OXYGEN ISOTOPES IN FORAMINIFERA: OVERVIEW AND HISTORICAL REVIEW
PAUL N. PEARSON
School of Earth and Ocean Sciences, Main Building, Cardiff University, Park Place, Cardiff, CF10 3AT, United Kingdom [email protected]
ABSTRACT.—Foraminiferal tests are a common component of many marine sediments. The oxygen iso- tope ratio (δ18O) of test calcite is frequently used to reconstruct aspects of their life environment. The δ18O depends mainly on the isotope ratio of the water it is precipitated from, the temperature of calcifica- tion, and, to a lesser extent, the carbonate ion concentration. Foraminifera and other organisms can poten- tially preserve their original isotope ratio for many millions of years, although diagenetic processes can alter the ratios. Work on oxygen isotope ratios of foraminifera was instrumental in the discovery of the orbital theory of the ice ages and continues to be widely used in the study of rapid climate change. Com- pilations of deep sea benthic foraminifer oxygen isotopes have revealed the long history of global climate change over the past 100 million years. Planktonic foraminifer oxygen isotopes are used to investigate the history of past sea surface temperatures, revealing the extent of past 'greenhouse' warming and global sea surface temperatures.
INTRODUCTION Note: A number of textbooks provide succinct accounts of oxygen isotope systematics and THE MEASUREMENT of oxygen isotope ratios measurement; this section is based mainly on of biogenic calcite is one of the longest- Faure and Mensing (2005), Allègre (2008), and established and most widely used of all paleocli- Hoefs (2009). mate proxies. It principally provides information Oxygen has three stable isotopes, 16O, 17O, on the temperature or oxygen isotope ratio of and 18O, which occur on Earth in the approximate seawater at the time of calcification if the other proportions 99.757%, 0.038%, and 0.205%, re- parameter is known or assumed. The signal can, spectively (Rosman and Taylor, 1998; other in principle, survive for hundreds of millions of sources give slightly different figures). These dif- years in fossils. Although many types of organ- ferent abundances reflect the fact that the three isms produce calcite skeletons, foraminifera have isotopes are produced by different synthetic path- been employed particularly widely because of ways in stars. The proportions vary somewhat in their abundance and diversity in marine sediment, natural Earth materials because each substance especially deep-sea oozes where many of the has its own prior history of fractionation (proc- longest and most continuous paleoclimate records esses that sorted or partitioned the isotopes) and are found. Here, the development of the proxy in mixing (processes that combined or assimilated both benthic and planktonic foraminifera is re- the isotopes). Fractionations occur in two main viewed in two parts. Part 1 is an overview of the ways, isotope-exchange reactions and kinetic ef- principles of the technique and its early develop- fects. Because molecules with a heavy isotope ment, together with some of its complications and have slightly greater covalent bond strengths and limitations. Part 2 outlines some of the major ap- lower vibrational frequencies than their lighter plications in paleoclimate studies from the 1970s counterparts, they are slightly less reactive. They to the present. are also slower to diffuse along concentration gradients and across membranes. Examples of PART 1: PRINCIPLES AND HISTORY fractionation processes that affect oxygen isotopes in water are evaporation, in which the light iso- Oxygen isotopes tope 16O is slightly preferred, and condensation, In Reconstructing Earth’s Deep-Time Climate—The State of the Art in 2012, Paleontological Society Short Course, November 3, 2012. The Paleontological Society Papers, Volume 18, Linda C. Ivany and Brian T. Huber (eds.), pp. 1–38. Copyright © 2012 The Paleontological Society. THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18 where the heavy isotope 18O is slightly preferred. torical reason for the adoption of this is described The physical mixing of two water masses with below). The original standards were simply different isotopic ratios (e.g. when a river flows SMOW and PDB, but because original supplies of into the sea) will produce water with an interme- both standards ran out and some inter-laboratory diate ratio. differences emerged in defining what the precise Because most natural fractionation and mix- values were relative to other available reference ing processes are strictly mass dependent, the iso- materials, a worldwide convention was deter- topes 17O and 18O fractionate and combine rela- mined (at a meeting in Vienna) (Gonfiantini, tive to 16O proportionally according to their re- 1984; see Coplen, 1994, for further discussion). spective masses (18O fractionating twice as much Deviations from VSMOW tend to be used in stud- as 17O), so for most applications, there is little to ies of the hydrological cycle and also for high- be gained by measuring both isotopes. The normal temperature processes, such as in metamorphic procedure is to measure the ratio of 18O to 16O in a rocks, and studies of phosphates and silica. Devia- sample. By convention, the isotope ratio, R, is tions from VPDB tend to be used in studies of defined as the abundance of the heavier isotope low-temperature carbonates, including foramini- over the abundance of the lighter isotope. For the fera. Note that carbon isotope ratios in carbonates, global average proportions given above, this is 13C/12C, are also quoted relative to the same expressed as follows: VPBD standard (as δ13C ‰). 18O 0.205 The oxygen isotope ratios of carbonate sam- R = = = 0.002055 ples are usually measured using a gas source mass 16O 99.757 spectrometer. The sample is reacted with 100% phosphoric acid, producing CO2 gas that is ion- Because natural fractionations are, in practice, ized in a vacuum chamber by electron bombard- quite small, and 18O is always very much rarer ment. The mass spectrometer accelerates the ions than 16O, the isotopic ratios of natural materials under high voltage, and a magnetic field splits are generally quite close to this average value. them into streams of different isotope ratio that Small differences between small numbers are un- generate electrical currents in the detectors. The wieldy, so (as is the convention for other stable ratio of these currents is proportional to the iso- isotope systems) oxygen isotope ratios are gener- tope ratio of the sample. By alternately switching ally quoted as deviations (delta values) from the between a standard of known isotope ratio (gener- 18 oxygen isotope ratio (δ O) of a standard sub- ally a CO2 reference gas supplied by the National stance in parts per thousand ('per mil', sometimes Institute of Standards and Technology) and an ‘per mille’, signified by the symbol, ‘‰’). unknown sample, the isotope ratio of the un- known can be calculated. In this way, the follow- ing masses are usually measured: 12C16O16O Rsample - Rstandard (mass 44), 13C16O16O (mass 45), and 12C18O16O δ18O = x 1000 (mass 46) (higher masses are created from other Rstandard combinations of 13C and 18O, but these are rare and not routinely measured). From these ratios, the δ18O and δ13C of the sample are calculated There are also good, practical reasons for this simultaneously and quoted relative to VPDB. In convention because isotope ratios of sample mate- reality, some complications have to be taken into rials are almost always measured relative to a account, notably subtraction of the contribution of laboratory standard rather than as an absolute ra- molecules containing 17O to the above masses, tio, which is much more difficult to determine and the effect of isotopic fractionation associated accurately. Note that if a sample has a positive with the phosphoric acid reaction. δ18O, it is said to be enriched in the heavy isotope relative to the standard and, if negative, it is said to be depleted. Foraminiferal calcite It so happens that two different standard val- Foraminifera are single-celled eukaryotic or- ues are in widespread use from which these delta- ganisms belonging to the Phylum Granuloreticu- values are quoted: Vienna Standard Mean Ocean losa, which comprises amoeboid organisms char- Water (VSMOW) and Vienna Pee Dee Belemnite acterized by pseudopodia that have a granular (VPDB, which is a carbonate standard; the his- texture to the flowing cytoplasm (Lee et al.,
2 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA
FIGURE 1.—Foram art: calcite tests of selected benthic (left) and planktonic (right) foraminifera. These are from exceptionally well-preserved Paleogene sediments of Tanzania (33–45 Ma). Scale is approximate; diameters are from about 0.20–0.75 mm. Images: P. N. Pearson and I. K. McMillan. Note the literature is split between those who use the adjective 'planktonic' versus 'planktic' and between those who use 'benthonic' versus 'benthic'. It so happens that planktonic and benthic are clearly in the ascendancy as of 2012, by roughly 10:1 and 30:1 respectively, as shown by a word search on abstracts. It has been argued that planktic is the correct Greek form of the adjective (Rodhe, 1974; Emiliani, 1991) but it has been pointed out that planktonic, like electronic, is perfectly good English, however ugly it may be in Greek (Hutchinson, 1974). The majority usage is followed here.
2000). Most foraminifera are marine, and many cessive chambers added episodically throughout secrete a test (or shell) made of calcium carbonate life, starting with a first-formed chamber (prolo- (CaCO3; generally low-Mg calcite, but high-Mg culus). The chambers all have small openings or calcite in porcelaneous species and aragonite in foramina (singular: foramen) that provide internal some groups). Some foraminifera live among the connections between them and allow cytoplasm ocean plankton distributed in the upper part of the flow inside and outside the test (hence 'foramini- water column; others are benthic, living directly fera', which means 'bearers of foramina'). Gener- on the sea floor or at shallow depths in the sedi- ally, the chambers increase in size through the life ment (Figure 1). Foraminifera are heterotrophs, cycle (ontogeny). Different groups of foraminifera feeding on organic matter rather than photosyn- have adopted a wide variety of chamber shapes thesizing, although many species live in symbiotic and geometries of chamber addition, with spiral association with photosynthetic algae. Foraminif- arrangements being the most common. This eral tests can occur in large numbers, and in many means that many types of foraminifera can be places, they form a significant component of the readily distinguished by eye (Figure 1). Foramini- sea-floor sediment (for quantitative estimates, see fer tests are large enough (mostly in the range Schiebel, 2002). The oxygen in foraminiferal cal- 0.1–1.0 mm) that they can be manually separated cite derives from the seawater (or, in the case of (picked) under a binocular microscope with a fine infaunal species, the pore water) in which the or- brush or needle, and their isotopic ratios can be ganism lived. Hence, the isotope ratios can pro- determined either individually or by combining vide information about the composition and his- relatively small numbers of tests of the same spe- tory of that water, and the environmental condi- cies. This is a convenient advantage in paleocli- tions in which the test was secreted. mate research over smaller biogenic particles such Most foraminiferal tests are built from suc- as coccoliths, which are also very common in ma-
3 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18 rine carbonates, but cannot easily be separated gen isotope ratio of a calcite shell to determine its into species, and are too small to analyze indi- temperature of calcification dates back to the vidually. work of Harold Urey and colleagues (Urey, 1947, Foraminiferal life cycles generally last weeks 1948; the latter paper reports results of collabora- or months, although there is evidence that some of tive work between himself, Charles McKinney, very large (> 1 cm) benthic species of the past John McCrea, and Samuel Epstein at the Univer- may have lived for several years (Purton and sity of Chicago; see also Epstein, 1997, for a per- Brasier, 1999). The complex process of test secre- sonal account of the period). Urey (1947) began tion (biomineralization) has been studied for his investigation by making calculations of the many living species, both benthic (Erez, 2003) isotope exchange reaction between water and car- and planktonic (Hemleben et al., 1989). The de- bonate ions. As he himself explained: “If calcium tails vary across the biological subgroups, as do carbonate is crystallized slowly in the presence of the distinctive test microstructures that result. water at 0°C, the calculations show that the ratio Chamber formation is always episodic with a pe- of the oxygen isotopes in the calcium carbonate riod of rest and feeding in between. The solid should be 1.026 to 500 if the ratio of the isotopes chambers are constructed in a matter of hours in the water is 1 to 500, i.e. oxygen 18 is very from many millions of minute calcite plaques, or slightly concentrated in the calcium carbonate in microgranules, which themselves are formed in- relation to the water. On the other hand, if the tracellularly (Hemleben et al., 1989). temperature is 25°C, the oxygen isotopes will be In some species, seawater is vacuolated and concentrated only to the extent of 1.022 as com- transported inside the cell (Bentov et al., 2009). pared with 1 in 500 on water. This shows that The vacuole in which the calcite is precipitated is there is a slight temperature coefficient for the separated from both seawater and cytoplasm by abundance of 18O isotope in the calcium carbon- membranes (Erez, 2003). The chamber itself is ate.” (Urey, 1948, p. 491; note that these figures formed on an organic template, and an organic would now be revised somewhat.) membrane is found within the test wall of most Urey (1948) pointed out that these differ- species. The temperature within the cell is set by ences, although small, were measurable and con- ambient conditions, but the observed rates of cal- sistent with the actual isotopic ratios of a variety cification are at least an order of magnitude of biogenic carbonates that he and his colleagues slower than inorganic precipitation (Erez, 2003). measured to investigate the effect. They did this Because the formation of calcite is so clearly bio- on a pioneering mass spectrometer built to the logically mediated, there is considerable potential design of Alfred Nier (see Nier, 1947, and for biological fractionation of the various isotopes McKinney et al., 1950, for early analytical devel- and trace metals incorporated into the calcite (see opments). Among the carbonates measured were Zeebe et al., 2008, for a review). Despite this, two samples of Globigerina ooze from the Pacific foraminifera have been found to secrete calcium Ocean at 60°S. Globigerina is a genus of plank- carbonate close to isotopic equilibrium, unlike tonic foraminifer, hence, these were the first of certain other types of calcifiers such as corals, the very many foraminifera to be dissolved and echinoderms, and arthropods (Wefer and Berger, analyzed in the pursuit of temperature informa- 1991). However, small but significant disequilib- tion. The δ18O values of these samples were re- rium effects (sometimes called vital effects) have ported as +2.14 and +1.85 relative to a mollusk been reported and are discussed further below. that grew at 13°C (note the PDB standard had yet to be defined). These enriched isotope ratios im- The heroic age of oxygen isotopes plied a cooler calcification temperature for the As mentioned above, a molecule with a heavy foraminifera than the mollusk, which was consis- isotope has a slightly greater covalent bond tent with the high-latitude location. “I suddenly strength and a lower vibrational frequency than its found myself with a geologic thermometer in my lighter counterpart so it is slightly less reactive. hands,” Urey is reported to have said (Emiliani, This difference in reactivity is influenced by the 1958, p. 2). Although he also admitted it was "not ambient temperature such that an increase in tem- useful" (Epstein, 1997, p. 8), it was evidently an perature lessens the difference in reaction rates interesting problem and, in the immediate post- (essentially, this is because everything vibrates war period, a welcome non-military application of faster at higher temperature). The idea that this isotope physics. Urey (1948) reported measure- effect could be exploited by measuring the oxy- ments of a variety of samples from the Cretaceous
4 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA chalk of England to underline his point, estimat- laboratory standard that is still in use today as ing temperatures of 17 to 27°C. However, he also VPDB (Craig, 1953, 1965). identified several caveats that remain centrally The application of oxygen isotope techniques important: that the isotope ratio of the water in to foraminifera was developed by another of Har- which the calcite grew could have changed old Urey's research associates, the micropaleon- through time, which would change the baseline tologist Cesare Emiliani. Emiliani first turned his for the calculation; that biological disequilibrium attention to planktonic foraminifera, manually effects might be expected (later called 'vital ef- separating various species for analysis. Using a fects' by Urey et al., 1951); and that the preserva- sample of sea-floor sediment from the Caribbean, tion of many fossils may not be good enough for a he showed that different species from the same reliable measurement to be made. sample had significantly different oxygen isotope The theory of mass-dependent fractionation of ratios, which he interpreted as representing differ- the isotopes applies well to gases, but only ap- ent depth habitats (Emiliani, 1954a). At that time, proximately to liquids and not at all to ionic crys- relatively little was known of planktonic forami- tals (Faure and Mensing, 2005). Hence, it is nec- niferal depth ecology, but Emiliani's insight essary to conduct experiments to empirically de- proved correct, as has been demonstrated subse- termine the actual temperature relationship by quently in innumerable plankton tows and nets growing calcite at different temperatures. This job (summarized by Hemleben et al., 1989; see also was given to Urey's graduate student, John Mulitza et al., 1997, 1999). McCrea, who published the first oxygen-isotope Emiliani then turned his attention to benthic paleotemperature equation based on inorganically foraminifera with the aim of determining the past grown calcite crystals, and showed that it was temperature of the sea floor, which had previously close to the theoretical expectation (McCrea, been regarded as a constant environment. He 1950). A revised version of this temperature scale compared the δ18O paleotemperatures of calcare- based on analyses of mollusk shells grown be- ous benthic foraminifera from Oligocene, Mio- tween 7 and 30°C was published by Epstein et al. cene, and Pliocene samples from the Pacific (1953). The historic Epstein et al. (1953) paleo- Ocean and estimated that deep-sea temperatures temperature equation is: had declined from 10.4 to 7.0 to 2.2°C across these epochs (Figure 2). This, he noted, was con- 18 18 18 T(°C) = 16.5− 4.3(δ Occ − δ Osw) + 0.14(δ Occ− sistent with abundant geological and paleobotani- 18 2 δ Osw) cal evidence for global cooling during the Ceno- zoic (Emiliani, 1954b; see also Emiliani, 1961). 18 where δ Occ is the measured value in calcium This discovery has been hailed as the founda- 18 carbonate and δ Osw is the isotope ratio of the tion point of the science of paleoceanography water from which it is precipitated. The slope of this relationship means that a 0.23‰ increase in 18 δ Occ corresponds to a difference of about 1°C. C)
There are many subsequent variants of this equa- o tion, but they mostly take the same general form. Refinements and developments of paleotempera- ture equations are discussed in the following sec- tion. Further Mesozoic analyses were reported by Urey et al. (1951) in a celebrated paper that in- cluded the presentation of a series of measure- TEMPERATURE ( ments through the rostrum of a Jurassic belemnite from Scotland, from which it was argued that the AGE (106 YEARS) organism lived through four seasonal cycles. The presentation of these results at the Geological So- FIGURE 2.—The birth of paleoceanography? Tempera- ciety of America annual meeting in 1950 report- ture estimates of Pacific deep waters in the Cenozoic edly caused a sensation (Allègre, 2008). Other from four oxygen isotope analyses of benthic fora- specimens analyzed were belemnites from the minifera (modified from Emiliani, 1954b). Note the Maastrichtian Pee Dee Formation of Carolina, numerical timescale has been modified considerably which subsequently became the basis for the PDB since initial publication.
5 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
Stages 1 2 3 4 5 6 7 8 9 10 11 12 13
30 m)
20
Weight % 10
fraction (>74 0
C) 30 o
25
Isotopic 20 temperature ( 15 0 50 100 150 200 250 300 350 400 450 500 550 600 650 700 Core depth (cm) FIGURE 3.—First view of the glacial cycles from oxygen isotopes: coarse sediment fraction (top panel) and plank- tonic foraminiferal oxygen-isotope paleotemperature estimates (bottom panel) from Piston Core A-1794 (Caribbean, Lamont Geological Observatory). Isotope 'stages' are labeled at the top. Temperature offsets between the planktonic foraminifer species are mainly related to depth habitat (and hence temperature) of calcification. This record is now considered to span ~400 kyr. (compare with Figure 12). Small letters represent different species. Modified from Emiliani, 1955.
