand, during regressions, less organic matter will be buried on the narrower continental shelves and the (jl3C of the marine HC03 decreases correspondingly. Also, erosion and oxidation of carbon in organic-rich shelf sediments after a marine offlap further contribute to lowering of oceanic HCO; ol3e.

It has been indicated that all Holocene benthic foraminifera seem to deposit calcium carbonate out of eqUilibrium with the bicarbonate in


the sea water overlying the sediment; inferred departures from equilibrium are fairly consistent from place to place for the majority of species.27 Consequently, changes downcore in the l3cl2C ratio of a species of benthic foraminifer reflect, at least primarily, changes in this ratio in the dissolved HC03 in the water in which it grew. In 1985, F. Woodruff and S. M. Savin investigated this matter by studying 166 samples from 21 DSDP localities and a piston core in the Miocene of the Pacific Ocean?8 Most of the relevant isotopic data were obtained from monospecific samples of the genus Cibicidoides which has been reported as depositing CaC03 with a S13C value of a few tenths per thousand lower than that of the dissolved HCO; of the water overlying the sediment. The Miocene benthic foraminiferalS l3C record demonstrated a consistent pattern at all depths through geological time. At almost every one of the sites involved, there was a 16-15 Ma S 13C maximum (1-4%0 to 1'8%0) and a 5·6 Ma minimum forS 13C (-0·2%0 to -0'7%0). These are extremes, however, and otherwise there is a uniformity of values over practically the whole of the Pacific and over water depths which ranged from 1·5 to 4 km. The major features are believed to reflect changes in Pacific and perhaps global geochemical cycling of carbon. The similarity of the S 13C values from geographically widespread sites in each Miocene time interval is considered to accord with measurements of this parameter in marine dissolved inorganic carbon (TDC) at many Pacific GEOSECS stations.29 The measurements suggest that the range of modern Sl3C values at the localities examined do not vary by more than 0'6%0.

The main features of the S13C record correlated with the marine onlap/offlap curve ofP. J. Vail andJ. Hardenbol in 1979, which showed a very high sea level between 16 and 13 Ma and a very low one between 7 and 5 Ma.30 Possibly this means that changes in the delivery of organic carbon to the deep sea are directly related to sea-level changes and critically affect the Miocene benthic S\3C record.

Of course, organic carbon deposited on the shelves is especially liable to erosion and later oxidation; hence the onset of a regression could well be characterized in the carbon isotope record by a short-term low S13C spike (cf. the very low S l3C minimum at 5·6 Ma mentioned above).31 If the correlation of the main features of the Miocene benthic foraminiferal oxygen isotope time series curves discussed by E. Barrera et al. in 1985 with those of the onlap/offlap curve mentioned above is valid, this would provide a mechanism for rapid sea-level lowerings through


inferred glaciations in the Miocene.32 In addition, the global albedo is high when the sea levels are low, and this promotes global cooling. Unfortunately, at almost every site, the correlation between 8 180 and onlap/offlap was poorer than that between 8l3C and onlap/offlap.28 Differences in the 813C between the sites are correlated with local differences in biological productivity in the overlying surface waters. Clearly, l3e/l2e values of benthic foraminifera are potentially useful indicators of marine palaeoproductivity.

In later, glacial times, changes have been recorded in the 8 I3e values from glacial to interglacial stages. Benthic foraminifera living in the former have lower I3e;t2c ratios by ca 0'7%o? This implies that reservoirs of organic carbon must have been smaller during glacial periods, reflecting a sea level fall of ca 140 m which left enlarged shelves subject to erosion. As the sea rose, reflooding them, a net deposition could have resulted. This could contribute to the glacial-to-interglacial cycle in organic carbon storage, and therefore to the S l3e change noted in benthic foraminifers. At the close of glacial time, removal of Le02

from the ocean in the form of organic matter induced a 2 km drop in the aragonite lysocline. Although the relative contributions of terrestrial reservoirs (trees, soils) and marine reservoirs (shelf sediments) to the removal of carbon from the ocean-atmosphere reservoir at the same time are not known, probably part of the loss was due to the terrestrial reservoirs.