(Hay and Zakevich, 1999). However, just as im- is itself depleted after evaporation and condensa- portant was his subsequent work on piston cores tion fractionations on its journey from the ocean of marine sediments spanning the late Pleistocene to the poles. Hence the growth of ice sheets would to Recent. Analyzing planktonic foraminifera leave the ocean isotopically enriched (and, simul- from these cores, Emiliani (1955) demonstrated taneously, lower sea level). Conversely, melting the existence of a series of glacial cycles, which of the ice would recirculate the isotopically de- he numbered as ‘isotope stages’ back through pleted fresh water. Subtracting this effect, Emili- time by taking successive isotopic maxima and ani (1955) estimated that sea-surface temperatures minima (Figure 3), a notation system that is still had varied by about 6°C between glacial minima in widespread use. These isotopic cycles corre- and interglacial maxima (Figure 3). sponded to micropaleontological assemblage Subsequent measurements of polar ice proved variations in the same cores that were interpreted to be more depleted than Emiliani had originally as representing surface ocean-temperature thought, so his calculations are now thought to changes (Ericson and Wollin, 1956). Emiliani have underestimated the ice volume effect. Alter- (1955) argued that his record supported the orbital native calculations by Olausson (1965) indicated theory of the ice ages that had been developed in that the entire amplitude of glacial/interglacial the nineteenth century (Croll, 1875), and elabo- δ18O variability could be explained by the ice ef- rated further in the earlier twentieth century (Mi- fect rather than temperature. This controversy ap- lankovitch, 1941). proached resolution when Shackleton (1967) ana- Emiliani (1955) knew that the waxing and lyzed benthic foraminifer δ18O data from one of waning of the great ice sheets would have had an Emiliani's cores and showed they had a similar 18 affect on whole-ocean δ Osw, thereby changing amplitude of variation to the planktonic data. The the baseline for his paleotemperature calculations. covariation of benthic and planktonic δ18O is ex- Specifically, ice sheets are expected to be isotopi- plained more easily by whole ocean changes in 18 18 cally depleted with respect to O because they δ Osw than it is by temperature; moreover, it is are formed from high-latitude water vapor, which simply not possible to lower the glacial tempera-
6 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA ture of bottom water by 6°C compared with mod- Global Marine Inc., named the GLOMAR Chal- ern ocean water, because that would put it below lenger (see Winterer, 2000, and National Research freezing (Shackleton, 1967; see also Dansgaard Council, 2011, for historical reviews). and Tauber, 1969). Scientific ocean drilling has continued ever Shackleton’s interpretation has tended to pre- since under continuations of this program: the vail, especially following the CLIMAP (Climate: Ocean Drilling Program (ODP; using the drill- Long-Range Investigation Mapping and Predic- ship JOIDES Resolution) and Integrated Ocean tion) project in the 1970s, which mapped ice age Drilling Program (IODP; using multiple drilling temperature estimates worldwide using microfos- platforms). These programs have completely sil assemblage data (CLIMAP Project Scientists, transformed our understanding of Earth's climate 1976). Hence, the deep-sea oxygen isotope re- history, and oxygen isotope analysis of foraminif- cords of the Pleistocene are now interpreted more eral tests has always been one of the most impor- as a signal of global ice-volume fluctuation rather tant tools in researching Earth's past climate (Na- than temperature, although a component of tem- tional Research Council, 2011). perature change is also present in the records and, in some instances, that component can be several Equations, corrections, offsets degrees (see, e.g., Schrag et al., 2002; Herbert et As described above, McCrea (1950) published al., 2010). Emiliani never fully accepted the the first laboratory oxygen-isotope paleotempera- Shackleton interpretation of the planktonic re- ture equation, which was subsequently revised by cords, arguing that the CLIMAP data overesti- Epstein et al. (1953). Since then, many other em- mated glacial sea-surface temperatures (see Emil- pirically derived equations have been published iani, 1992, and Herbert et al., 2010, for more on using either inorganic calcite precipitates or vari- the controversy). ous sorts of organisms grown in culture or col- Emiliani’s great discoveries in the 1950s were lected in the field across a temperature range. In made possible, in part, by the piston-coring sys- general, the paleotemperature equations take the tem developed by Borje Kullenberg and famously quadratic form: deployed aboard the motor-schooner Albatross in 18 18 18 18 2 the worldwide Swedish Deep Sea Expedition T(°C) = a+b(δ Occ − δ Osw) + c(δ Occ−δ Osw) (1947–8). It was then used by the Lamont Geo- 18 logical Observatory in the Caribbean (see Olaus- where δ Occ is the measured value in calcite and 18 son, 1996, for review). These cores could be up to δ Osw is the isotope ratio of the water from which about 20 m long. The potential of even deeper the calcite is precipitated (standardized to the drilling was proven by Project Mohole, which, in VPDB scale). The coefficients a, b, and c are de- 1961, drilled 183 m of sediment and basalt from a termined experimentally. The term a represents 18 18 single hole in the Guadeloupe Basin (as reviewed the temperature when δ Occ = δ Osw, b is the by National Research Council, 2011). The follow- slope (always negative), and c is the second-order ing year, Emiliani proposed to develop long pa- term for the (always slight) curvature of the slope leoclimate records using an innovative drilling arising from the fact that isotopic fractionation platform, the M/V Submarex. This vessel had been decreases with increasing temperature. Some of a wartime patrol boat, but was later converted by the published equations are simplified by the fact an oil-industry consortium into an experimental that c is set at zero, and therefore, the relationship drill ship—allegedly the first of its kind—with a between T and δ18O is linear. The reason for this derrick mounted on the port side. Its name alludes is that the curvature is so slight, a straight line to the fact that it was a sub-marine explorer, not provides just as good a statistical fit over the de- an ex-submarine chaser! Emiliani's project was sired temperature range. (Note that even the stan- called LOCO (for Long Cores). In 1963, the Sub- dard quadratic is only an approximation to the marex successfully recovered 55 m of pelagic theoretically expected relationship, which is loga- sediment from a single hole. Emiliani's LOCO rithmic.) advisory group developed into the JOIDES (Joint A tricky aspect of the paleotemperature equa- 18 Oceanographic Institutes for Deep Earth Sam- tions is that δ Osw values are always measured on pling) committee that supervised further deep the VSMOW scale (or, before that was defined, drilling off of Florida. This soon led to the forma- the SMOW scale or something equivalent), but to tion of the Deep Sea Drilling Project (DSDP) in be used in equations that are expressed in the 1967, with a drill ship specially designed by form given above (which is usual in paleoclimate
7 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18 studies) they must be standardized to the VPDB (or PDB) scale. This means a small but significant correction or conversion factor must be applied to 18 the δ Osw value. This conversion simultaneously corrects for the following: 1) that the absolute scales are offset from each other; and 2) that there is an experimental difference in the fractionation of the oxygen isotopes that occurs when reacting carbonate with phosphoric acid at 25°C (as is done with calcite samples) versus that caused in the equilibrium between water and CO2 at 25.3°C (as is done with water samples) (see Grossman, 2012, for more details). If this was not compli- cated enough, it is important to appreciate that the accepted value for the (V)SMOW to (V)PDB conversion has changed through time. The origi- nal value taken by Epstein et al. (1953) was thought to be -0.20‰. This was subsequently re- vised to -0.22‰ (Friedmann and O'Neil, 1977) and -0.27‰ (Hut, 1987). When using old equa- FIGURE 4.—The carbonate chemistry effect unmasked. tions, it is generally necessary to use the value Experimental calibrations of the planktonic foramini- that was assumed in the original study in order to fer Orbulina universa in high light (with symbiont be faithful to the original calibration (see Bemis et activity) and low light (without symbiont activity). al., 1998, for review). However, as Grossman Numbers in brackets represent numbers of specimens (2012) points out, that rule does not apply to the measured. The offset in the calibrations is attributed to the non-ambient carbonate chemistry of the microen- original Epstein et al. (1953) equation because it vironment of calcification (see text). Modified from was directly standardized with PDB-derived CO2, Spero (1988) after Bemis et al. (1988). Note that the so in that case, the currently accepted value of δ18O scale is not directly comparable with Figure 5 -0.27‰ is appropriate. because the VSMOW-VPDB conversion is not incor- A problem that gradually became apparent porated on this figure. from field studies is that the paleotemperature equations were repeatedly found to give tempera- ferent equations for a symbiotic species (Orbulina tures that are slightly too high by ~1–2°C com- universa) grown in different irradiance levels (and pared with in-situ temperature measurements hence levels of symbiotic activity) (Spero, 1992; (Shackleton et al., 1973; Kahn, 1979; Williams et Bemis et al., 1998; Bijma et al., 1999). Interest- al., 1981; Deuser and Ross, 1981; Sautter and ingly, the temperature:δ18O relationship for O. Thunell, 1991; Mulitza et al., 2003; see discussion universa grown in low light is very close to that in Bemis et al., 1998). Hence, it seems that some obtained by Kim and O'Neil (1997) for inorganic significant (albeit fairly minor overall) vital-effect calcite (see also Spero and Lea, 1993), but there is fractionation may occur in some foraminifera a clear and consistent offset of about -0.3‰ for (Shackleton et al., 1973). An experimental break- high light conditions (Figure 4). The offset was through was made by Spero (1992), Spero and even greater for O. universa grown in artificially Lea (1993), and Spero et al. (1997), who found 2- 2- 2- high [CO3 ]. More recently, similar [CO3 ] ef- that the carbonate ion concentration [CO3 ] and fects have been suggested for benthic foramini- pH local to the planktonic foraminifer test was fera (Rathmann and Kuhnert, 2008; Rollion-Bard associated with a significant, additional, negative et al., 2008), and experimentally measured in a fractionation of the oxygen isotopes. These pa- coccolithophore and a calcareous dinoflagellate rameters might be affected by various biological (Ziveri et al., 2012). However, the slopes of some processes local to the foraminifera. The most no- of the experimentally determined relationships table of these is the presence of a cloud of photo- 18 2- between δ O and [CO3 ] differ markedly. synthetic symbionts in some species that would be 2- A simple explanation for the underlying proc- expected to lower the pH and increase the [CO3 ] 2- ess responsible for the [CO3 ] effect was pro- locally (Rink et al., 1998). In an elegant series of posed by Zeebe (1999), who related it to isotopic experiments, efforts were made to determine dif-
8 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA fractionation between the species of dissolved bly because of biological evolution (Vincent et al., - 2- inorganic carbon (H2CO3, HCO3 and CO3 ) as a 1985; Katz et al., 2003). function of pH (Zeebe, 1999). Zeebe (2001) cal- Part of the reason for the offsets between spe- culated a theoretical value of -1.42‰ in δ18O for cies of benthic foraminifera may be that some are every 1 unit increase in pH, although experimen- infaunal and some epifaunal, so they calcify in tal determinations of the relationship are still not environments with different pH levels (Bemis et closely constrained (Spero et al., 1997; Beck et al., 1998; Cramer et al., 2011). Also likely is that al., 2005). This explanation indicates that the use while making their test, some foraminifera incor- of the equations such as Kim and O'Neil (1997) porate a component of isotopically light metabolic would yield temperatures too high by ~1.5°C in oxygen that originates from their own respired symbiotic species. It also suggests that for paleo- CO2 (Erez, 1978). It has been suggested that spe- 18 applications, a value for pH as well as δ Osw cies adapted to low oxygen conditions on the sea should be estimated or assumed, especially if floor may have special adaptations enabling oxy- large differences in the carbonate system might be gen exchange with their environment, and so tend expected during the geological past (e.g., Wilson to precipitate their tests closer to equilibrium than et al., 2002; Bice et al., 2006; Zeebe, 2012), or other species (Grossman, 1987). Several studies across some past climatic event (Zeebe et al., have now been published reporting experimental 2008; Uchikawa and Zeebe, 2010). results from deep-sea benthic foraminifera grown Although many shallow-water benthic fora- in culture (Wilson-Finelli et al., 1998; McCorkle minifera have photosymbiotic relationships with et al., 2008; Barras et al., 2009; Fillipson et al., algae, foraminifera that are commonly used in 2010). The temperature relationships of the cul- deep-sea paleoclimate research do not because tured species are encouragingly close to the pub- they live well below the photic zone. One might lished inorganic paleotemperature relationship. expect, then, that their paleotemperature curves However, there appears to be an additional onto- would lie close to the inorganic relationship. This genetic (growth) effect whereby young/small in- is indeed the case for some species, but others dividuals precipitate their tests with a slight nega- were found to have consistent offsets (Duplessey tive offset. This may be because they recycle a et al., 1970; Shackleton and Opdyke, 1973; higher proportion of oxygen from metabolic CO2 Shackleton, 1974; Graham et al., 1981; and many (Barras et al., 2009; Fillipson et al., 2010). subsequent studies). The initial view was that the In general, the oxygen-isotope technique is families Buliminacea (including the common very good at identifying relative temperature deep-sea genus Uvigerina) and Cassidulinacea changes because the slopes of the published rela- give values that are generally close to equilib- tionships are all quite similar, but the possibility rium, whereas the Discorbacea and Rotaliacea of a systematic offset from equilibrium should (including Cibicidoides) usually give variable always be considered in the light of the literature. negative offsets averaging about -1‰ (Grossman, Clearly, a general rule in any oxygen isotope 1987). Hence, for deep-sea compilations, it be- study based on foraminifera that is designed to came standard practice to adjust Cibicidoides val- determine variations down-core is that a single ues by a fixed number to bring them into line with species should be measured from a restricted size Uvigerina data (see review by Wefer and Berger, range. This means the sometimes subtle morpho- 1991). Subsequently, it has been pointed out (Be- logical distinctions between species need to be mis et al., 1998; Lynch-Stieglitz et al., 1999; known and appreciated. Costa et al., 2006) that if one uses the more recent Considering all this, we come to the question: paleotemperature equation of Kim and O’Neil which paleotemperature equation should one use? (1997), Cibicidoides gives values that are closer There is no right answer to this because it depends to equilibrium than does Uvigerina, hence current on the application. For example, does one prefer studies adjust data from other genera to bring an inorganic calibration or one appropriate for the them into line with Cibicidoides, and use a paleo- target organism? Which is best, a laboratory cul- temperature equation calibrated to that genus ture study or one based on in-situ collections and (e.g., Cramer et al., 2009, 2011). Another problem measurements taken in the ocean? Is a new and that must be confronted when using benthic fo- more tightly calibrated equation to be preferred to raminiferal δ18O in deep time is that the offsets a more widely used historical equation, even if the between common species and genera may have differences between them are small? To help in changed over multimillion-year timescales, possi- this decision, some of the most prominent histori-
9 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
TABLE 1.—Comparison of some prominent paleotemperature equations for calcite δ18O. Note that in each 18 case the δ Osw term is quoted relative to VSMOW so it is also necessary to add to it the appropriate VSMOW to VPDB conversion current at the time the equation was calibrated, as given in the last column (see text for details).
Calibra- VSMOW Reference Material tion range Equation to VPDB (oC) conversion
o 18 T( C) = 16.5 - 4.3 (δ Occ- Epstein et al. (1953) Mollusk shell 7.2–29.5 18 18 18 -0.27 δ Osw) + 0.14 (δ Occ-δ Osw)
Synthetic calcite and o 18 O'Neil et al. (1969) reformu- T( C) = 16.9 - 4.38 (δ Occ- calcite-water ex- 0–500 18 18 18 -0.20 lated by Shackleton (1974) δ Osw) + 0.10 (δ Occ-δ Osw) change o 18 T( C) = 17.04 - 4.34 (δ Occ- 18 18 Horibe and Oba (1972) Cultured mollusks 4.5–23.3 δ Osw) + 0.16 (δ Occ- -0.20 18 δ Osw) Cultured o 18 Globigerinoides T( C) = 17.0 - 4.52 (δ Occ- Erez and Luz (1983) 14–30 18 18 18 -0.22 sacculifer (plank- δ Osw) + 0.03 (δ Occ-δ Osw) tonic foraminifera)
Kim and O'Neil (1997) re- o 18 T( C) = 16.1 - 4.64 (δ Occ- formulated by Bemis et al. Synthetic calcite 10–40 18 18 18 -0.27 δ Osw) + 0.09 (δ Occ-δ Osw) (1998)
Cultured Orbulina o 18 universa (high light) T( C) = 14.9 - 4.80 (δ Occ- Bemis et al. (1998) 15–25 18 -0.27 (planktonic fora- δ Osw) mininifera)
Lynch-Stieglitz et al. (1999) In-situ Cibicidoides o 18 T( C) = 16.1 - 4.76 (δ Occ- as arranged by Cramer et al. and Planulina (ben- 4–26 18 -0.27 δ Osw) (2011) thic foramininifera)
Cultured Globoro- o 18 Mielke (2001) in Spero et al. talia menardii T( C) = 14.9 - 5.13 (δ Occ- ? 18 -0.27 (2003) (planktonic fora- δ Osw) minifera)
Cultured Globigeri- o 18 Spero et al. (unpublished) in noides sacculifer T( C) = 12.0 - 5.57 (δ Occ- ? 18 -0.27 Spero et al. (2003) (high light) (plank- δ Osw) tonic foraminifera)
In-situ Globigeri- o 18 noides sacculifer T( C) = 14.91 - 4.35 (δ Occ- Mulitza et al. (2003) 16–31 18 -0.27 (high light) (plank- δ Osw) tonic foraminifera)
10 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA
Erez & Luz (1983)
C)
o Horibe & Oba (1997) (1972)
Lynch-Stieglitz et al. (1999) Epstein et Shackleton al. (1953) (1974) Temperature ( Bemis et al. (1988; HL)
18 18 18 18 Occ - Osw (‰) Occ - Osw (‰)
FIGURE 5.—Seven leading paleotemperature equations compared. Note that the VSMOW to VPDB conversions for 18 the δ Osw term of the respective equations (see text) have been incorporated, making them directly comparable. The equations are shown extending beyond their calibration ranges for illustrative purposes. Note that when the Uvigerina-Cibicdoides offset is applied, the benthic equations of Shackleton (1974) and Lynch-Steiglitz et al. (1999) are brought much closer together than they appear on this plot. They do diverge at the warm end, which makes the choice of equation more critical for Paleogene deep-water temperatures than it is for the colder Neogene. cal and current equations are given in Table 1, and foraminifera in culture (specifically Globigerinoi- a subset is compared graphically in Figure 5. des sacculifer, a widely used species) between 14 The original Epstein et al. (1953) equation and 30°C. More recent culture experiments (Spero was based on mollusks and it was only calibrated et al., 2003) and in-situ field observations (e.g., down to 7°C. Shackleton (1974) argued that the Mulitza at al., 2003) have suggested that the Erez equation of O'Neil et al. (1969), as rearranged by and Luz (1983) equation significantly overesti- him, was more appropriate to use at low tempera- mates temperatures by about 2°C (Mulitza et al., tures typical of deep-sea benthic foraminifera. 2003). One possible reason for this is the carbon- The Shackleton (1974) equation arguably has ate chemistry of the culture experiments: as been superseded by a more recent set of experi- pointed out by Bemis et al. (1998), precise illumi- mental results made on inorganic calcite by Kim nation data and carbonate ion/pH data are not and O'Neil (1997), and converted to quadratic available for them. A variety of species-specific form by Bemis et al (1998). That equation is, equations have been developed using culture ex- however, defined only by measurements made at periments (Table 1). For studies that are based on 10, 25, and 40°C. For those working with benthic data from these particular species, the best choice foraminifera, the equation calculated for core-top is probably to use the most recent appropriate samples down to 4.1°C by Lynch-Stieglitz et al. equation. For studies that use other planktonic (1999) as rearranged by Cramer et al. (2011) cur- foraminifera species, including the many studies rently is the best-constrained field calibration, as that use extinct ones in deep time, a reasonable it is based on a large data set. approach would be to use the inorganic relation- For planktonic foraminifera, the equation of ship of Kim and O'Neil (1997), and acknowledge Erez and Luz (1983) has been used in the litera- that offsets equivalent to about -2oC are likely for ture frequently (note that a minor variant of this the photosymbiotic species if, by analogy with the equation is sometimes quoted in the literature recent, they calcified in a low-pH microenviron- with the coefficients a = 16.998, b = -4.52, and c ment. If precise absolute temperatures are critical = 0.028 based on Table 1 of Erez and Luz, 1983). in a particular study, it would be prudent to use This equation was the first to be calibrated using more than one equation in order to express the
11 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18 uncertainty associated with the choice of equa- Cooke, 1999, for review). Despite this, within tion. certain geographic regions, the relationship be- 18 tween surface δ Osw and salinity is reasonably 18 Regional and temporal variations in δ Osw close, and a set of regional equations have been The most significant complicating factor of proposed (LeGrande and Schmidt, 2006). 18 using δ Occ as a paleotemperature indicator is, of The variability of E-P is exemplified with 18 course, that it depends on knowing δ Osw (Urey, reference to the Hadley circulation of the atmos- 1947). However, this can also be viewed posi- phere. In general, E-P is negative in the equatorial tively because if temperature is known, assumed, regions because maximum solar insolation causes or determined (for example, by using another air masses to rise. Adiabatic expansion of the ris- proxy), the method can be used in reverse to in- ing air masses promotes heavy precipitation. 18 vestigate regional variations in δ Osw, or to track Conversely, the descending limbs of the Hadley the growth and decay of ice sheets through time. cells at ~30°N and S latitude are very dry, causing The δ18O of seawater can be thought of as reflect- local E-P to be positive in these regions. As the ing two principal factors: 1) interregional variabil- trade winds flow back towards the equatorial low ity that exists at any one time; and 2) the mean pressure, they pick up moisture from the ocean 18 value of δ Osw for the oceans as a whole, which surface once again, only to dump it in the Inter- can change over geological time. tropical Convergence Zone (ITCZ). On an even The δ18O of bottom waters depends on the broader scale, the global transport of water vapor source area or areas for that water and its history from the tropics toward the poles (with condensa- of advection: the transport, mixing, upwelling, tion continually favoring the removal of 18O) and downwelling of ocean currents (Rohling and means that the δ18O of precipitation and seawater Cooke, 1999). Consequently, the δ18O of epifau- tends to become progressively more depleted at 18 nal benthic foraminifera can be used (along with high latitudes. The regional δ Osw pattern is also several other geochemical tools) to fingerprint strongly affected by global thermohaline circula- deep water masses (Broecker and Peng, 1982). tion. For example, deep-water formation at the One especially important process that oxygen iso- present time is focused around Antarctica and the topes can help identify is sea-ice formation, which high-latitude North Atlantic. There is a compensa- expels dense brines with strongly depleted δ18O tory net flow of surface water from the Pacific to values. The volume and distribution of such the Atlantic, which evaporates along the way, be- brines influences patterns of deep-water formation coming saltier as it does so. This, in combination and is of potential significance in influencing the with a net vapor transport from the Atlantic to global thermohaline circulation of the oceans Pacific via the trade winds, means that, in general, 18 (Dokken and Jansen, 1999). The influence of the δ Osw of Atlantic surface water is considera- 18 18 brine formation on bottom-water δ O, as re- bly more enriched than the δ Osw of the Pacific flected in benthic foraminiferal data, is especially and Indian oceans at similar latitudes (Broecker, marked in enclosed basins such as the Weddell 1989; Schmidt et al., 1999). Sea in the Southern Hemisphere and the Barents The modern variability of surface ocean 18 and Nordic seas in the Northern Hemisphere (e.g., δ Osw is such that most places are within about Vidal et al., 1998; Mackensen, 2001; Rasmussen 0.5‰ of VSMOW, but the total range in open and Thomsen, 2009). ocean waters is from about -1.5% to +1.5%, with The principal factor governing the regional even greater variation in semi-enclosed basins δ18O of modern surface seawater is the local such as the Mediterranean, Baltic, and Red seas evaporation-precipitation (E-P) balance. Because (Schmidt, 1999; Schmidt et al., 1999; Bigg and the salinity of seawater also depends on E-P, the Rohling, 2000, LeGrande and Schmidt, 2006). two are often correlated, and the variability is Note that a 3‰ range corresponds to a tempera- sometimes referred to as a salinity effect on ture equivalent of over 10°C in the paleotempera- 18 δ Osw. This is clearly only shorthand, however, ture equations, therefore, this is an important issue because the δ18O of precipitation varies system- if one seeks to establish patterns of sea-surface atically with latitude (which changes the slope of temperature distributions in the geological past. the correlation; Craig, 1965), and additional vari- Zachos et al. (1994) proposed a correction factor 18 ability is caused by factors such as river waters for latitudinal variations in δ Osw for use in flowing into the ocean, iceberg melting, the local global compilations of past SST variation by fit- climate regime, and advection (see Rohling and ting a polynomial function through modern
12 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA
18 Southern Hemisphere δ Osw data:
18 2 3 δ Osw= 0.576 + 0.041x − 0.0017x + 0.0000135x where x is absolute latitude (N or S) in the range 0–70° (Figure 6). Note that this equation is quite deliberately a blunt tool: it is based only on data surface waters from the Southern Hemisphere, but is intended for O (‰, VSMOW) of use in both hemispheres and in deep geological 18 time. It avoids the poles because of the unknown influence of ice (or lack of it). The idea is that it can be used in non-analogue conditions, such as Latitude the Paleogene oceans, when different global cir- culation patterns may have operated and a net FIGURE 6.—Correcting for latitude in paleo- 18 flow of surface water to the Atlantic cannot be applications: Modern surface δ Osw (VSMOW) plot- assumed. ted against latitude for the Atlantic and Pacific Oceans The Zachos et al. (1994) latitude correction (data of Broecker, 1989). The dashed line is the Zachos continues to find favor in deep-time paleoclimate et al. (1994) polynomial function for use in paleocli- mate studies, which averages the southern hemisphere studies (e.g., Hollis et al., 2012). It is, however, data (see text). Modified from Crowley and Zachos less appropriate for assessing sites influenced by (2000). boundary currents where the latitudinal relation- ship barely applies, and it was never intended for responsible (e.g. Veizer et al., 1997, 2000; marginal or semi-isolated basins. Moreover, the 18 Prokoph et al., 2008; also see Grossman, 2010, fundamental relationship between δ Osw and lati- 2012). However, there is much less difference tude is very likely to have changed through time, between Mesozoic and Cenozoic values, suggest- especially in the transition from a greenhouse to ing that the effect is either small or negligible for icehouse climate in the early Oligocene (Bice et at least the last ~ 200 m.y.. (Hudson and Ander- al., 2000; Huber et al., 2003). A promising alter- 18 son, 1989; Veizer et al., 1997; Jaffrés et al., 2007). native is to extract surface δ Osw values for study Even so, do not forget that even a 0.5‰ difference sites from General Circulation Models. Such 18 in, say, mid-Cretaceous average δ Osw (which models use imposed paleogeographies and green- could be hard to detect given the current data) house gas forcings to simulate past atmospheric would cause a systematic bias of about two de- and oceanic circulation patterns. In modern grees. isotope-enabled GCMs, the oxygen isotope frac- The calculation of the influence of global ice tionations are also modeled (Schmidt, 1999; Hu- 18 volume on average δ Osw depends on knowing ber et al., 2003; Tindall et al., 2010; Roberts et al., how much continental ice was on the planet at a 2011). However, as pointed out by Roberts et al. given time, and what its mean δ18O was (meteoric (2011) and Hollis et al. (in press), these values are waters also have a slight effect on this budget). strongly dependent on the particular simulation Shackleton and Kennett (1975) made a simple used. 18 (and much quoted) estimate of the isotopic ratio The mean value of δ Osw varies through geo- of an ice-free world (Table 2). Some subtleties in logical time mainly because of changes in the 18 this calculation should be clarified. The δ Osw amount of continental ice, and changes in the rock figure of -0.28‰ is from Craig (1965), and it con- cycle, principally chemical weathering of rocks sists of two components: the conversion factor and the interaction of seafloor basalts and seawa- from SMOW to PDB of -0.20‰ (as then under- ter. The rock-cycle processes are slow, but could stood) and an additional -0.08‰ to account for become significant at the ~1‰ level over times- 8 the difference between the whole ocean (i.e., deep cales of 10 years or more (Veizer et al., 1999; water included) and the SMOW standard. The Wallmann, 2001; Jaffrés et al., 2007). Much evi- 18 value for the ice-volume effect on its own (the dence exists, for example, that δ O values of transition from the modern world to a world with early Paleozoic brachiopods are significantly de- no ice) is almost exactly -1.0‰ by this calculation pleted by several parts per mil compared to com- (see also Zachos et al., 1994). Hence for the pa- parable Mesozoic and Cenozoic fossils, and a 18 leotemperature equation of Shackleton (1974), the very long-term change in δ Osw may be partly
13 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
TABLE 2.—Shackleton and Kennett (1975)’s es- depth habitats and, therefore, across a wide range timate of the isotopic ratio of an ice-free world. * of temperatures. Moreover, at maturity, they typi- Note that the actual calculation yields -1.196, cally descend to a particular depth in the ocean which is rounded to -1.2. where they reproduce. The style of calcification often changes markedly during reproduction (ga-
3 18 metogenesis), when a thick layer of additional Body Volume (km ) δ O‰ calcite is often laid down (called gametogenic calcite; Bé, 1980; Duplessy et al., 1981). If whole (PDB) tests are analyzed, this depth migration will be averaged out (Lohmann, 1995). As the volume of Modern ocean 1, 360,000,000 -0.28 gametogenic calcite can be >50% of the test, tem- Antarctica 24,000,000 -50 peratures can be biased very easily to the cold end-member in species that produce a thick cal- Other ice 2,600,000 -30 cite crust. A recent example of depth stratification of Preglacial 1,386,600,000 -1.2* species is shown in Figure 7 (from Birch et al., ocean submitted). Here, several species of foraminifera from a single core-top sample were analyzed for δ18O, from which a temperature was determined. 18 appropriate value for the δ Osw term (including Estimates of their calcification depths were then the VPDB conversion) is -1.2‰ VSMOW for an made by fitting the temperature estimates against ice-free world, whereas for the Erez and Luz the measured water-column temperature profile. (1984) equation, it would be -1.22‰ VSMOW, Also shown on the figure are vertical profiles of and for Kim and O'Neil (1997) it would be oxygen concentrations and fluorescence (a meas- -1.27‰ VSMOW (see the explanation of appro- ure of chlorophyll concentrations). Bearing in priate VSMOW to VPDB conversions, above). mind that many species of planktonic foramini- The above calculation explains the values for fera occur as juveniles in the surface ocean and an ice-free world that commonly are encountered descend during their life cycle, the foraminifera in the literature (which are sometimes quoted with divide into four main ecological groups: 1) those the VPDB correction included and sometimes living exclusively in the surface mixed layer (0– not). However, a more sophisticated approach has 40 m at this site); 2) those living and calcifying now been taken by modeling the growth of mostly in the upper thermocline where peak pri- present-day ice sheets through their climate his- mary production takes place (40–100 m); 3) those 18 tory and allowing the δ O of precipitation (which calcifying at depth around the shallow oxygen is now ice buried in the ice sheets) to vary with minimum zone (100–200 m); and 4) a single, past climate. Initial calculations indicated that deeper-dwelling species living in sub-thermocline melting of the East Antarctic, West Antarctic, and intermediate water. Groups 1 and 2 are composed Greenland ice sheets would respectively contrib- mainly of carnivores and photosymbiotic species ute -0.91‰, -0.13‰, and -0.07‰ to the ocean, whereas Groups 3 and 4 survive by feeding on totaling -1.11 ± 0.03‰ for an ice-free ocean sinking phytodetritus (e.g., Hemleben et al., (L’Homme et al., 2005). However, it should be 1989). These results are broadly consistent with noted that Cramer et al. (2011) recalculated the many other geochemical and field studies of these total ice-volume effect using the ice-sheet compo- species, including depth-stratified plankton nets sitions of L’Homme et al. (2005) and updated val- and sediment traps (Fairbanks et al., 1982; Curry ues for the mass of the oceans and ice sheets, re- et al., 1983; Hemleben et al., 1989; Mulitza et al., sulting in an estimate of just -0.89‰. At the time 1997, 1999; Faul et al., 2000; Spero et al., 2003; of writing, this discrepancy seems unresolved. In Mohtadi et al., 2011). either case, for use in a paleotemperature equa- Figure 7 shows that most species calcify in tion, the appropriate VSMOW to VPDB conver- subsurface waters. The total reconstructed tem- sion term needs to be added (see Table 1). perature range of this particular assemblage varies from about 29°C to 8°C at depth (Birch et al., Depth habitats of planktonic foraminifera submitted). Hence, if the aim is to reconstruct sea As Emiliani (1954a) discovered, planktonic surface temperature (SST), it is necessary to pick foraminifera are adapted to live across a range of a species that lives its entire life cycle in the
14 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA Depth (m)
FIGURE 7.—Reconstructing the water column:depth habitats of a modern tropical assemblage of planktonic fora- minifera. A core-top sample of 13 species had their mean depths of calcification reconstructed by fitting their δ18O to the locally measured temperature profile (red line). Also shown is the oxygen concentration (related to the bal- ance between photosynthesis and respiration) and fluorescence (related to the chlorophyll concentration). Horizon- tal bars indicate the range of measured paleotemperature for each species; vertical bars indicate the corresponding depth range. Group 1 are species that calcify in the surface mixed layer and approximate the sea surface tempera- ture. Group 2 live at maximum biological production corresponding to the broad peak in oxygen. Group 3 are oxygen-minimum specialists and Group 4 (just one species) calcifies in deep water. mixed layer, the best known of which is Globi- Lipps, 2009; and Aurahs et al., 2009). To what gerinoides ruber. Other species (including the extent these genotypes represent varieties or spe- commonly studied G. sacculifer and Orbulina cies in the biological sense is not yet entirely universa) yield slightly more enriched δ18O values clear, although some genotypes do correspond to even though they are sometimes referred to as morphologically distinct types and almost cer- surface dwellers in the literature. They do occur in tainly do represent separate species with substan- surface waters, but many, if not all of the indi- tial divergence times in the geologic past (e.g., viduals descend to reproduce in the upper thermo- Huber et al., 1997; de Vargas et al., 1999). The cline where gametogenic calcite is added to the presence of unrecognized cryptic species in pa- test. leoceanographic samples could affect the accu- Genetic work has shown that the traditional racy of the δ18O proxy, especially if those species species of planktonic foraminifera as recognized inhabited different environmental conditions by their distinctive morphological features en- (Wang, 2000; Kucera and Darling, 2002). Despite compass a number of genotypes (see recent com- potential problems of distinguishing very closely pilations by Darling and Wade, 2008; Ujiié and related species, the rich fossil record of planktonic
15 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18 O (‰, VPDB) 18
Age (kyr)
FIGURE 8.—Oxygen isotope ratios of individual tests of the planktonic foraminifera Globigerinoides ruber (red circles) and Globigerina bulloides (blue circles; slightly offset to left to aid visibility) in a core that spans the last deglaciation. The single specimen results are plotted against multi-specimen records from the same core by Ivanova (2000) (solid lines). Note the wide range of δ18O values given by individual tests. (Modified from Ganssen et al., 2011.) foraminifera and their use in biostratigraphy als of the same species calcified at different means that it is possible to reconstruct their evolu- depths in the water column or different times of tionary history with impressive detail (see Aze et year. For many applications, particularly time- al., 2011, for a recent review). Individual morpho- series work, this is desirable because it helps species typically persist in the record for several eliminate noise from seasonal and inter-annual millions of years (benthic foraminifera persist for variation, and provides a clearer picture of long- longer—typically tens of millions of years). The term secular changes in the species average value more common species tend to be present continu- that might result from climate or global ice- ously in core samples, making them ideal for the volume variations. However, it also destroys in- construction of geochemical time series. formation of inter-individual variability that might also be of interest, such as determining seasonal Single vs. multispecimen analysis temperature ranges. Unsurprisingly, this is more Early work on the δ18O of foraminiferal tests of an issue with planktonic than it is with benthic required many specimens of the same species to foraminifera, which occupy a much more constant be combined to provide enough calcite for a sin- habitat. gle analysis. Initially, in order to provide about 5 Some mass spectrometers are sufficiently sen- mg of sample, several hundred tests were needed sitive to measure individual planktonic foramini- (Emiliani 1955; Shackleton, 1967; Olausson, fer tests. The first to achieve this were Killingley 1996). Now, instrumental improvements mean et al. (1981), who measured large-sized tests of that typically, just one or a few specimens are Orbulina universa and two other species through needed depending on species and size fraction. a box core in the equatorial Pacific Ocean, span- Combining specimens averages out much of the ning the Holocene and Last Glacial Maximum. variability between individuals that might result, They found inter-specimen variability of about for example, from the fact that different individu- 2‰, which is about twice that which could result
16 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA
still desirable to combine 20+ specimens for an 0.5 analysis, even if the instrument can run just two or three. Schiffelbein and Hills (1984) calculated that to achieve a reproducibility of 0.1‰ at 90% 0.4 confidence one would typically have to combine > 400 tests. Single-specimen results similar to those of 0.3 Killingley et al. (1981) and Ganssen et al. (2011) have been obtained for a number of common spe- cies in various places. Some authors have suc- 0.2 cessfully used inter-specimen variability to isolate environmental signals such as seasonal salinity and/or temperature variation (Tang and Stott, Analytical precision (‰) 0.1 1993; Ganssen et al., 2011; Friedrich et al., 2012a) and changes in the amplitude of the El Niño climate oscillation (Koutavas et al., 2006; 0.0 Leduc et al., 2009). Single-specimen analyses are also useful for distinguishing tests that may have 0 10 20 30 40 50 been reworked across major environmental shifts Foraminifers per sample such as the last deglaciation (Killingley et al., 1981), Cretaceous/Paleogene boundary (Kaiho FIGURE 9.—How many foraminifera should one com- and Lamolda, 1999) and Paleocene/Eocene bine? Examples of precision: reliability sampling boundary (Kelly et al., 1996; Zachos et al., 2007). curves for δ18O based on the variability of single specimen data measurements from one species (Glo- bigerinoides sacculifer, 355–420 µm) at one location. Diagenesis Curves for other species, size fractions, and locations When he first proposed the oxygen isotope will differ, depending on their inherent variability paleothermometer, Urey remarked: “We cannot (standard deviation) of the measurements. Various expect that every fossil which has been or will be confidence levels are shown. The curves show that found on the surface of the earth can be tested by even if fifty tests are combined in one sample, the our methods for the temperature at which it lived. 50% reproducibility level is still less precise than ma- We do hope, however, to find a small fraction of chine precision (dashed line). (Modified and relabeled such fossil remains preserved to such a degree from Schiffelbein and Hills, 1984.) that such measurements can be made” (Urey, 1946, p. 495–496). The initial focus of the Chi- from seasonal temperature change at the site, so cago group was on belemnites because it was they suggested that variability in individual depth thought that the heavily calcified rostra (or habitats, and possibly vital effects, must be re- guards) would be protected from the worst diage- sponsible for the spread of values. They also sug- netic effects. However, as we have seen, when gested that if an accurate species-average δ18O Emiliani began the systematic study of foraminif- value was to be obtained, about 50 individuals eral δ18O, the results were very encouraging: would have to be combined. Similar results to species-specific offsets were found and related to Killingley et al. (1991) were obtained recently by habitat (Emiliani, 1954a), and a coherent signal Ganssen et al. (2011) for two species, and are re- could be extracted going back millions of years produced here in Figure 8. (Emiliani, 1954b). However, just because a signal Schiffelbein and Hills (1984) statistically in- can be extracted, it does not mean that it is pre- vestigated the sampling issue to demonstrate how served without alteration. the number of specimens that are combined for After death, planktonic foraminiferal tests one measurement relates to the precision of a sink through the water column until they either measurement given the standard deviation found reach the sea floor or dissolve in corrosive bottom in a particular population. An example of one of waters (Berger, 1971, 1979a; Hemleben et al., their precision:reliability sampling curves is 1989). It is also possible that some dissolution shown in Figure 9. It emphasizes that noise in occurs in the low pH waters of the oxygen- multi-specimen records is more likely to be be- minimum zone (Milliman et al., 1999; but see cause of sampling than measurement, and that it is also Bissett et al., 2011, and Friedrich et al.,
17 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
FIGURE 10.—Recrystallization textures. These images contrast well-preserved planktonic (a–d) and benthic (i–l) foraminifera under light microscope and SEM dissections versus recrystallized specimens of the same species (e–h, m–p). Diagenesis causes complete internal recrystallization of the test walls. (Modified from Pearson et al., 2007.)