A problem with foraminifera is that different species demonstrate a different fractionation factor due to vital effects (ironically defined by W. S. Broecker in 1982 as a name used by geochemists toconcenl the fact that 'we don't understand them', in relation to work on the geochemistry of oceanic phosphate34). This is significant because, if deep sea water were to be brought to the surface and warmed without any biological activity (i.e. heated isochemically), its CO2 partial pressure would be ca three times higher than that in the atmosphere, the removal ofLC02 by plants drawing this down to an extent determined by the POi-ILC02

ratio in deep water. The vital effects mentioned prevent an estimation of the difference in the 8 l3e between glacial phosphorus-free surface water and glacial mean deep water directly from the difference between the S l3e values for the tests of planktonic organisms grown in phosphatefree surface water and for the mean tests of benthic organisms grown in various places in the deep sea. However, the difference can be determined indirectly from the following equation:


(c5i3CPFSW - c5 i3CMDw)a = (c5 13CPFSW - c5 13CMDWhoD

+ (c5 13Ca - c5 13CI)plankt - (c5 i3Ca - c5 i3CI)benthic

(12.1)

where the subscript PFSW denotes phosphate-free surface water, MDW mean deep water, and TOO today, and the term on the left-hand side is the one required. On the right-hand side, the first term is the measured difference between phosphorus-free surface water and mean deep water in the oceans now, the second the glacial-to-interglacial change for a given species of planktonic foraminifer as measured in individual deepsea cores, and the third the glacial-to-interglacial change for a given species of benthic foraminifer as measured in individual deep-sea cores. The procedure obviates the vital effect for a given species if this parameter is assumed to be constant from glacial to interglacial time.34

As noted earlier, the glacial-to-interglacial c5 13C difference for benthic species averages O· 7%0, the Bcl2c ratio being lower during glacial times. The same effect has been observed in planktonic species, but the difference in their case is far smaller, not much over a zero value. Of course, it is not possible to determine how glacial planktonic foraminifers differed from their interglacial forebears or descendants either in vital effect or depth habitat, so caution is essential in the matter.

Measurements of CO2/air ratios in gas trapped in bubbles in glacialage ice imply that the CO2 content of the atmosphere then was much lower than in the Holocene, perhaps because of changes in the nutrient element chemistry of sea water. W. S. Broecker in 1982 proposed two feasible scenarios.34 One entails alternate storage and erosion of phosphorus, leaving residues from shelf sediments and the second involves alterations in the C/P ratio in organic debris falling down to the deep sea. The data presented above (Section 12.4) concerning both planktonic and benthonic foraminifera, as well as the oxygen record inferred from distributions of the latter and the early post-glacial CaC03 preservation event given by the aragonitic pteropods, accord with either. However, if an early post-glacial spike in the i3C record for planktonic shells could be found, it would be possible to eliminate the hypothesis entailing shelf storage.

Broecker pointed out important implications of the nutrient hypotheses.34 If the shelf storage caused the glacial-to-interglacial CO2

increase, then this change represents an amplifier of some primary cause. This is because sea-level changes are necessary to drive deposition on to and erosion from the shelves. Alternatively, if the

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524 ISOTOPES IN THE EARTH. SCIENCES

changes in the C/P ratio for falling debris are responsible, then the CO2

change could either be an amplifier or a primary cause for the major glacial-to-interglacial climatic cycle. This latter is possible, as selfsustained oscillations in oceanic chemistry could perhaps be driven by interactions between ocean ecology and ocean nutrient chemistry.

Interestingly, if there were no life in the oceans, their surface chemistry could be expected to change over a few tens of years into a type reflecting the isochemical heating mentioned earlier. Table 12.2 illustrates this change. Broecker referred to it as the 'Strangelove' effect and mentioned that such a disaster might have taken place at the end of the Cretaceous time interval as the result of the impact of either an asteroid or a comet on the Earth; see Section 8.2.5.34

12.10.1. Changes in the Oceanic Carbon-13/Carbon-12 Ratio During the Last 140 000 Years

L. D. Labeyrie and J. C. Duplessy in 1985 compared records of the carbon and oxygen isotope ratios of Neogloboquadrina pachyderm a between nine high-latitude sediment cores from the northern and southern hemispheres covering the last 140000 years. The strong analogies between the O\3C ratios allowed an appropriate O\3C stratigraphical scale to be defined.35 In this, there are three transitions which coincide with the oxygen isotope transitions 6/5 (125 000 years), 5/4 (65000 years) and 2/1 (13 000 years). The S13C records of this species in the high-altitude cores follow the changes in O\3C of the surface water TDC near areas of deep-water formation; they present trends similar to the benthic foraminifera 0 \3C records in other cores. However, the amplitude of the isotopic shifts varies, which implies that a great deal of the variability noted is connected with global changes in the carbon distribution between the biosphere and the ocean.