2012a, who suggest water-column effects gener- assessing the mass balance between chamber cal- ally are minimal). Further dissolution in the sedi- cite and gametogenic crust when considering both ment of both benthic and planktonic foraminiferal the original habitat and the potential effects of tests can be promoted by low pH associated with dissolution. the aerobic mineralization of organic matter Precipitation of calcite cements from pore (Jahnke et al., 1997), or, for oxygen-deficient waters is a frequently encountered problem. It can sediments, exposure to oxygen during storage of occur on the outside of a test, overgrowing sur- the cores themselves (Self-Trail and Seefelt, face features, or on the inside, commonly filling 2004). The external gametogenic calcite often is the test completely. If this happens during shallow more resistant to dissolution, so tests can become burial, the cements are likely to have a more en- hollowed out with the internal chamber wall dis- riched δ18O than the original plankton tests, but if solved away (Bé et al., 1975). If different phases cements are formed at high temperatures after of the foraminiferal calcite were produced at dif- substantial burial, the opposite will be the case ferent times in the life cycle, such differential dis- (Schrag, 1999). Diagenetic cements formed from solution potentially could alter the isotopic ratios. percolating meteoric waters (which are often en- This effect can be especially important if an iso- countered in geological sections emplaced on topic record is produced from a site that shows land) also will tend to have depleted δ18O, often varying dissolution intensities through time. by several parts per mil (e.g., Corfield et al., Lohmann (1995) highlighted the importance of 1990). Clearly, foraminifer tests with substantial
18 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA C (‰, VPDB) 3
18O (‰, VPDB)
FIGURE 11.—Comparison of the δ18O and δ13C of a well-preserved middle Eocene assemblage of foraminifera from Tanzania (right) with a recrystallized assemblage of the same age from tropical Pacific drill core ODP Site 865 (cen- tre). Species are divided into mixed layer, thermocline, lower thermocline and unknown habitats. Lines join data from the same species and size fraction. Data of the recrystallized assemblage converge on the expected composi- tion of seafloor diagenetic calcite precipitated at ~11oC with average δ13C. Assuming the recrystallized assemblage started out with similar isotopic ratios to the well preserved foraminifera, the data suggests a ~50–60% contribution of diagenetic O and C in the recrystallized tests (see text). ODP = Ocean Drilling Program; TDP =Tanzania Drilling Project. See Appendix 1 for data and further information. diagenetic overgrowth and infilling are not suit- so claims in the literature that foraminifera are able for extracting accurate paleoenvironmental well preserved cannot always be relied upon. information. Fortunately, such effects are usually Pearson and Burgess (2009) proposed the follow- plainly visible under the light microscopic and/or ing criteria for recognizing exceptionally pre- scanning electron microscope (SEM) (Pearson served tests: (1) tests should be reflective and and Burgess, 2009). translucent under reflective light, especially the Micron-scale recrystallization of tests, which smaller, thin-walled species; (2) when placed in occurs commonly during the shallow burial of water or oil, ultrafine features smaller than a mi- deep-sea carbonate oozes, is more of a problem. cron should survive (e.g. spines, if originally pre- The biogenic texture of most foraminifer tests is sent); (3) originally smooth parts of the test sur- microgranular, with individual granules or face and interior should still be smooth on a sub- plaques in the ~0.1 µm size range and having ir- micron scale; and (4) the original submicron regular shapes. However, this texture is com- granular texture should be visible in cross-section monly replaced entirely by larger, more-equant when the test is broken. neomorphic calcite crystals in the ~1 µm range These features are illustrated in Figure 10. It (Pearson et al., 2001). This wholesale recrystalli- was originally hoped that benthic foraminifera zation can cause very substantial alteration of the might be less prone to recrystallization than oxygen-isotope ratios (Pearson et al., 2001, 2007; planktonic forms because of their denser calcite Williams et al., 2004; Sexton et al., 2006). Per- microstructures (Pearson et al., 2001), but it haps because a component of the original isotopic seems that this is not the case (Sexton et al., 2006; signature remains after recrystallization, this Pearson et al., 2007; Figure 10). Unfortunately, problem was not fully appreciated for many years, recrystallization is nearly ubiquitous in deep-sea
19 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18 carbonate oozes and chalks of substantial geo- In this section, some of the main applications of logic age (millions of years). On the other hand, oxygen isotope analyses of foraminifera are re- well-preserved foraminifera generally are found viewed under three headings: 1) climate cycles in hemipelagic clays, where the low permeability and cyclostratigraphy, 2) deep-time benthic fo- presumably inhibits the recrystallization process raminiferal compilations, and 3) planktonic fo- (e.g., Norris and Wilson, 1998; Pearson et al., raminiferal habitats and sea-surface temperatures. 2001; Wilson et al., 2002; Sexton et al., 2006; Happily, these three subject areas correspond to Pearson and Wade, 2009). Emiliani's troika of papers that helped define the The effect of diagenetic recrystallization on discipline of paleoceanography in its earliest days the oxygen and carbon isotope ratios of plank- (Emiliani, 1954b, 1955, and 1954a, respectively). tonic foraminiferal assemblages is shown in Fig- ure 11, which compares two assemblages of mid- Climate cycles and cyclostratigraphy dle Eocene foraminifera, one well-preserved and In the 1950s and 1960s, ongoing piston coring the other recrystallized. This type of comparison, and Emiliani's deep coring strategy immediately although never exact, suggests that recrystallized paid dividends as isotopic investigations found assemblages are partially homogenized in their the same series of δ18O stages in many different isotope ratios, reducing interspecies differentials, locations and extending further back into time. and converging on the expected value for diage- These cycles were similar to the warm/cold cycles netic calcite. Because the δ13C of diagenetic cal- that had been identified previously using forami- cite is similar to an average value of the foramini- niferal assemblage data including the coiling di- fera, the absolute values do not change much. rections and abundance of certain species (e.g., However, the oxygen isotope ratio of inorganic Ericson et al., 1956). The impact of these records calcite on the sea floor is very different from the was revolutionary, especially when one considers original planktonic values because of the colder the extensive labors of land-based glacial geolo- temperatures, so the shift in δ18O is much more gists over the previous century had managed to substantial. This point was made by Pearson et al. identify only four major cycles of ice advance and (2001, their fig. 4) in a similar diagram; data for retreat (see review by Imbrie and Imbrie, 1979). Figure 11 and further details are given in the Ap- Profound discoveries are not always immediately pendix and online spreadsheet at accepted, and Emiliani’s orbital interpretation of
20 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA *+'"#,$-./0/1#$.0(20-"(213'$4'$5(-66$7/(#
!"#$%&'()
FIGURE 12.—Pacemakers of the ice ages: A stack of 57 globally distributed δ18O records from benthic foraminifera, correlated to a single timescale. Each record has negative (low ice, warm temperature) values to the top. The short cores are piston cores; the longer ones are Deep Sea Drilling Project and Ocean Drilling Program sites. The apparent curvature of the long records is real. (Modified from Liesiecki and Raymo, 2005)
δ18O data back to 5.3 Ma was presented by myr. This intensification is accompanied by the Lisiecki and Raymo (2005), and was based on a development of the characteristic saw-tooth pat- large number of previous studies from laborato- tern of the late Pleistocene glacial cycles, in ries around the world (see Lisiecki and Raymo, which trends towards glacial intensification are 2005, for references). Plotted alongside one an- suddenly and repeatedly terminated in abrupt de- other (Figure 12), the 57 globally distributed sites glaciations. An intriguing recent suggestion is that provide eloquent testimony to the ubiquitous in- these late Pleistocene changes are the early signs fluence of changes in the ice sheets and deep-sea or prelude of a future transition toward permanent temperatures over this period. The data indicate a glaciation of the Northern Hemisphere continental long-term cooling/ice growth trend and an in- masses (Crowley and Hyde, 2008). crease in the amplitude of the cycles towards the The δ18O cycles relate to changes in deep-sea present day. Changes in the dominant frequency temperature and ice volume (and hence sea level) of variations are also clear, especially the intensi- (Cutler et al., 2003), and various authors have fication of ~100 kyr cyclicity over the past ~1.2 attempted to separate these effects (e.g., Shackle-
21 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18 ton, 2000; Raymo et al., 2006). The shape of the Pälike et al., 2006). Other long records from the cycles is very different from the much smoother Pacific and South Atlantic have helped extend the orbital forcing of climate (e.g., calculated solar- coverage, and emphasized the importance of the insolation curves). It is clear that although the long eccentricity cycle (~405 k.y.r) on both δ18O orbital cycles set the frequency of variability, the and δ13C (Wade and Pälike, 2004; Westerhold et climatic response is highly nonlinear. Much more al., 2005). Combined with high-resolution biostra- than orbital theory is required to understand gla- tigraphy (Wade et al., 2011), such studies have cial cycles, which must be strongly influenced by allowed the calibration of the geological timescale feedbacks and thresholds internal to the Earth sys- (based on radiometric decay constants) to orbital tem. As Hays et al. (1976) originally put it, the chronology (based on gravitational physics of the orbital forcing is the ‘pacemaker’ of the ice ages, solar system). The excellent fit between these not the sole explanation for their pattern of onset geological chronometers validates both ap- and retreat, or their severity. proaches to geochronology with great precision. Oxygen-isotope stratigraphy has become an Although chaotic interactions mean that the solu- indispensable tool for stratigraphic correlation and tions for past orbital dynamics (e.g., interactions timescale refinement, especially when combined between precession, obliquity, and eccentricity) in the same down-core records with magnetostra- cannot be 'retrodicted' beyond about 40 Ma (Las- tigraphy (e.g., Shackleton and Opdyke, 1973; kar, 1999; Laskar et al., 2010), the geological re- Bassinot et al., 1994) and biostratigraphy (often cords themselves can help reveal what those in- based on the foraminifera themselves; see Wade et teractions were (Pälike et al., 2004). al., 2011, for review). In the 1960s, a different source of long-term (100 k.y.r.+) climate archives Deep-time benthic compilations was discovered in ice cores, the δ18O variations of When Emiliani (1954b) compared the first which over the last glacial cycle were found to δ18O paleotemperatures of Oligocene, Miocene, closely resemble the saw-tooth pattern already and Pliocene foraminifera (see above, Figure 2) known from deep-sea foraminifer records (Dans- he relied on cores from the Albatross expedition gaard et al., 1969). Ice core records extend back that happened to sample sediment of those ages. almost to a million years (Luthi et al., 2008), re- He remarked: “It is unfortunate that more paleo- vealing how intimately the global ice-volume temperature data from the ocean bottom are not fluctuations first observed in foraminiferal δ18O available, particularly from the Eocene, Upper relate to atmospheric greenhouse-gas concentra- Miocene, and Lower Pliocene. The type of mate- tions. Similar patterns also have been found in the rial that is needed, however, makes rather remote δ18O of speleothems and compared to foraminifer the probability of securing additional, suitable records (e.g., Bar-Matthews at al., 2003). High- samples in the near future” (Emiliani, 1954b, p. resolution records of climate variability from such 855). Systematic ocean drilling would eventually archives has helped illuminate the geographic change that! variability of climate change on land and in the DSDP Legs 6 and 17 produced several long oceans, notably the so-called 'bipolar seesaw' of drill sites from the tropical Pacific Ocean extend- Northern and Southern Hemisphere surface tem- ing as far back as the late Cretaceous. Douglas peratures during the last deglaciation (Broecker, and Savin (1971, 1973) analyzed benthic and 1998; Barker et al., 2009). The causes and conse- planktonic foraminifera at intervals from the Cre- quences of rapid climate change are being studied taceous through Pleistocene, and found a long- intensively, with an increasing emphasis on un- term increase in δ18O that they interpreted as a derstanding spatial variability and climatic tele- cooling trend. They also noted the inherent uncer- connections (Cronin, 1999; Clark et al., 2002; tainty associated with likely changes in the oxy- Clement and Peterson, 2006). gen isotope ratio of seawater, which they assumed This influence of orbital cyclicity on forami- to be constant for the sake of calculation. niferal δ18O and δ13C records from deep-sea cores Shackleton and Kennett (1975) produced a (and other parameters, such as CaCO3 content) somewhat higher resolution record from the high- extends back throughout the available record. For latitude Campbell Plateau south of New Zealand example, drilling on the Ceara Rise, an aseismic and, as seen above (Table 2), took a more sophis- ridge in the tropical Atlantic, has produced more- ticated approach to calculating the oxygen isotope or-less continuously cyclic records back to the ratio of seawater in an ice-free world. Their ben- early Oligocene, ~30 Ma (Shackleton et al., 1999; thic isotope record shows a long-term cooling
22 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA
-1 0
EECO 1 PETM K/P MECO 2
3 O (‰ VPDB) EOT MMCO 18 4
5 INHG
0 10 20 30 40 50 60 70 80 Age (Ma)
FIGURE 13.—The big stack: Compilation of deep-sea benthic δ18O values obtained by laboratories worldwide for the last 80 m. yr., normalized to Cibicidoides (see text). Black: global mean; blue: southern ocean; orange: Pacific; red: Atlantic. INHG = Intensification of Northern Hemisphere Glaciation; MMCO = Middle Miocene Climate Optimum; EOT = Eocene–Oligocene transition; MECO = Middle Eocene Climate Optimum; EECO = Early Eocene Climate Optimum; PETM = Paleocene/Eocene Thermal Maximum; K/P = Cretaceous / Paleogene boundary. Note that the timescale of the Lisiecki and Raymo (2004) stack (Figure 12) corresponds to the last 5 myr. of this record. (Modi- fied from Cramer et al., 2009, with acronyms added.) Trend lines represent different oceans. trend from the late Paleocene to the Plio–Pleisto- Cenozoic benthic foraminiferal isotope data were cene. Shackleton and Kennett (1975) resolved presented by Miller et al. (1987), Zachos et al. several additional features, such as an early Eo- (2001, 2008), and Cramer et al. (2009), and for cene warming and sharp steps towards more posi- the Cretaceous, by Grossman (2010) and Frie- tive values in the early Oligocene and middle drich et al. (2012b). When comparing these com- Miocene. The early Oligocene δ18O shift to more pilations, it is important to appreciate that they enriched values was interpreted as due to cooling, include different correction factors for vital-effect not ice growth, because the bottom-water tem- fractionations of some benthic foraminiferal gen- perature was deemed too warm to be compatible era (see discussion above). The recent Cramer et with an ice cap in the early Oligocene. Instead, al. (2009, 2011) and Friedrich et al. (2012b) com- Shackleton and Kennett (1975) suggested that the pilations also differ from earlier versions by cal- Antarctic ice cap grew between the middle and culating separate trends for different ocean basins late Miocene—a conclusion that has subsequently in order to distinguish the influences of source been challenged, as is discussed further below. waters with different temperatures (an effect that More data were published by Savin et al. (1975), becomes very noticeable after the Eocene–Oligo- who independently reached similar general con- cene transition; Figure 13, and in the Cretaceous, clusions. Data from these early records were for which see Friedrich et al., 2012b). augmented and synthesized by Savin (1977). The deep-sea benthic foraminiferal oxygen As isotope laboratories proliferated and a isotope stack is the best record we have for trends wider community became involved in deep-sea in climate since the Early Cretaceous. It illustrates drilling, data sets expanded, a process that contin- how profoundly the temperature and global ice ues to the present day. Significant compilations of volume must have changed over this period. The
23 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
-2.0 24.5
-1.0 20.2
0.0 13.6 16.0 C) o +1.0 9.7 12.0
+2.0 6.1 8.2 O (‰, PDB) 18
+3.0 2.8 Temperature (
+4.0 0.7 BOTTOM SURFACE
0 10 20 30 40 50 60 70 80 Age (Ma) FIGURE 14.—The first glimpse of deep climate history: oxygen isotope paleotemperature estimates for various planktonic (open symbols) and benthic (filled symbols) foraminifera from DSDP Legs 6 and 17, tropical Pacific Ocean (From Douglas and Savin, 1973, axes relabeled for clarity). Note the different 'bottom' and 'surface' tempera- ture scales. current consensus is that the planet was close to Planktonic habitats and paleo sea-surface being ice-free in the early Eocene, and seafloor temperatures temperatures were about 14°C, compared to only The benthic foraminiferal oxygen-isotope 2–3°C today (Cramer et al., 2011). Global tem- compilation (e.g,. Zachos et al., 2008) is some- peratures may have been even higher in the mid- times interpreted as a global climate curve (e.g., Cretaceous. The long-term Cenozoic climate vari- Hansen et al., 2010), but actually, it mainly is a ability was memorably described by Zachos et al. record of high-latitude climate. Even though the (2001) as consisting of trends, rhythms, and aber- benthic foraminifera span the latitudes, the deep rations—the trends include the warming in the ocean water in which they calcified was mostly Paleocene and coolings in the Eocene, Oligocene, derived from high-latitude source regions where and Miocene to Recent. The aberrations can be the surface water was dense enough to sink. The divided into transient events such as the growth and decay of continental ice sheets also Paleocene/Eocene Thermal Maximum, and occurs mainly at high latitudes. Hence, the ben- switches, such as the Eocene–Oligocene transi- thic record ideally should be matched with sea tion. The rhythms are the orbital cycles which, surface temperature (SST) δ18O records from although ever-present, are too short in wavelength planktonic foraminifera from all latitudes, and to be individually resolved on Figure 13 (but see especially the tropics, to determine a fuller picture Figure 12 for comparison of the last 5 my.). of oceanic temperature variation through time. First, however, it is necessary to determine
24 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA the relative depth stratification of fossil plank- tonic species so that SSTs can be investigated. Most common species have been analyzed for δ18O (and δ13C, which gives additional informa- tion), allowing relative depth rankings to be estab- lished as far back as the mid-Cretaceous (notably Douglas and Savin, 1978; Boersma et al., 1979,
1987; Savin et al., 1985; Shackleton et al., 1985; Age (Ma) Corfield and Cartlidge, 1991; Pearson et al., 1993; Pearson and Shackleton 1995; D’Hondt and Ar- thur, 1995; Huber et al., 1995; Lu and Keller, 1996; Norris and Wilson, 1998; Wilson and Nor- ris, 1998; Abramovich et al., 2002; Wilson et al., 2002; Sexton et al., 2006; Pearson and Wade, 2008). Taking advantage of the excellent fossil Proxy temperature (oC) record of foraminifera, stable-isotope evidence of depth habitats has been used frequently in evolu- FIGURE 15.—The evolution of late Paleocene and Eo- tionary paleobiology, for example, to study spe- cene paleotemperatures as reconstructed from fora- ciation (Norris et al. 1996; Pearson et al., 1997), minifera from deep ocean sites (ODP 1219; ODP 865) habitat evolution (Norris et al., 1993; Hodell and and the exceptionally well-preserved succession in Vayavananda, 1993; Schneider and Kennett, Tanzania with 'TEX-86' paleotemperature estimates for 1996; Coxall et al., 2000, 2007; Ando et al., Tanzania are also shown. Open symbols: benthic fora- 2010), and macroevolution in relation to envi- minifera (not available from Tanzania) and filled sym- ronmental change (Ezard et al. 2011). bols: surface dwelling planktonic foraminifera. The The first long-term records of planktonic fo- difference in tropical SST is larger than can be ex- raminifer δ18O were made during the first phase plained by regional variation and was attributed to the recrystallized state of the open ocean sites, which of deep-sea drilling (Douglas and Savin, 1971, makes planktonic foraminiferal records converge on 1973; Shackleton and Kennett, 1975). Compari- the benthic foraminiferal records. (From Pearson et al., sons between planktonic and benthic records sug- 2007) gested that the δ18O difference between them was relatively low in the early Cenozoic and had in- creased with time (Douglas and Savin, 1973; see A similar picture of low tropical temperatures Figure 14; see also Shackleton and Kennett, 1975, and reduced latitudinal temperature gradients also Savin et al., 1975, Boersma et al., 1979, Berger, began to emerge from Miocene compilations 1979). In these studies, planktonic foraminiferal (Savin et al., 1985; Keller, 1985) and the Creta- paleotemperature estimates for the Paleogene and ceous (Douglas and Savin, 1978; Shackleton, Late Cretaceous were found to be in the range 20– 1984; D’Hondt and Arthur, 1996) (see Grossman, 25°C for tropical sites, which is considerably 2010, for a recent compilation of data). However, cooler than modern tropical SSTs, which peak the pattern seemed at odds with paleontological around 30°C. Berger (1979b) pointed out that the data, which indicated warm tropical temperatures low global temperature gradients implied by these during these climate states (reviewed by Adams et studies indicate that mean wind velocities, ocean al., 1990). When GCM simulations began to be currents, upwelling, and downwelling would all applied to past climates, the low-gradient tem- be very much slower at these times in the past. perature pattern demanded by the Berger/ Shackleton and Boersma (1981) reviewed the Shackleon/Boersma theory of Eocene climate then available data, finding (like Berger, 1979b) proved impossible to reproduce (Barron, 1987, that high-latitude surface temperatures were Bush and Philander, 1997; Bice et al., 2000), warmer than modern (consistent with an absence which remains very much the case today. As Bar- of large ice sheets), but tropical temperatures were ron (1987, p. 729) put it: “The lack of consistency substantially cooler, with maximum values of between model results, oceanic data, and terres- about 23°C. They suggested that more efficient trial data defines the nature of the problem of Eo- transport of heat by ocean currents might have cene climates. The solution must involve reinter- warmed the poles at the expense of the tropics, pretation of at least one major source of informa- creating the ice-free Eocene climate. tion.” D’Hondt and Arthur (1996) called it the
25 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
‘cool tropic paradox’ in relation to the Cretaceous. most different from that of early diagenesis on the In their review of the problem, Crowley and sea floor. The effect would be much reduced at Zachos (2000) pointed out that the lack of evi- high latitudes, hence diagenesis can contribute to dence for warmer than modern tropical SSTs in an apparent flattening of the meridional tempera- the past, if substantiated, would provide an impor- ture gradient as well as a reduction of tropical tant constraint on climate models. temperatures (Schrag, 1999; Pearson et al., 2001). The possibility that diagenetic alteration of Other evidence for warm δ18O temperatures from foraminiferal calcite biases measured paleotem- well-preserved foraminifera supported, in some peratures towards cold estimates has long been cases, by a range of other proxies has been pro- suspected (Killingley, 1983) but was not generally vided by Huber et al. (2002), Bice et al. (2003, thought to be a major effect, partly because inter- 2006), Stewart et al. (2004), Zachos et al. (2006), species isotopic differentials are consistently ob- Burgess et al. (2008), Bornemann et al. (2008), served in the deep-sea record (Crowley and Pearson and Wade (2009), and Wade et al. (2012). Zachos, 2000). However, some evidence for warm Unlike much of the data from recrystallized fora- tropical SSTs did begin to emerge (e.g., Wilson minifera, such results are generally in line with and Opdyke, 1996, based on coral aragonite and GCM simulations with high CO2 forcing (Huber, early diagenetic cements; also Norris and Wilson, 2008; Tindall et al., 2010; Roberts et al., 2011; 1998; Wilson and Norris, 2001, from exception- Huber and Caballero, 2011). The prevailing view ally well-preserved planktonic foraminifera from is that the Cretaceous and Paleogene warm cli- the mid-Cretaceous claystones). Pearson et al. mate states were extreme greenhouse climates (2001) analyzed exceptionally well-preserved (Huber, 2008), and are comparable to simulations Late Cretaceous and Eocene foraminiferal assem- of future climate with anthropogenic greenhouse blages from clay-rich sediments from Tanzania warming. and elsewhere (Figures 1, 10), and argued that seafloor diagenetic micron-scale recrystallization The multiproxy revolution in carbonate oozes had the effect of partially ho- In paleoceanographic studies, we would fre- mogenizing the δ18O and δ13C differences, caus- quently like to know some parameter from the ing substantial 18O enrichment but not destroying past that cannot be measured directly (e.g., tem- interspecies differences altogether (Figure 11). perature, pH, pCO2, pO2, circulation patterns, nu- Pearson et al. (2007) reported data from more trients, levels of biological productivity) so we exceptionally well-preserved foraminiferal as- must find something that can be measured that semblages from a series of drill cores and outcrop would have responded to the target variable. samples from the hemipelagic clays of Tanzania These are paleoclimate proxies, and oxygen iso- (Figure 15), and compared the data with tropical topes of biogenic carbonates are a classic exam- Eocene δ18O records from the Pacific (Bralower ple. They have been joined gradually by a wide et al., 1995; Dutton et al., 2005). Whereas the range of other isotopic and geochemical proxies, deep-sea records suggest that temperatures peaked some of which are measured on foraminifera and around 25°C in the early Eocene and declined others on bulk sediment or organic compounds in thereafter to about 17°C in the late Eocene (simi- the cores. It is increasingly common to find two lar to Douglas and Savin, 1973, see Figure 14; or more proxies used together in an attempt to and many subsequent studies), the Tanzania tem- reduce the degrees of freedom in determining pa- peratures are mostly > 30°C (i.e., warmer than leoenvironmental conditions. A common example modern) and do not show a marked decline. The is the combined use of Mg/Ca ratios and δ18O: the warm temperature reconstructions were supported former as an indicator of temperature and the lat- 18 by the TEX-86 organic paleothermometer (also ter, taking that value into account, of δ Osw. shown in Figure 15). Wilson and Norris (2002) Multiproxy studies involving δ18O have tack- compared glassy and recrystallized foraminifera led many of the foremost problems in paleocean- from the mid-Cretaceous and found that the ographic research from the temperature and salin- glassy specimens yielded much warmer paleo- ity of the glacial ocean (Kucera et al., 2007) to the temperature estimates. The implication of these temperature and CO2 content of the early and studies is that diagenetic micro-recrystallization mid-Cretaceous (McArthur et al., 2007; Bice et can reduce apparent planktonic foraminiferal al., 2006). Every proxy has its limitations, so the temperatures by > 10°C in the tropics, where the use of multiple proxies is essential if important temperature of calcification in the surface ocean is variables, such as sea-surface temperatures, sea-
26 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA level variations, and ice fluctuations are to be de- providing fresh directions for the application of termined, especially for the distant past. Using oxygen isotopes: first, the ability to analyze small proxies together can help highlight when one or sample spots by laser ablation ICP-MS (e.g. Rei- another is compromised. An ideal study would chart et al., 2003) or ion microprobe (Kozdon et use a range of inorganic and organic proxies in al., 2009, 2011); and second, the measurement of combination with model results. The literature is the CO2 isotopologues that can, in principle, pro- too vast to review here, but the approach can be vide simultaneous information about both tem- 18 exemplified by studies of the Eocene-Oligocene perature and δ Osw without the need for assuming transition (EOT), the largest permanent shift in one or the other (Eiler, 2007). These develop- the global climate state of the Cenozoic (high- ments are already affecting the interpretation of lighted on Figure 13). This transition corresponds the more traditional whole-specimen analyses, but to a large ~1.5% shift in deep-sea benthic forami- are reviewed by other papers in this volume so niferal δ18O. When initially discovered will not be discussed further here. (Devereux, 1967; Shackleton and Kennett, 1975; Kennett and Shackleton, 1976; Savin, 1977), the CONCLUSIONS entire isotope shift was attributed to deep-sea cooling. Interpretations of later records invoked a Harold Urey (1893–1981) was awarded the combination of cooling and ice growth (e.g. Arthur L. Day Medal of the Geological Society of Keigwin, 1980; Keigwin and Corliss, 1986; Poore America and the Goldschmidt Medal of the Geo- and Matthews, 1984; Shackleton, 1986; Miller et chemical Society for his geologic thermometer to al., 1987). The discovery of ice-rafted debris and sit alongside his 1934 Nobel Prize for the discov- glaciomarine sediments of Oligocene age (Ham- ery of deuterium. The main caveats identified by 18 brey et al., 1991; Zachos et al., 1992) suggested Urey (1948)—variations in δ Osw, vital-effect that at least some component of ice growth was fractionations, and diagenesis—have all been the likely to be in the signal. Surprisingly, the first subject of intense debate. Application to foramini- combined measurements of δ18O and Mg/Ca ra- fera and deep-sea sediments might have devel- tios in foraminifera failed to detect any deep-sea oped slowly had it not been for the industry and cooling, so the isotope shift was interpreted as a insights of Cesare Emiliani, who accelerated the signal of ice growth alone, the exact opposite of field rapidly in the 1950s, and identified the main the initial interpretation (Lear et al., 2000). This research themes of paleoceanography. The result was problematic, however, because it would oxygen-isotope method has turned out to be just seem to demand growth of an Oligocene ice sheet as important for determining fluctuations in past larger than the modern one (Coxall et al., 2005). global ice volume as it has past temperatures. The discovery that Mg/Ca ratios are affected by Data from foraminifera have provided us with 2- [CO3 ] (Elderfield et al., 2006) led to a reap- invaluable records of global change on astro- praisal of the records (Lear and Rosenthal, 2006). nomical timescales, proving the orbital theory of Subsequent combined proxy studies have tended the ice ages, and, over multimillion-year times- to suggest a roughly equal component of both ice cales, revealing, in detail, how global climate has growth and cooling (Lear et al., 2008; Katz et al., evolved from a greenhouse to an icehouse state. 2008; Peck et al., 2010; Pusz et al., 2011; Bohaty et al., 2012; Wade et al., 2012), although very ACKNOWLEDGMENTS precise estimates are difficult to achieve because of the combined analytical uncertainties and re- Ben Cramer and Dick Norris provided de- producibility (Billups and Schrag, 2002). It is tailed and insightful reviews. The manuscript has possible to accommodate all of this early Oligo- also benefited from the comments of Helen Cox- cene ice on Antarctica, especially if Antarctic ice all, Ethan Grossman, Brian Huber, Howie Spero, was considerably less depleted in 18O than it is and Bridget Wade. now (Deconto et al., 2008) or if there was a larger continental area above sea level than has gener- REFERENCES ally been assumed (Wilson et al., 2012). Oxygen-isotope analysis of foraminifera con- ABRAMOVICH, S., G. KELLER, D. STUBEN, AND tinues to be one of the most important tools in Z. BERNER. 2003. Characterization of late Cam- paleoclimate research. In recent years, two other panian and Maastrichtian planktonic foraminiferal significant developments have occurred that are depth habitats and vital activities based on stable
27 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