The l3C/J2C ratios of the species in the North Atlantic cores display larger regional variations at 18000 years BP than now. These may be attributable to the activity of a strong divergent cyclonic cell centred roughly 55°N and 15°W during the last ice maximum. This is reminiscent of the existing Weddell Sea.

12.10.2. The Carbon-13/Carbon-12 Ratio Now An interesting study was made in 1985 by T. Torgersen and A. R. Chivas, who examined the Ol3C values of organic material in marine sediments off Hinchinbrook Island, Australia, and found a widespread (> 230 km2)

influence of terrestrial organic carbon, Cr.36 Extensive mangrove forests on this island are known to export <50 Mg CT day-l of leaf litter and


7· 5 Mg Cr day-I of microparticulate organic carbon. A sedimentation rate of 13-29 Mg CT day-I was calculated from 210Pb sedimentation rates, organic carbon contents and the stable isotope ratios of organic matter in the sediment (1 Mg = 106 g). The balance shows that terrestrial organic carbon in the marine environment can be accounted for by terrestrial export, although there are significant losses as well. The shoreline re-deposition of floating leaves and the like, as well as organic matter degradation to CO2 and/or CH4 succeeded by bubble ebullition, probably constitute major loss terms. Torgersen and Chivas considered that leaf litter could well be the most important single source of terrestrial organic carbon in the marine environment here, although they realized that mangrove represents an extreme instance of terrestrial productivity.36

Depth-related changes in the l3C/12C ratio of skeletal carbonate deposited by a Caribbean reef-frame building coral, Montastrea annularis, were investigated in 1976 by J. N. Weber et alY They correlated systematic variations in the isotopic composition of such. skeletal carbonate with water depth, location of the corallites within the corallum and polyp packing density, this being based upon 426 samples. Coral skeletons are not isotopically homogeneous, but when large samples are homogenized and analysed the within-sample variability becomes rather unimportant. In any case, for oxygen isotope palaeotemperature measurement, at least one year of growth must be included in the analysed sample if the mean annual water temperatures are to be estimated reliably. However, many reef corals extend their colonies at different rates in different directions and systematic isotope ratio variations within a corallum can be useful in assessing what effect, if any, growth rate may have on skeletaI8 l3e. The present species was especially interesting because of major changes in colony form as a function of depth. Main results of the work were to show that, in Acropora cervicornis, there are small-scale systematic variations in 8 l3C and 8180 within a given coral skeleton, and to compare top and side (edge) of Montastrea annularis colonies. Their forms change with depth and their tops seem invariably, at least at all the depths examined, to be relatively enriched in l3C with respect to the sides; the difference is greatest at 4·6 m. At 18·3 m the effect was negligible. The same relationship was found to apply for 8180, although in this case the differences between tops and sides were much smaller than for 8 l3e. There is a positive correlation between the latter parameter and polyp packing density. In addition, there is a decrease in8 l3e with depth down to 18·3 m which is succeeded at 22·9 m by an abrupt shift to higher 0 l3C


values, in turn followed by a slight decrease at 27-4 m. The 8 180 values demonstrate a similar trend, but interpretation of these data is rendered difficult because of the temperature dependence of the 180/60 ratio. Weber and his associates inferred that growth rate as such does not exert any significant influence upon skeletal isotopic composition and that the results mentioned above suggest that there are two ecotypes of Montastrea annularis involved, i.e. a deep-water and a shallow-water one.37 Further interesting observations showed that in Acropora cervicornis growth rate and 8 13e are inversely correlated.