isotopes. Palaeogeography, Palaeoclimatology, BECK, W. C., E. L. GROSSMAN, AND J. W. MORSE. Palaeoecology, 202:1–29. 2005. Experimental studies of oxygen isotope ADAMS, C. G., D. E. LEE, AND B. R. ROSEN. 1990. fractionation in the carbonic acid system at 15o, Conflicting isotopic and biotic evidence for tropi- 25o, and 40oC. Geochimica et Cosmochimica Acta, cal sea-surface temperatures during the Tertiary. 69:3493–3503. Palaeogeography, Palaeoclimatology, Palaeoecol- BEMIS, B. E., H. J. SPERO, J. BIJMA, AND D. W. ogy, 77:289–313. LEA. 1998. Reevaluation of the oxygen isotopic ALLÈGRE, C. 2008. Isotope Geology. Cambridge composition of planktonic foraminifera: Experi- University Press, Cambridge, 512 p. mental results and revised palaeotemperature ANDO, A., B.T. HUBER, AND K.G. MACLEOD. equations. Paleoceanography, 13:150–160. 2010. Depth-habitat reorganization of planktonic BENTOV, S., C. BROWNLEE, AND J. EREZ. 2009. foraminifera across the Albian/Cenomanian The role of seawater endocytosis in the biominer- boundary. Paleobiology, 36:357–373. alization process in calcareous foraminifera. Pro- AURAHS, R., M. GÖKER, G. W. GRIMM, V. HEM- ceedings of the National Academy of Sciences, LEBEN, C. HEMLEBEN, R. SCHIEBEL, AND M. 106:21500–21504. KUČERA. 2009. Using the multiple analysis ap- BERGER, W. H. 1971. Sedimentation of planktonic proach to reconstruct relationships among plank- foraminifera. Marine Geology, 11:325–358. tonic foraminifera from highly divergent and BERGER, W. H. 1979a. Preservation of foraminifera, length-polymorphic SSU rDNA sequences. Bioin- p. 105–155. In J. LIPPS, W. H. BERGER, M. A. BU- formatics and Biology Insights, 3:155–157. ZAS, R. G. DOUGLAS, AND C. A. ROSS (eds.). Fo- AZE, T., T. H. G. EZARD, A. PURVIS, H. K. COX- raminiferal Ecology and Paleoecology, SEPM ALL, D. R. M. STEWART, B. S. WADE, AND P. N. Short Course No. 6. Society of Economic Paleon- PEARSON. 2011. A phylogeny of Cenozoic mac- tologists and Mineralogists, Houston, TX. roperforate planktonic foraminifera from fossil BERGER, W. H. 1979b. Stable isotopes in foramini- data. Biological Reviews, 86:900–927. fera, p. 156–198. In J. LIPPS, W. H. BERGER, M. A. BAR-MATTHEWS, M., A. AYALON, M. GILMOUR, BUZAS, R. G. DOUGLAS, AND C. A. ROSS (eds.). A. MATTHEWS, AND C. J. HAWKESWORTH. Foraminiferal Ecology and Paleoecology, SEPM 2003. Sea-land oxygen isotopic relationships from Short Course No. 6. Society of Economic Paleon- planktonic foraminifera and speleothems in the tologists and Mineralogists, Houston, TX. Eastern Mediterranean region and their implica- BICE, K. L., D. BIRGEL, P. A. MEYERS, K. A. tion for paleorainfall during interglacial intervals. DAHL, K.-U. HINRICHS, AND R. D. NORRIS. Geochimica et Cosmochimica Acta, 67, 3181– 2006. A multiple proxy and model study of Creta- 3199. ceous upper ocean temperatures and atmospheric BARRAS, C., J.-C. DUPLESSY, E. GESLIN, E. MI- CO2 concentrations. Paleoceanography, 21: CHEL, AND F. J. JORISSEN. 2010. Calibration of doi:10.1029/2005PA001203. δ18O of cultured benthic foraminiferal calcite as a BICE, K. L., B. T. HUBER, AND R. D. NORRIS. 2003. function of temperature. Biogeosciences, 7:1349– Extreme polar warmth during the Cretaceous 1356. greenhouse? Paradox of the late Turonian δ18O BARRON, E. J. 1987. Eocene equator-to-pole surface record at Deep Sea Drilling Project Site 511. Pa- ocean temperatures: a significant climate problem? leoceanography, 18, doi:10.1029/2002PA000848. Paleoceanography, 2:729–739. BICE, K. L., C. R. SCOTESE, D. SEIDOV, AND E. J. BASSINOT, F. C., L. D. LABEYRIE, E. VINCENT, X. BARRON. 2000. Quantifying the role of geo- QUIDELLEUR, N. J. SHACKLETON, AND Y. graphic change in Cenozoic ocean transport using LANCELOT. 1994. The astronomical theory of uncoupled atmosphere and ocean models. Palaeo- climate change and the age of the Brunhes- geography, Palaeoclimatology, Palaeoecology, Matuyama magnetic reversal. Earth and Planetary 161:295–310. Science Letters, 126:91–108. BIGG, G. R., AND E. J. ROHLING. 2000. An oxygen BÉ, A. W. H. 1980. Gametogenic calcification in a spi- isotope dataset for marine waters. Journal of Fo- nose planktonic foraminifer, Globigerinoides sac- raminiferal Research, 105:8527–8536. culifer (Brady). Marine Micropaleontology, BIJMA, J., H. J. SPERO, AND D. W. LEA. 1999. Reas- 5:283–310. sessing foraminiferal stable isotope geochemistry: BÉ, A. W. H., J. W. MORSE, AND S. M. HARRISON. Impact of the oceanic carbonate system (experi- 1975. Progressive dissolution and ultrastructural mental results), p. 489–512. In G. FISCHER (ed.). breakdown of planktonic foraminifera, p. 27–55. Use of Proxies in Paleoceanography: Examples for In W. V. SLITER, A. W. H. BÉ, AND W. H. BERGER the South Atlantic. Springer Verlag, Berlin. (eds.). Dissolution of Deep Sea Carbonates. BILLUPS, K., AND D. P. SCHRAG. 2002. Paleotemper- Cushman Foundation for Foraminiferal Research, tures and ice volume of the past 27 Myr revisited Special Publication, No. 13. with paired Mg/Ca and 18O/16O measurements on
28 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA
benthic foraminifera. Paleoceanography, logical Observatory, 690 pp. 17:doi:10.1029/2000PA000567 BURGESS, C. E., P. N. PEARSON, C. H. LEAR, H. E. BIRCH, H., H. K. COXALL, P. N. PEARSON, AND D. G. MORGANS, L. HANDLEY, R. D. PANCOST, KROON. Submitted. Planktonic foraminiferal AND S. SCHOUTEN. 2008. Middle Eocene cli- stable isotopes: ecological niches and disequilib- mate cyclicity in the southern Pacific: Implications rium fractionation effects. Marine Micropaleon- for global ice volume. Geology, 36:651–654. tology. BUSH, A. B. G., AND S. G. H. PHILANDER. 1997. BISSETT, A., T. R. NEU, AND D. DE BEER. 2011. The late Cretaceous: simulation with a coupled Dissolution of calcite in the twilight zone: bacte- atmosphere-ocean general circulation model. Pa- rial control of dissolution of sinking planktonic leoceanography, 12:495–516. carbonates is unlikely. PLoS ONE 6(11):e26404. CLARK, P. U., N. G. PISIAS, T. F. STOCKER, AND A. doi:10.1371/journal.pone.0026404. J. WEAVER. 2002. The role of the thermohaline BOERSMA, A., I. PREMOLI SILVA AND N. J. circulation in abrupt climate change. Nature, SHACKLETON. 1987. Atlantic Eocene plank- 415:863–869. tonic foraminiferal paleohydrographic indicators CLEMENT, A. C., AND L. C. PETERSON. 2006. and stable isotope paleoceanography. Paleocean- Mechanisms of abrupt climate change of the last ography, 2:287–331. glacial period. Reviews of Geophysics, 46: BOERSMA, A., N. J. SHACKLETON, M. A. HALL, RG4002, doi:10.1029/2006RG000204. AND Q. GIVEN. 1979. Carbon and oxygen isotope CLIMAP PROJECT MEMBERS. 1976. The surface of the records at DSDP Site 384 (North Atlantic) and ice age earth. Science, 191:1131–1137. some Paleocene paleotemperatures and carbon COPLEN, T. B. 1994. Reporting of stable hydrogen, isotope variations in the Atlantic Ocean. Initial carbon, and oxygen abundances. Pure and Applied Reports of the Deep Sea Drilling Project, 43:695– Chemistry, 2:273–276. 715. CORFIELD, R. M., AND J. E. CARTLIDGE. 1991. BOHATY, S. M., J. C. ZACHOS, AND M. L. DE- Isotopic evidence for the depth stratification of LANEY. 2012. Foraminiferal Mg/Ca evidence for fossil and recent Globigerinina: a review. Histori- Southern Ocean cooling during the Eocene–Oligo- cal Biology, 5:37–63. cene transition. Earth and Planetary Science Let- CORFIELD, R. M., M. A. HALL, AND M. D. ters, 317–318:251–261. BRASIER. 1990. Stable isotope evidence for fo- BORNEMANN, A., R. D. NORRIS, O. FRIEDRICH, raminiferal habitats during the Cenomanian / Tu- B. BECKMANN, S. SCHOUTEN, J. S. SINN- ronian anoxic event. Geology, 18:175–178. NINGE DAMSTÉ, J. VOGEL, P. HOFMANN, COSTA, K. B., F. A. L. TOLEDO, M. A. G. PIVEL, C. AND T. WAGNER. 2008. Isotopic evidence for A. V. MOURA, AND F. CHEMALE. 2006. Evalua- glaciation during the Cretaceous supergreenhouse. tion of two genera of benthic foraminifera for Science, 319:189–192. down-core paleotempertaure studies in the western BOUVIER-SOUMAGNAC, Y., AND J.-C. DU- South Atlantic. Brazilian Journal of Oceanogra- PLESSY. 1985. Carbon and oxygen isotopic com- phy, 54:75–84. position of planktonic foraminifera from labora- COXALL, H. K., P. N. PEARSON, N. J. SHACKLE- tory culture, plankton tows and recent sediment; TON, AND M. A. HALL. 2000. Hantkeninid depth implications for the reconstruction of paleocli- adaptation: an evolving life strategy in a changing matic conditions and of the global carbon cycle. ocean. Geology, 28:87–90. Journal of Foraminiferal Research, 15: 302–320. COXALL, H., P. A. WILSON, H. PÄLIKE, C. H. BRALOWER, T. J., J. C. ZACHOS, E. THOMAS, M. LEAR, AND J. BACKMAN. 2005. Rapid stepwise PARROW, C. K. PAULL, D. C. KELLY, I. PRE- onset of Antarctic glaciation and deeper calcite MOLI SILVA, W. V. SLITER, AND K. C. LOH- compensation in the Pacific Ocean. Nature, MANN. 1995. Late Paleocene to Eocene paleo- 433:53–57. ceanography of the equatorial Pacific Ocean: Sta- COXALL, H. K., P. A. WILSON, P. N. PEARSON, ble isotopes recorded at Ocean Drilling Program AND P. F. SEXTON. 2007. Iterative evolution of Site 865, Allison Guyot. Paleoceanography, digitate planktonic foraminifera. Paleobiology, 10:841–865. 33:495–516. BROECKER, W. S. 1989. The salinity contrast be- CRAIG, H. 1953. The geochemistry of stable isotopes tween the Atlantic and Pacific oceans during gla- of carbon. Geochimica et Cosmochimica Acta, cial time. Paleoceanography, 4:207–212. 3:53–72. BROECKER, W. S. 1998. Paleocean circulation during CRAIG, H. 1965. Measurement of oxygen isotope pa- the last deglaciation: a bipolar seesaw? Paleocean- leotemperatures, p. 162–182. In E. TONGIORGI ography, 13:doi:10.1029/97PA03707. (ed.). Stable Isotopes in Oceanographic Studies BROECKER, W. S., AND T.- H. PENG. 1982. Tracers and Paleotemperatures. Cons. Naz. Delle Ric., in the Sea. Eldigio Press, Lamont Doherty Geo- Spoleto, Italy.
29 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
CRAMER, B. S., K. G. MILLER, P. J. BARRETT, AND Proceedings of the National Academy of Sciences, J. D. WRIGHT. 2011. Late Cretaceous–Neogene 96:2864–2868. trends in deep ocean temperature and continental DECONTO, R. M., D. POLLARD, P. A. WILSON, H. ice volume: Reconciling records of benthic fo- PALIKE, C. H. LEAR, AND M. PAGANI. 2008. raminiferal geochemistry (δ18O and Mg/Ca) with Thresholds for Cenozoic bipolar glaciation. Na- sea level history. Journal of Geophysical Research, ture, 455:652–656. 116:C12023, doi:10.1029/2011JC007255. DEUSER, W. G., AND E. H. ROSS. 1989. Seasonally CRAMER, B. S., J. R. TOGGWEILER, J. D. WRIGHT, abundant planktonic foraminifera of the Sargasso M. W. KATZ, AND K. G. MILLER. 2009. Ocean Sea: Succession, deep-water fluxes, isotopic com- overturning since the Late Cretaceous: Inferences positions, and paleoceanographic implications. from a new benthic foraminiferal isotope compila- Journal of Foraminiferal Research, 19:268–293. tion. Paleoceanography, 24:PA4216, DEVEREUX, L. 1967. Oxygen isotope paleotempera- doi:10.1029/2008PA001683. ture measurements on New Zealand Tertiary fos- CROLL, J. 1875. Climate and Time, in their Geologi- sils. New Zealand Journal of Science, 10:988– cal Relations: A Theory of Secular Changes of the 1011. Earth's Climate. Daldy, Tbister and Company, D’HONDT, S., AND M. A. ARTHUR. 1995. Interspe- London, 577 p. cies variation in stable isotopic signals of Maas- CRONIN, T. 1999. Principles of Paleoclimatology, trichtian planktonic foraminifera. Paleoceanogra- Columbia University Press, New York, 560 p. phy, 10:123–135. CROWLEY, T. J., AND W. T. HYDE. 2008. Transient D’HONDT, S., AND M. A. ARTHUR. 1996. Late Creta- nature of late Pleistocene climate variability. Na- ceous oceans and the cool tropic paradox. Science, ture, 456:226–230. 271:1838–1841. CROWLEY, T. J., AND J. C. ZACHOS. 2000. Compari- DOKKEN, T. M., AND E. JANSSEN. 1999. Rapid son of zonal temperature profiles for past warm change in the mechanism of ocean convection time periods, p. 50–76. In B. T. HUBER, K. G. MA- during the glacial period. Nature, 401:458–461. CLEOD, AND S. L. WING (eds.). Warm Climates in DOUGLAS, R. G., AND S. M. SAVIN. 1971. Isotopic Earth History, Cambridge University Press, Cam- ananlyses of planktonic forminifera from the Ce- bridge UK. nozoic of the northwest Pacific, Leg 6. Initial Re- CURRY, W. B., AND T. J. CROWLEY. 1987. The δ13C ports of the Deep Sea Drilling Project, 6:1123– of equatorial Atlantic surface waters: Implications 1127. for Ice Age pCO2 levels. Paleoceanography, DOUGLAS, R. G., AND S. M. SAVIN. 1973. Oxygen 2:489–517. and carbon isotope analysis of Cretaceous and CURRY, W. B., R. C. THUNELL, AND S. HONJO. Tertiary foraminifera from the central north Pa- 1983. Seasonal changes in the isotopic composi- cific. Initial Reports of the Deep Sea Drilling Pro- tion of planktonic foraminifera collected in Pan- ject, 17:591–605. ama Basin sediment traps. Earth and Planetary DOUGLAS, R. G., AND S. M. SAVIN. 1978. Oxygen Science Letters, 64:33–43. isotopic evidence for the depth stratification of the CUTLER, K. B., R. L. EDWARDS, F. W. TAYLOR, H. Tertiary and Cretaceous planktonic foraminifera. CHENG, J. ADKINS, C. D. GALLUP, P. M. CUT- Marine Micropaleontology, 3:175–196. LER, G. S. BURR, AND A. L. BLOOM. 2003. DUNBAR, R. B., AND G. WEFER. 1984. Stable iso- Rapid sea-level fall and deep ocean temperature tope fractionation in benthic foraminifera from the change since the last interglacial period. Earth and Peruvian continental margin. Marine Geology, Planetary Science Letters, 206:253–271. 59:215–225. DANSGAARD, W., S. J. JOHNSON, AND J. MILLER. DUPLESSY, J. C., P. BLANC, AND A. W. H. BÉ. 1981. 1969. One thousand centuries of climatic record Oxygen-18 enrichment of planktonic forminifera from Camp Century on the Greenland ice sheet. due to gametogenic calcification below the Science, 166:377–380. euphotic zone. Science, 213:1247–1250. DANSGAARD, W., AND H. TAUBER. 1969. Glacier DUPLESSY, J. C., C. LALOU, AND A. C. VINOT. oxygen-18 content and Pleistocene ocean tempera- 1970. Differential isotopic fractionations in ben- tures. Science, 166:499–502. thic foraminifera and paleotemperatures reas- DARLING, K. F., AND C. M. WADE. 2008. The genetic sessed. Science, 138:250–251. diversity of planktonic foraminifera and the global DUTTON, A., K. C. LOHMANN, AND R. M. LECKIE. distribution of ribosomal genotypes. Marine Mi- 2005. Insights from the Paleogene tropical Pacific: cropaleontology, 67:216–238. Foraminiferal stable isotope and trace elemental DE VARGAS, C., R. NORRIS, L. ZANINETTI, S. W. results from Site 1209, Shatsky Rise. Paleocean- GIBB, AND J. PAWLOWSKI. 1999. Molecular ography, 20, doi:10.1029/2004PA001098. evidence of cryptic speciation in planktonic fora- ELDERFIELD, H., J. YU, P. ANAND, T. KIEFER, minifers and their relation to oceanic provinces. AND B. NYLAND. 2006. Calibrations for benthic
30 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA
foraminiferal Mg/Ca paleothermometry and the foraminifera from the Panama Basin. Nature, carbonate ion hypothesis. Earth and Planetary Sci- 298:841–844. ence Letters, 250:633–649. FAUL, K. L., A. C. RAVELO, AND M. L. DELANEY. EMILIANI, C. 1954a. Depth habitats of some pelagic 2000. Reconstructions of upwelling, productivity, foraminifera as indicated by oxygen isotope ratios. and photic zone depth in the eastern equatorial American Journal of Science, 252:149–158. Pacific Ocean using planktonic foraminiferal sta- EMILIANI, C. 1954b. Temperature of Pacific bottom ble isotopes and abundances. Journal of Forami- waters and polar superficial waters during the Ter- niferal Research. 30:110–125. tiary. Science, 119:853–855. FAURE, G., AND T. M. MENSING. 2005. Isotopes: EMILIANI, C. 1955. Pleistocene temperatures. Journal Principles and Applications, Third Edition. John of Geology, 63:538–578. Wiley and Sons, Hoboken, New Jersey, 897 p. EMILIANI, C. 1958. Ancient temperatures. Scientific FILLIPSON, H. L., J. M. BERNHARD, S. A. LIN- American, 198: 54–66. COLN, AND D. C. MCCORCKLE. 2010. A EMILIANI, C. 1961. The temperature decrease of sur- culture-based calibration of benthic foraminiferal face water in high latitudes and of abyssal-hadal paleotemperature proxies: δ18O and Mg/Ca results. water in open oceanic basins during the past 75 Biogeosciences, 7:1335–1347. million years. Deep Sea Research, 8:144–147. FRIEDRICH, O., R. D. NORRIS, AND J. ERBACHER. EMILIANI, C. 1966. Isotopic paleotemperatures. Sci- 2012b. Evolution of middle to Late Cretaceous ence, 154:851–857. oceans—a 55 m.y. record of Earth's temperature EMILIANI, C. 1978. The cause of the ice ages. Earth and carbon cycle. Geology, 40:107–110. and Planetary Science Letters, 37:349–352. FRIEDRICH, O., R. SCHIEBEL, P. A. WILSON, S. EMILIANI, C. 1991. Planktic/planktonic, nektic/ WELDEAB, C. J. BEER, M. J. COOPER, AND J. nektonic, benthic/benthonic. Journal of Paleontol- FIEBIG. 2012a. Influence of test size, water depth, ogy, 65:329. and ecology on Mg/Ca, Sr/Ca, δ18O and δ13C in EMILIANI, C. 1992. Pleistocene paleotemperatures. nine modern species of planktic foraminifera. Science, 257:1462. Earth and Planetary Science Letters, 319– EPSTEIN, S. R. 1997. The role of stable isotopes in 320:133–145. geochemistries of all kinds. Anuual Review of FRIEDMAN, I., AND J. R. O’NEIL. 1977. Compilation Earth and Planetary Sciences, 25:1–21. of stable isotope fractionation factors of geo- EPSTEIN, S., R. BUCHSBAUM, H. A. LOWEN- chemical interest, p. 1–12. In M. FLEISCHER (ed.). STAM, AND H. C. UREY. 1953. Revised Data of Geochemistry, U. S. Government Printing carbonate-water isotopic temperature scale. Geo- Office, Washington, DC. logical Society of America Bulletin, 64:1315– GANSSEN, G. M., F. J. C. PEETERS, B. METCALFE, 1325. P. ANAND, S. J. A. JUNG, D. KROON, AND G.-J. EREZ, J. 1978. Vital effect on stable-isotope composi- BRUMMER, 2011. Quantifying sea surface tem- tion seen in foraminifera and coral skeletons. Na- perature ranges of the Arabian Sea for the past ture, 273:199–202. 2000 years. Climate of the Past, 7:1337–1349. EREZ, J. 2003. The source of ions for biomineraliation GASPERI, J. T., AND J. P. KENNETT. 1993. Vertical in foraminifera and their implications for paleo- thermal structure of Miocene surface waters: ceanographic proxies. Reviews in Mineralogy and western equatorial paciic DSDP Site 289. Marine Geochemistry, 54:115–149. Micropaleontology, 22:235–254. EREZ, J., AND B. LUZ. 1983. Experimental paleotem- GRAHAM, D. W. B., B. H. CORLISS, M. L. perature equation for planktonic foraminifera. BENDER, AND L. D. KEIGWIN. 1981. Carbon Geochimica et Cosmochimica Acta, 47:1025– and oxygen isotopic disequilibria of Recent ben- 1031. thic foraminifera. Marine Micropaleontology, ERICSON, D. B., W. S. BROECKER, J. L. KULP, 6:483–497. AND G. WOLLIN. 1956. Late Pleistocene climates GROSSMAN, E. 1987. Stable isotopes in modern ben- and deep-sea sediments. Science, 124:385–389. thic foraminifera: a study of vital effect. Journal of ERICSON, D. B., AND G. WOLLIN. 1956. Micropale- Foraminiferal Research, 17:48–61. ontological and isotopic determinations of Pleisto- GROSSMAN, E. 2010. Oxygen isotope stratigraphy. In cene climates. Micropaleontology, 2:257–270. F. M. GRADSTEIN, J. G. OGG, AND A. SMITH (eds.). EZARD, T. H. G., T. AZE, P. N. PEARSON, AND A. A New Geologic Time Scale. Cambridge Univer- PURVIS. 2011. Interplay between changing cli- isty Press. mate and species’ ecology drives macroevolution- HAMBREY, M. J., W. U. EHRMANN, AND B. ary dynamics. Science, 332:349–351. LARSEN. 1991. Cenozoic glacial record of the FAIRBANKS, R. G., M. SVERDLOVE, R. FREE, P. H. Prydz Bay continental shelf, East Antarctica. Pro- WIEBE, AND A. W. H. BÉ. 1982. Vertical distribu- ceedings of the Ocean Drilling Program, Scientific tion and isotopic fractionation of living planktonic Results, 119:77–132.
31 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
HANSEN, J., M. SATO, P. KHARECHA, D. BEER- Paleogene. Geological Society of America Special LING, R. BERNER, V. MASSON-DELMOTTE, Paper, 369. M. PAGANI, M. RAYMO, D. L. ROYER, AND J. HUDSON, J. D., AND T. F. ANDERSON. 1989. Ocean C. ZACHOS. 2010 Target atmospheric CO2: temperatures and isotopic compositions through Where should humanity aim? Open Atmospheric time. Transactions of the Royal Society of Edin- Science Journal, 2:217–231. burgh: Earth Sciences, 80:183–192. HAY, W. W., AND E. ZAKEVICH. 1999. Cesare Emili- HUT, G. 1987. Consultants group meeting on stable ani (1922–1995): the founder of paleoceanogra- isotope reference samples for geochemical and phy. International Microbiology, 2:52–54. hydrological investigations. International Atomic HAYS, J. D., J. IMBRIE, AND N. J. SHACKLETON. Energy Agency, Vienna, 42 p. 1976. Variations in earth's orbit: Pacemaker of the HUTCHINSON, G. E. 1974. De rebus planktonicis. ice ages. Science, 194:1121–1132. Limnology and Oceanography, 19:360–361. HEMLEBEN, C., M. SPINDLER, AND O.R. ANDER- IMBRIE, J., A. BERGER, E. A. BOYLE, S. C. CLE- SON. 1989. Modern Planktonic Foraminifera. MENS, A. DUFFY, W. R. HOWARD, G. KUKLA, Springer-Verlag, New York, 363 p. J. KUTZBACH, D. G. MARTINSON, A. MCIN- HERBERT, T. D., L. C. PETERSON, K. T., LAW- TYRE, A. C. MIX, B. MOLFINO, J. J. MORLEY, RENCE, AND Z. LIU. 2010. Tropical ocean tem- L. C. PETERSON, N. G. PISIAS, W. L. PRELL, peratures over the past 3.5 million years. Science, M. W. RAYMO, N. J. SHACKLETON, AND J. R. 328:1350–1354. TOGGWEILER. 1993. On the structure and origin HODELL, D. A., AND A. VAYAVANANDA. 1993. of major glaciation cycles 2. The 100,000-year Middle Miocene paleoceanography of the western cycle. Paleoceanography, 8:699–735. equatorial Pacific and the evolution of Globoro- IMBRIE, J., AND K. P. IMBRIE. 1979. Ice ages: solving talia (Fohsella). Marine Micropaleontology, the mystery. Harvard University Press, Cambridge 22:279–310. Massachusetts, 224 p. HOEFS, J. 2009. Stable isotope geochemistry, 6th Edi- IVANOVA, E. 2000. Late Quaternary monsoon history tion. Springer-Verlag, Berlin, 285 p. and paleoproductivity of the western Arabian Sea. HOLLIS, C. J., TAYLOR, K. W. R., L. HANDLEY, R. PhD thesis, Free University, Amsterdam, 172 p. D. PANCOST, M. HUBER, J. B. CREECH, B. JAFFRÉS, J. B. D., G. A. SHIELDS, AND K. WALL- HINES, E. M. CROUCH, H. E. G. MORGANS, J. MANN. 2007. The oxygen isotope evolution of S. CRAMPTON, S. GIBBS, S. SCHOUTEN, P. N. seawater: A critical review of a long-standing con- PEARSON, AND J. C. ZACHOS. 2012. Southwest troversy and an improved geological water cycle Pacific marine temperature history from late Pa- model for the past 3.4 billion years. Earth Science leocene to middle Eocene: Revisited. Earth and Reviews, 83:83–122. Planetary Science Letters. JAHNKE, R. A., D. B. CRAVEN, D. C. HUBER, M. 2008. A hotter greenhouse? Science, MCCORCKLE, AND C. E. REIMERS. 1997. Ca- 321:353–354. CO3 dissolution in California continental margin HUBER, B. H., J. BIJMA, AND K. DARLING. 1997. sediments: The influence of organic matter remin- Cryptic speciation in the living planktonic fora- eralisation. Geochimica et Cosmochimica Acta, minifer Globigerinella siphonifera (d'Orbigny). 61:3587–3604. Paleobiology, 23: 33–62. KAHN, M. I. 1979. Non-equilibrium oxygen and car- HUBER, M., AND R. CABALLERO. 2011. The early bon isotope fractionation in tests of living plank- Eocene equable climate problem revisited. Cli- tonic foraminifera. Oceanologia Acta, 2:195–208. mate of the Past Discussions, 7:241–304. KAIHO, K., AND M. LAMOLDA. 1999. Catastrophic HUBER, B. H., D. A. HODELL, AND C. P. HAMIL- extinction of planktonic foraminifera at the Cret- TON. 1995. Middle–Late Cretaceous climate of caeous–Tertiary boundary evidenced by stable the southern high latitudes: Stable isotopic evi- isotopes and foraminiferal abundance at Caravaca, dence for minimal equator-to-pole thermal gradi- Spain. Geology, 27:355–358. ents. Geology, 107:1164–1191. KATZ, M., D. R. KATZ, J. D. WRIGHT, K. G. HUBER, B. T., R. D. NORRIS, AND K. G. MACLEOD. MILLER, D. K. PAK, N. J. SHACKLETON, AND 2002. Deep-sea paleotemperature record of ex- E. THOMAS. 2003. Early Cenozoic benthic fo- treme warmth during the Cretaceous. Geology, raminiferal isotopes: species reliability and inter- 30:123–126. species correction factors. Paleoceanography, 18: HUBER, M., L. C. SLOAN, AND C. SHELLITO. 2003. doi:10.1029/2002PA000798. Early Paleogene oceans and climate: fully coupled KATZ, M., K. G. MILLER, J. D. WRIGHT, B. S. modeling approach using the NCAR CCSM, p. WADE, J. V. BROWNING, B. S. CRAMER, AND 25–47 In S. L. WING, P. D. GINGERICH, B. Y. ROSENTHAL. 2008. Stepwise transition from SCHMITZ, AND E. THOMAS (eds.). Causes and con- the Eocene greenhouse to the Oligocene icehouse. sequences of globally warm climates in the early Nature Geoscience, 1:329–333.
32 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA
KEIGWIN, L. D. 1980. Paleoceanographic changes in LASKAR, J. 1999. The limits of Earth orbital calcula- the Pacific across the Eocene–Oligocene bound- tions for geological time-scale use. Philosophical ary. Nature, 287:722–725. Transactions of the Royal Society, A, 357:1735– KEIGWIN, L. D., AND B. H. CORLISS. 1986. Stable 1759. isotopes in late middle Eocene to Oligocene fora- LASKAR, J., A. FIENGA, M. GASTINEAU, AND H. minifera. Geological Society of America, Bulletin, MANCHE. 2011. La2010: A new orbital solution 97:335–345. for the long term motion of the Earth. Earth and KELLER, G. 1985. Depth stratification of planktonic Planetary Astrophysics: arXiv:1103.1084v1. foraminifera in the Miocene ocean, p. 177–195. In LEA, D. W. 2011. Elemental and isotopic proxies of J. P. KENNETT (ed.). The Miocene Ocean: Paleo- past ocean temperatures, p. 227–253. In H. D. ceanography and Biogeography. Geological Soci- HOLLAND, AND K. K. TUREKIAN (eds.). Isotope ety of America, Memoir, No.163. geochemistry from the treatise on geochemistry. KELLY, D. C., T. J. BRALOWER, J. C. ZACHOS, I. Elsevier, London. PREMOLI SILVA, AND E. THOMAS. 1996. Rapid LEAR, C. H., T. R. BAILEY, P. N. PEARSON, H. K. diversification of planktonic foraminifera in the COXALL, AND Y. ROSENTHAL. 2008. Cooling tropical Pacific (ODP Site 865) during the late and ice growth across the Eocene–Oligocene tran- Paleocene thermal maximum. Geology, 24:423– sition. Geology, 36:251–254. 426. LEAR, C. H., H. ELDERFIELD, AND P. A. WILSON. KENNETT, J. P., AND SHACKLETON, N. J. 1976. 2000. Cenozoic deep-sea temperatures and global Oxygen isotopic evidence for the development of ice volumes from Mg/Ca in benthic foraminiferal the psychrosphere 38 Myr ago. Nature, 260:513– calcite. Science, 287:269–272. 515. LEAR, C. H., AND Y. ROSENTHAL. 2006. Benthic KILLINGLEY, J. S. 1983. Effects of diagnetic recrys- foraminiferal Li/Ca: insights into seawater carbon- tallization on 18O/16O values in deep sea sedi- ate saturation state. Geology, 34:985–988. ments. Nature, 201:594–597. LEDUC, G., L. VIDAL, O. CARTAPANIS, AND E. KILLINGLEY, J. S., R. F. JOHNSON, AND W. H. BARD. 2009. Modes of eastern equatorial Pacific BERGER. 1981. Oxygen and carbon isotopes of thermocline variability: Implications for ENSO individual shells of planktonic foraminifera from dynamics over the last glacial period. Paleocean- Ontong-Java Plateau, Equatorial Pacific. Palaeo- ography, 24, PA3202, doi:10.1029/2008PA001701. geography, Palaeoecology, Palaeoclimatology, LEE, J. L., J. PAWLOWSKI, J.- P. DEBENAY, J. 33:193–204. WHITTAKER, F. BANNER, A. J. GOODAY, O. KIM, S.-T., AND J. R. O'NEIL. 1997. Equilibrium and TENDAL., J. HAYNES, AND W. W. FABER. non-equilibrium oxygen isotope effects in syn- 2000. Phylum Granuloreticulosa. In J. L. LEE, G. thetic carbonates. Geochimica et Cosmochimica F. LEEDALE, AND P. BRADBURY (eds.). An Illus- Acta, 61:3461–3475. trated Guide to the Protozoa (2nd edn.): Society of KOUTAVAS, A., P. B. DE MENOCAL, G. C. OLIVE Protozoologists, Lawrence, Ks. 2 vols., 1432 p. COL, AND J. LYNCH-STEIGLITZ. 2006. Mid- LEGRANDE, A. N., AND G. A. SCHMIDT. 2006. Holocene El Nino-Southern Oscillation (ENSO) Global gridded data of the oxygen isotopic com- attenuation revealed by individual foraminifera in position of seawater. Geophysical Research Let- eastern tropical Pacific sediments. Geology, ters, 33: doi:10.1029/2006gl026011. 34:993–996. L’HOMME, N., G. K. C. CLARKE, AND C. RITZ. KOZDON, R., D. C. KELLY, N. T. KITA, J. H. FOUR- 2005. Global budget of water isotopes inferred NELLE, AND J. W. VALLEY. 2011. Planktonic from polar ice sheets. Geophysical Research Let- foraminiferal oxygen isotope analysis by ion mi- ters, 32:L20502, doi:10.1029/2005GL023774. croprobe technique suggests warm tropical sea LISIECKI, L. E., AND M. E. RAYMO. 2005. A Plioce- surface temperatures during the Early Paleogene. ne–Pleistocene stack of 57 globally distributed Paleoceanography, 26:PA3206, benthic δ18O records. Paleoceanography, doi:10.1029/2010PA002056. 20:PA1003, doi:10.1029/2004PA001071. KUCERA, M., AND K. F. DARLING. 2002. Cryptic LOHMANN, G. P. 1995. A model for variation in the species of planktonic foraminifera: their effect on chemistry of planktonic foraminifera due to sec- palaeoceanographic reconstructions. Philosophical ondary calcification and selective dissolution. Pa- Transactions of the Royal Society of London, A, leoceanography, 10:445–447. 360:695–718. LUTHI, D., M. LE FLOCH, B. BEREITER, T. BLU- KUCERA, M., A. ROSELL-MELÉ, R. SCHNEIDER, NIER, J.-M. BARNOLA, U. SIEGENTHALER, C. WAELBROECK, AND M. WEINELT. 2005. D. RAYNAUD, J. JOUZEL, H. FISCHER, K. Multiproxy approach for the reconstruction of the KAWAMURA, AND T. F. STOCKER. 2008. High glacial ocean (MARGO). Quaternary Science Re- resolution carbon dioxide concentration record views, 24:813–819. 650,000–800,000 years before present. Nature,
33 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
453:379–382. MEGGERS, A. PAUL, AND G. WEFER. 2003. LYNCH-STIEGLITZ, J., W. B. CURRY, AND N. Temperature:δ18O relationships of planktonic SLOWEY. 1999. A geostrophic transport estimate foraminifera collected from surface waters. Palae- for the Florida current from the oxygen isotopic ogeography, Palaeoclimatology, Palaeoecology, composition of benthic foraminifera. Paleoceanog- 202:143–152. raphy, 14:360–373. MULITZA, S., B. DONNER, G. FISCHER, A. PAUL, MACKENSEN, A. 2001. Oxygen and carbon stable J. PÄTZOLD, C. RÜHLEMANN, AND M. SEGL. isotope tracers of Weddell Sea water masses: New 1999. The South Atlantic oxygen isotope record of data and some paleoceanographic implications. planktic foraminifera, p. 121–142. In G. WEFER. S. Deep-Sea Research Part 1, Oceanographic Re- MULITZA, AND V. RATMEYER (eds.). The South search Papers, 48:1401–1422. Atlantic in the Late Quaternary: Reconstruction of MCARTHUR, J. M., N. M. M. JANSSEN, S. REBOU- Material Budgets and Current Systems, Springer- LET, M. J. LENG, M. F. THIRLWALL, AND B. Verlag, Berlin. VAN DE SCHOOTBRUGGE. 2007. Paleotem- MULITZA, S., A. DÜRKOOP, W. HALE, G. WEFER, peratures, polar ice-volume, and isotope stratigra- AND H. S. NIEBLER. 1997. Planktonic foramini- phy (Mg/Ca, δ18O, δ13C, 87Sr/86Sr): The Early Cre- fera as recorders of past surface-water stratifica- taceous (Berriasian, Valanginian, Hauterivian). tion. Geology, 25:335–338. Palaeogeography, Palaeoclimatology, Palaeoecol- NATIONAL RESEARCH COUNCIL. 2011. Scientific ogy, 248:391–430. Ocean Drilling: Accomplishments and Challenges. MCCORKLE, D. C., J. M. BERNHARD, C. J. HINTZ, National Academies Press, Washington DC, 145 p. J. K. BLANKS, G. T. CHANDLER, AND T. J. NIER, A. O. 1947. A mass spectrometer for isotopes SHAW. 2008. The carbon and oxygen stable iso- and gas analysis. Review of Scientific Instru- topic composition of cultured benthic foramini- ments, 18:398–411. fera, p. 135–154. In W. E. N. AUSTIN, AND R. J. NORRIS, R. D., R. M. CORFIELD, AND J. E. JAMES (eds.). Biogeochemical Controls on Paleo- CARTLIDGE. 1993. Evolution of depth ecology ceanographic Environmental Proxies, Geological in the planktonic foraminifera lineage Globoro- Society of London, Special Publications, No. 303. talia (Fohsella). Geology, 11:975–978. MCCREA, J. M. 1950. On the isotopic chemistry of NORRIS, R. D., R. M. CORFIELD, AND J. E. carbonates and a paleotemperature scale. Journal CARTLIDGE. 1996. What is gradualism? Cryptic of Chemical Physics, 18:849–857. speciation in globorotaliid foraminifera. Paleo- MCKINNEY, C. R., J. M. MCCREA, S. EPSTEIN, H. biology, 22:386–405. A. ALLEN, AND H. C. UREY. 1950. Improve- NORRIS, R. D., AND P. A. WILSON. 1998. Low- ments in mass spectrometers for the measurement latitude sea-surface temperatures for the mid- of small differences in isotopic abundance ratio. Cretaceous and the evolution of planktic foramini- Review of Scientific Instruments, 21:724–730. fera. Geology, 26:823–826. MILANKOVITCH, M. 1941. Kanon der Erdbestrah- OLAUSSON, E. 1965. Evidence of climatic changes in lung und seine Anwendung auf das Eiszeitenprob- North Atlantic deep-sea cores, with remarks on lem. Königlich Serbische Akademie, Belgrade, isotopic temperature analysis. Progress in Ocean- 132 p. ography, 3:221–252. MILLER, K. G., R. G. FAIRBANKS, AND G. S. OLAUSSON, E. 1996. The Swedish Deep-Sea Expedi- MOUNTAIN. 1987. Tertiary oxygen isotope syn- tion with the “Albatross” 1947–1948: A Summary thesis, sea level history, and continental margin of Sediment Core Studies. Novum , Grafiska AB, erosion. Paleoceanography, 2:1–19. Göteborg, 98 p. MILLIMAN, J. D., P. J. TROY, W. M. BALCH, A. K. O'NEIL, J. R., R. N. CLAYTON, AND T. K. MAYEDA. ADAMS, Y.-H. LI, AND F. T. MACKENZIE. 1999. 1969. Oxygen isotope fractionation in divalent Biologically mediated dissolution of calcium car- metal carbonates. Journal of Chemical Physics, bonate above the chemical lysocline? Deep Sea 51:5547–5558. Research Part I: Oceanographic Research Papers, PÄLIKE, H., J. FRAZIER, AND J. C. ZACHOS. 2006. 46:1653–1699. Extended orbitally forced palaeoclimate records MOHTADI, M., D. W. OPPO, A. LCKGE, R. DEPOL- from the Atlantic Ceara Rise. Quaternary Science 34HOLZ, S. STEINKE, J. GROENEVELD, N. Reviews, 25:3138–3149. HEMME, AND D. HEBBELN. 2011. Reconstruct- PÄLIKE, H, J., L. LASKAR, AND N. J. SHACKLE- ing the thermal structure of the upper ocean: In- TON. 2004. Constraints on astronomical parame- sights from modern planktonic foraminifera shells ters from the geological record of the past 25 Ma. chemistry and alkenones in modern sediments of Earth and Planetary Science Letters, 182:1–14. the tropical eastern Indian Ocean. Paleoceanogra- PEARSON, P. N. 1998. Stable isotope and the study of phy, 24, PA3219, doi:10.1029/2011PA002132. evolution in planktonic foraminifera. The Paleon- MULITZA, S., D. BOLTOVSKOY, B. DONNER, H. tological Society Papers, 4:138–178.