12.11. 8 J3e AND ANIMAL DIETS

M. J. DeNiro and S. Epstein in 1978 investigated the effects of diet on the distribution of carbon isotopes in animals.38 They stated that the isotopic composition of the whole body of an animal is related to the isotopic composition of its diet, but the animal is enriched in 813e on average by ca I %0. In some cases which they examined, this is counterbalanced by the Be depletion arising from the respired e02•

The isotopic relationships between the whole bodies of animals and their diets are similar for different species raised on the same diet and for the same species raised on different diets, but the 8 13e values of whole bodies of individuals of a species raised on the same diet may differ by as much as 2%0. The relationship between the Be/12e ratio of a tissue and that of the diet depends upon the type of tissue as well as the nature of the diet involved. Many of the isotope relationships between major biochemical fractions (lipid, carbohydrate and protein) are qualitatively preserved as diet carbon becomes incorporated into the animal. However, it was found that the difference between the 8 13e values of a biochemical fraction in an animal and in its diet can be as great as 3%0. The 8 13e values of the biochemical constituents collagen, chitin and the insoluble organic fraction of shells, all of which are frequently preserved in fossil material, relate to the isotopic composition of the diet. DeNiro and Epstein stated that it is feasible to perform dietary analysis based upon the determination of the l3e/12e ratio of animal carbon, opining that analysis of the total animal carbon will usually give a better measure of diet than the analysis of individual tissues, biochemical fractions or biochemical components.38

This isotopic method of dietary analysis is applicable to fossil material providing that the original 13e value of some component has been preserved. Often, the best-preserved animal carbon is included in


the carbonate fraction of either inve11ebrate shells or vertebrate bones. In both marine and non-marine molluscs, the ol3e values of the shell carbonates reflect mostly those of the e02 dissolved in the water rather than the diet ol3e values and indeed aquatic carbonate-depositing organisms of other phyla may well obtain the carbon in their shell carbonate from this source too. In terrestrial snails, the carbon in their shell carbonate derives from an uncertain source. Even if there is some dietary influence involved, this will be hard to evaluate because of the possible contribution of carbon from the atmosphere and the sure ingestion of eaeo3• Both of these have ol3e values far more positive than dietary sources, which would be swamped out in the ol3e value of the resultant shell carbonate. A similar situation is encountered in the case of vertebrate bone because the source of the carbon in the relevant carbonate fraction is again uncertain. DeNiro and Epstein suggested that there may be some dietary influence on the basis of analyses of bones obtained from mice raised on known diets.38 However, measurements of radiocarbon in fossil bones demonstrated that the carbon in the carbonate fraction is exchangeable with the e02 dissolved in groundwater and/or atmospheric e02 so that the originalo I3C value of the carbonate fraction is unlikely to be preserved in fossil bone.

Important to palaeoclimatology is the observation that the relative amounts of potential diet sources ingested by an animal are determinable from the Ol3C value of the carbon in the animal if these sources possess sufficiently differento 13C values. For instance, the ol3e values of aquatic plants and animals frequently do not overlap those of terrestrial organisms, and it should be feasible to utilize this difference i!1 order to ascertain the relative contributions of these two kinds of organisms to the diets of animals living in near-shore environments. In addition, C3

and C4 plants act as potential diet sources. The 0 13e values of most ofthe former range from -24%0 to - 34%0, while those of the latter lie between -6%0 and -19%0, a difference sufficiently large to enable the relative quantities of C3 and C4 plants eaten by an animal to be determinable from the ol3e value of its carbon. These plant categories have been discussed in detail in Section 10.2.

REFERENCES

1. EpSTEIN, S., BUCHSBAUM, R., LOWENSTAM, H. A and UREY, H. C, 1953. Revised carbonate-water isotopic temperature scale. Bull. Geol. Soc. Amer., 64, 1315-6.


2. KAUFFMAN, E. G., 1979. Cretaceous. In: Treatise on Invertebrate Paleontology, Part A, Ed. C. Teichert and R. A Robison. University of Kansas Press, Geol. Soc.Amer.

3. BOERSMA, A and SHACKLETON, N., 1979. Some oxygen and carbon isotope variations across the Cretaceous/Tertiary boundary in the Atlantic Ocean. In: Cretaceous-Tertiary Boundary Events, Copenhagen Univ. Proc. Symp., Ed. W. K Christensen and T. Birkelund, Vol. 2, pp. 50-3.

4. BUCKHARDT, B. and J0RGENSEN, N. 0., 1979. Stable isotope variations at the CretaceouslTertiary boundary in Denmark. In: Cretaceous-Tertiary Boundary Events, Copenhagen Univ. Proc. Symp., Ed. W. K Christensen and T. Birkelund, Vol. 2, pp. 54-61.

5. DOUGLAS, R. G. and SAVIN, S. M., 1973. Oxygen and carbon isotope analysis of the Cretaceous and Tertiary foraminifera from the central North Pacific. In: Initial Reports of the Deep Sea Drilling Project, Vol. 17, pp. 591-605. US Govt Printing Office, Washington, DC.

6. KAUFFMAN, E. G., 1977. Geological and biological overview: western interior Cretaceous Basin. In: Cretaceous facies, faunas, and paleoenvironments across the Western Interior Basin, Ed. E. G. Kauffman, Rocky Mt. Assoc. Geol. Mountain Geol., 14(3,4), 75-99.

7. KAUFFMAN, E. G., 1984. The fabric of Cretaceous marine extinctions. In: Catastrophes and Earth History, Ed. W. A Berggren and 1. A van Couvering, pp. 151-246. Princeton University Press.

8. BOERSMA, A, 1984. Campanian through Paleocene paleotemperature and carbon isotope sequence and the Cretaceous-Tertiary boundary in the Atlantic Ocean. In: Catastrophes and Earth History, Ed. W. A Berggren and J. A van Couvering, pp. 247-78. Princeton University Press.

9. BROECKER, W. S., 1971. A kinetic model for the chemical composition of sea water. Quat. Res., 1, 188-207.

to. KROOPNlCK, P. M., MARGOLIS, S. and WONG, c., 1977. CSl3C variations in marine carbonate sediments as indicators of the CO2 balance between the atmosphere and oceans. In: Fate of Fossil Fuel CO2 , pp. 295-321. Plenum, New York.

11. DODD, J. R., STANTON, R. J., Jr and JOHNSON, M., 1984. Oxygen isotopic composition of Neogene molluscan fossils from the Eel River Basin of California. Bull. Geol. Soc. Amer., 95, 1253-8.

12. KAHN, M. I., OBA, T. and Ku, T. L., 1981. Paleotemperatures and the glacially-induced changes in the oxygen-isotope composition of sea water during late Pleistocene and Holocene time in Tanner Basin, California. Geology, 9, 485-90.

13. CROWLEY, T. J. and MATTHEWS, R. K, 1983. Isotope-plankton comparisons in a late Quaternary core with a stable temperature history. Geology, 11, 275-8.

14. EMILIANI, c., 1955. Pleistocene temperatures. J Geol., 63,538-78. 15. IMBRIE, J. and KIpp, N. J., 1971. A new micropaleontological method for

quantitative paleoclimatology: application to a late Pleistocene Caribbean core. In: The Late Cenozoic GlacialAges, Ed. K K Turekian,pp. 71-179. Yale Univ. Press, New Haven, CT.

16. IMBRIE,J., VAN DONK,J. and KIpp, N.J., 1973. Paleoc1imatic investigation of


a late Pleistocene Caribbean deep-sea core: comparison of faunal and isotopic methods. Quat. Res., 3, 10-38.

17. KIPP, N. 1., 1976. New transfer function for estimating sea-surface conditions from sea-bed distribution of planktonic foraminiferal assemblages in the North Atlantic. In: Investigations of Late Quaternary Paleoceanography and Paleoclimatology, Ed. R. M. Cline and 1. D. Hays, pp. 3-42. Geol. Soc. Amer. Memoir 145.

18. CURRY, W. B. and MATIHEWS, R. K, 1981. Paleo-oceanographic utility of oxygen isotopic measurements on planktonic foraminifera: Indian Ocean core-top evidence. Paleogeog., Paleoclimat. Paleoecol., 33, 173-91.

19. BROECKER W. S. and VAN DONK, 1., 1970. Insolation changes, ice volumes and the Ol8 record in deep-sea sediments. Rev. Geophys. Space Phys., 8, 169-98.

20. DANSGAARD, W. and TAUBER, H., 1969. Glacier oxygen-18 content and Pleistocene ocean temperatures. Science, 166,499-502.

21. SHACKLETON, N. 1. and OPDYKE, N., 1973. Oxygen isotope and paleomagnetic stratigraphy of equatorial Pacific core V28-238: oxygen isotope tempera tures and ice volumes on a 105 year and 106 year scale. Quat. Res., 3, 39-55.

22. CLIMAP PROJECT MEMBERS, 1976. The surface of the ice-age Earth. Science, 191, 1131-7.

23. HUTSON, W. H., 1980. Bioturbation of deep-sea sediments: oxygen isotopes and stratigraphic uncertainty. Geology, 8, 127-30.

24. MATIHEWS, R. K and POORE, R. Z., 1980. Tertiary8 l80 record and glacioeustatic sea-level fluctuations. Geology, 8,501-4.

25. DELAl"lEY, M. L. and BOYLE, E. A, 1986. Lithium in foraminiferal shells: implications for high-temperature hydrothermal circulation fluxes and oceanic crustal generation. Earth Planet. Sci. Lett., 80, 91-105.

26. DURAZZI, J. T., 1977. Stable isotopes in the ostracod shell: a preliminary study. Geochim. Cosmochim. Acta, 41, 1168-70.

27. GRAHAM, D. W., CORLISS, B. H., BENDER, M. L. and KEIGWIN. L. D.. Jr, 1981. Carbon and oxygen disequilibria of Recent deep-sea benthic foraminifera. Mar. Micropal., 6, 483-97.

28. WOODRUF, F. and SAVIN, S. M., 1985. 8l3C values of Miocene Pacific benthic foraminifera: correlations with sea level and biological productivity. Geology, 13, 119-22.

29. KROOPNICK, P. M, 1984. The distribution of C-13 of TC02 in the world oceans. Deep-Sea Res., 32(1),57-84.

30. VAIL, P.1. and HARDENBOL, 1.,1979. Sea-level changes during the Tertiary. Oceanus, 22, 71-9.

31. BERGER, W. H., 1982. Deep-sea stratigraphy: Cenozoic climatic steps and the search for chemo-climatic feedback. In: Cyclic and Event Stratification, Ed. G. Einsele and A Seilacher, pp. 121-57. Springer-Verlag, Berlin.

32. BARRERA, E., KELLER, G. and SAVIN, S. M., 1985. Evolution of the Miocene ocean in the eastern north Pacific as inferred from oxygen and carbon isotope ratios of foraminifera. In: The Miocene Ocean: Paleoceanography and Biogeography, Ed. J. P. Kennett. Geol. Soc. Amer. Memoir.

33. SHACKLETON, N. J., 1977. Tropical rainforest history and the equatorial


Pacific carbonate dissolution cycles. In: The Fate of Fossil Fuel CO2 in the Oceans, Ed. N. R. Anderson and A Malahoff, pp. 401-28. Plenum Press, New York.

34. BROECKER, W. S., 1982. Ocean chemistry during glacial time. Geochim. Cosmochim. Acta, 46, 1689-705.

35. LABEYRIE, L. D. and DUPLESSY, 1. C, 1985. Changes in the oceanic 13CPC ratio during the last 140 000 years: high-latitude surface water records. Paleogeog. Paleoclimat. Paleoecol., 50,217-40.

36. TORGERSEN, T. and CHIVAS, A R., 1985. Terrestrial organic carbon in marine sediment: a preliminary balance for a mangrove environment derived from Be. Chem. Geol., 52, 379-90.

37. WEBER,1. N., DEINES, P., WEBER, P. H., BAKER, P. A, 1976. Depth related changes in the BC/12C ratio of skeletal carbonate deposited by the Caribbean reef-frame building coral Montastrea annularis: further implications of a model for stable isotope fractionation by scleractinian corals. Geochim. Cosmochim. Acta, 40,31-9.

38. DENIRO, M. J. and EpSTEIN, S., 1978. Influence of diet on the distribution of carbon isotopes in animals. Geochim. Cosmochim. Acta, 42, 495-506.

isotopes in the earth sciences || isotopes in palaeoclimatology

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