34 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA
PEARSON, P. N., AND C. E. BURGESS. 2008. Fora- cene sea surface temperatures and global ice minifer test preservation and diagenesis: compari- volumes. Marine Micropaleontology, 9:111–134. son of high latitude sites, p. 59–72. In W. E. N. PROKOPH, A., G. A. SHIELDS, AND J. VEIZER. AUSTIN, AND R. H. JAMES (eds.). Biogeochemical 2008. Compilation and time-series analysis of Controls on Paleoceanographic Environmental marine carbonate δ18O, δ13C, 87Sr/86Sr and δ34S Proxies, Geological Society of London, Special database through Earth history. Earth Science Re- Publications, No. 303. views, 87:113–133. PEARSON, P. N., P. DITCHFIELD, AND N. J. PURTON, L., AND M. D. BRASIER. 1999. Giant pro- SHACKLETON. 2002. Palaeoclimatology (Com- tist Nummulites and its Eocene environment: Life munications arising): Tropical temperatures in span and habitat insights from δ18O and δ13C data greenhouse climates. Nature, 419:898. from Nummulites and Venericardia, Hampshire PEARSON, P. N., P. W. DITCHFIELD, J. SINGANO, basin, UK. Geology, 27:711–714. K. G. HARCOURT-BROWN, C. J. NICHOLAS, PUSZ, A. F., R. C. THUNELL, AND K. G. MILLER. R. K. OLSSON, N. J. SHACKLETON, AND M. A. 2011. Deep water temperature, carbonate ion, and HALL. 2001. Warm tropical sea surface tempera- ice volume changes across the Eocene–Oligocene tures in the Late Cretaceous and Eocene epochs. climate transition. Paleoceanography, 26, PA2205. Nature, 413:481–487. doi:10.1029/2010PA001950. PEARSON, P. N., AND N. J. SHACKLETON. 1995. RASMUSSEN, T. L., AND E. THOMSEN. 2009. Stable Neogene multispecies planktonic foraminifer sta- isotope signals from brines in the Barents Sea: ble isotope record, Site 871, Limalok guyot. Pro- implications for brine formation during the last ceedings of the Ocean Drilling Program, Scientific glaciation. Geology, 37:903–906. Results, 144:21–59. RATHMANN, S., AND H. KUHNERT. 2008. Carbonate PEARSON, P. N., N. J. SHACKLETON, AND M. A. ion effect on Mg/Ca, Sr/Ca and stable isotopes on HALL. 1993. Stable isotope paleoecology of mid- the benthic foraminifera Oridorsalis umbonatus dle Eocene planktonic foraminifera and multi- off Namibia. Marine Micropaleontology, 66:120– species isotope stratigraphy, DSDP Site 523, South 133. Atlantic. Journal of Foraminiferal Research, RAYMO, M. E., L. LISIECKI, AND K. NISANCIO- 23:123–140. GLU. 2006. Plio–Pleistocene ice volume, Antarc- PEARSON, P. N., N. J. SHACKLETON, AND M. A. tic climate, and the global δ18O record. Science, HALL. 1997. Stable isotopic evidence for the 313:492–495. sympatric divergence of Globigerinoides trilobus REICHART, G.-J., F. JORISSEN, P. ANSCHUTZ, AND and Orbulina universa (planktonic foraminifera). P. R. D. MASON. 2003. Single foraminiferal test Journal of the Geological Society, 154:295–302. chemistry records the marine environment. Geol- PEARSON, P. N., B. E. VAN DONGEN, C. J. NICHO- ogy, 31:355–358. LAS, R. D. PANCOST, S. SCHOUTEN, J. M. RINK, S., M. KÜHL, J. BIJMA, AND H. SPERO. 1998. SINGANO, AND B. S. WADE. 2007. Stable warm Microsensor studies of photosynthesis and respira- tropical climate through the Eocene epoch. Geol- tion in the symbiotic foraminifer Orbulina uni- ogy, 35:211–214. versa. Marine Biology, 131:583–595. PEARSON, P. N., AND WADE, B. S. 2009. Taxonomy ROBERTS, C. D., A. N. LEGRANDE, AND A. K. TRI- and stable isotope paleoecology of well-preserved PATI. 2011. Sensitivity of seawater oxygen iso- planktonic foraminifera from the uppermost Oli- topes to climatic and tectonic boundary conditions gocene of Trinidad. Journal of Foraminiferal Re- in an early Paleogene simulation with GISS Model search, 39:191–217. E-R. Paleoceanography 26:PA4203, PECK, V. L., J. YU, S. KENDER, AND C. R. RIES- doi:10.1029/201pa002025. SELMAN. 2010. Shifting ocean carbonate chem- RODHE, W. 1974. Plankton, planktis, planktonic. istry during the Eocene–Oligocene climate transi- Limnology and Oceanography, 19:360. tion: implications for deep ocean Mg/Ca paleo- ROHLING, E. J., AND S. COOKE. 1999. Stable oxygen thermometry. Paleoceanography, 25, PA4219. and carbon isotopes in foraminiferal carbonate doi:10.1029/2009PA001906. shells, p. 239–258. In B. SEN GUPTA (ed.). Modern PISIAS, N. G., D. G. MARTINSON, T. C. MOORE, Foraminifera. Kluwer, Dordrecht. N. J. SHACKLETON, W. PRELL, J. HAYS, AND ROLLION-BARD, C., J. EREZ, AND T. ZILBERMAN. G. BODEN. 1984. High resolution stratigraphic 2008. Intra-shell oxygen isotope ratios in the ben- correlation of benthic oxygen isotopes spanning thic foraminifera genus Amphistegina and the in- the last 300,000 years. Marine Geology, 56:119– fluence of seawater carbonate chemistry and tem- 136. perature on this ratio. Geochimica et Cosmo- POORE, R.Z., AND R. K. MATTHEWS. 1984. Oxygen chimica Acta, 72:6006–6014. isotope ranking of late Eocene and Oligocene ROSMAN, J. R., AND P. D. TAYLOR. 1998. Isotopic planktonic foraminifers: implications for Oligo- compositions of the elements (technical report):
35 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
commission on atomic weights and isotopic abun- ‘Glassy’ versus ‘Frosty’. Geochemistry, Geophys- dances. Pure and Applied Chemistry, 70:217–235. ics, Geosystems, 7:Q12P19, SAUTTER, L. R., AND R. C. THUNELL. 1991. Sea- doi:10.1029/2006GC001291. sonal variability in δ18O and δ13C of planktonic SHACKLETON, N. J. 1967. Oxygen isotope analyses foraminifera from an upwelling environment: and Pleistocene temperatures reassessed. Nature, sediment trap results from the San Pedro Basin, 215:15–17. Southern California Bight. Paleoceanography, SHACKLETON, N. J. 1974. Attainment of isotopic 6:307–334. equilibrium between ocean water and the ben- SAVIN, S. M. 1977. The history of Earth's surface thonic foraminifera genus Uvigerina, p. 203–209. temperature during the past 100 million years. In Isotopic changes in the ocean during the last Annual Reviews of Earth and Planetary Science, glacial. Cent. Nat. Rech. Sci. Colloq. Int., No. 219. 5:319–355. SHACKLETON, N. J. 1986. Paleogene stable isotope SAVIN, S. M., L. ABEL, E. BARRERA, D. A. events. Palaeogeography, Palaeoclimatology, Pa- HODELL, G. KELLER, J. P. KENNETT, J. KILL- laeoecology, 57:91–102. INGLEY, M. MURPHY, AND E. VINCENT. 1985. SHACKLETON, N. J. 2000. The 100,000-year ice-age The evolution of Miocene surface and near-surface cycle identified and found to lag temperature, car- temperatures, oxygen isotopic evidence, p. 49–82. bon dioxide, and orbital eccentricity. Science, In J. P. KENNETT (ed.). The Miocene Ocean: Pa- 289:1897–1902. leoceanography and Biogeography. Geological SHACKLETON, N., AND A. BOERSMA. 1981. The Society of America, Memoir, No.163. climate of the Eocene ocean. Journal of the Geo- SAVIN, S. M., R. G. DOUGLAS, AND F. G. STEHLI. logical Society of London, 138:153–157. 1975. Tertiary marine paleotemperatures. Geologi- SHACKLETON, N. J., R. M. CORFIELD, AND M. A. cal Society of America, Bulletin, 86:1499–1510. HALL. 1985. Stable isotope data and the ontogeny SCHIEBEL, R. 2002. Planktonic foraminiferal sedi- of Paleocene planktonic foraminifera. Journal of mentation and the marine calcite budget. Global Foraminiferal Research, 15:321–336. Biogeochemical Cycles, 11:125–133. SHACKLETON, N. J., S. J. CROWHURST, G. P. SCHIFFELBEIN, P., AND S. HILLS. 1984. Direct as- WEEDON, AND J. LASKAR. 1999. Astronomical sessment of stable isotope variability in planktonic calibration of Oligocene–Miocene time. Philo- foraminifera populations. Palaeogeography, Pa- sophical Transactions of the Royal Society of laeoclimatology, Palaeoecology, 48:197–213. London, Series A, 357:1907–1929. SCHMIDT, G. A. 1999. Forward modeling of carbon- SHACKLETON, N. J., AND J. P. KENNETT. 1975. ate proxy data from planktonic foraminifera using Paleotemperature history of the Cenozoic and the oxygen isotope tracers in a global ocean model. initiation of Antarctic glaciation: oxygen and car- Paleoceanography, 14:482–497. bon isotope analyses in DSDP Sites 277, 279, and SCHMIDT, G. A., G. R. BIGG, AND E. J. ROHLING. 281. Initial Reports of the Deep Sea Drilling Pro- 1999. Global seawater Oxygen-18 database— ject, 29:743–755. v1.21: http://data.giss.nasa.gov/o18data/ SHACKLETON, N. J., AND N. D. OPDYKE. 1973. SCHNEIDER, C. E., AND J. P. KENNETT. 1996. Iso- Oxygen isotope and palaeomagnetic stratigraphy topic evidence for interspecies habitat differences of equatorial Pacific core V28-238: Oxygen iso- during evolution of the Neogene planktonic fora- tope temperatures and ice volumes on a 105 year minferal clade Globoconella. Paleobiology, and 106 year scale. Quaternary Research, 3:39–55. 22:282–303. SHACKLETON, N. J., J. D. H. WISEMAN, AND H. A. SCHRAG, D. P. 1999. Effects of diagenesis on the iso- BUCKLEY. 1973. Non-equilibrium isotopic frac- topic record of late Paleogene tropical sea surface tionation between seawater and planktonic fo- temperatures. Chemical Geology, 1999:215–224. raminiferal tests. Nature, 242:177–179. SCHRAG, D. P., J. F. ATKINS, K. MCINTYRE , J. L. SHIEH, Y.-T., C.-F. YOU, K.-S. SHEA, AND C.-S. ALEXANDER, D. A. HODELL, C. D. CHAR- HONG. 2002. Identification of artifacts in forami- LES, AND J. F. MCMANUS. 2002. The oxygen niferal tests using carbon and oxygen isotopes: isotopic composition of seawater during the Last Journal of Asian Earth Sciences, 21:1–5. Glacial Maximum. Quaternary Science Reviews, SPERO, H. J. 1992. Do planktic foraminifera accu- 21:331–342. rately record shifts in the carbon isotopic composi- SELF-TRAIL, J. M., AND E. L. SEEFELT. 2004. Rapid tion of sea water ΣCO2? Marine Micropaleontol- dissolution of calcareous nannofossils from freshly ogy, 19:275–285. cored sediments, USA. Journal of Nannoplankton SPERO, H. J. 1998. Life history and stable isotope Research, 26:94. geochemistry of planktonic foraminifera. The Pa- SEXTON, P. F., P. A. WILSON, AND P. N. PEARSON. leontological Society Papers, 4:7–36. 2006. Microstructural and geochemical perspec- SPERO, H. J., J. BIJMA, D. W. LEA, AND B. E. BE- tives on planktic foraminiferal preservation: MIS. 1997. Effect of seawater carbonate concen-
36 PEARSON: OXYGEN ISOTOPES IN FORAMINIFERA
tration on foraminiferal carbon and oxygen iso- AND H. STRAUSS. 1999. 87Sr/86Sr, δ13C and δ18O topes. Nature, 390:497–500. evolution of Phanerozoic seawater. Chemical Ge- SPERO, H. J., AND D. W. LEA. 1993. Intraspecific sta- ology, 161:59–88. ble isotope variability in the planktic foraminifera VEIZER, J., P. BRUCKSCHEN, F. PAWELLEK, A. Globigerinoides sacculifer: results from laboratory DIENER, G. PODLAHA, G.A.F. CARDEN, T. experiments. Marine Micropaleontology, 22:221– JASPER, C. KORTE, H. STRAUSS, K. AZMY, 234. AND D. ALA. 1997. Oxygen isotope evolution of SPERO, H. J., K. M. MIELKE, E. M. KALVE, D. W. Phanerozoic seawater. Palaeogeography, Palaeo- LEA, AND D. K. PAK. 2003. Multispecies ap- climatology, Palaeoecology, 132:159–172. proach to reconstructing eastern equatorial Pacific VEIZER, J., Y. GODDERIS, AND L. M. FRANÇOIS. thermocline hydrography during the past 360 kyr. 2000. Evidence for decoupling of atmospheric Paleoceanography, 18, CO2 and global climate during the Phanerozoic doi:10.1029/2002PA000814. eon. Nature, 408:698–701. STEWART, D. R. M., P. N. PEARSON, P. W. DITCH- VIDAL, L., L. LABEYRIE, AND T. C. E. VAN WEER- FIELD, AND J. M. SINGANO. 2004. Miocene ING. 1998. Benthic δ18O records in the North At- ocean temperatures: evidence from three excep- lantic over the last glacial period (60–10 k.yr.): tionally preserved foraminiferal assemblages from Evidence for brine formation. Paleoceanography, Tanzania. Journal of African Earth Sciences, 40, 13:245–251. 173–189. VINCENT, E., J. S. KILLINGLEY, AND W. H. STOTT, L. D., AND C. M. TANG. 1996. Reassessment BERGER. 1985. Miocene oxygen and carbon iso- of tropical sea surface δ18O temperatures. Paleo- tope stratigraphy of the tropical Indian Ocean. p. ceanography, 11:37–56. 103–130. In J. P. KENNETT (ed.). The Miocene TANG, C. M., AND L. D. STOTT. 1993. Seasonal salin- Ocean: Paleoceanography and Biogeography. ity changes during Mediterranean sapropel deposi- Geological Society of America, Memoir, No.163. tion 9,000 years B.P.: Evidence from isotopic WADE, B. S., J. P. HOUBEN, W. QUAIJTAAL, S. analyses of individual planktonic foraminifera. SCHOUTEN, Y. ROSENTHAL, K. G. MILLER, Paleoceanography, 8:473–494. M. E. KATZ, J. D. WRIGHT, AND H. TINDALL, J., R. FLECKER, P. VALDES, D. N. BRINKHUIS. 2012. Multiproxy record of abrupt SCHMIDT, P. MARKWICK, AND J. HARRIS. sea-surface cooling across the Eocene–Oligocene 2010. Modelling the oxygen isotope distribution of transition in the Gulf of Mexico. Geology, 40:159– ancient seawater using a coupled ocean- 162. atmosphere GCM: Implications for reconstructing WADE, B. S., AND H. PÄLIKE. 2004. Oligocene cli- early Eocene climate. Earth and Planetary Science mate dynamics. Paleoceanography, 19:PA4019, Letters, 292:265–273. doi:10.1029/2004PA001042. UCHIKAWA, J., AND R. E. ZEEBE. 2010. Examining WADE, B. S., P. N. PEARSON, W. A. BERGGREN, possible effects of seawater pH decline on forami- AND H. PÄLIKE. 2011. Review and revision of niferal stable isotopes during the Paleocene–Eo- Cenozoic tropical planktonic foraminiferal biostra- cene thermal maximum. Paleoceanography, tigraphy and calibration to the geomagnetic polar- 25:PA2216, doi:10.1029/2009PA001864. ity and astronomical timescale. Earth Science Re- UJIIÉ, Y. AND J. H. LIPPS. 2009. Cryptic diversity in views, 104:111–142. planktic foraminifera in the northwest Pacific WAELBROECK, C., S. MULITZA, H. SPERO, T. Ocean. Journal of Foraminiferal Research, DOKKEN, T. KIEFER, AND E. CORTIJO. 2005. A 39:145–154. global compilation of late Holocene planktonic UREY, H. C. 1947. The thermodynamic properties of foraminiferal δ18O: relationship between surface isotopic substances. Journal of the Chemical Soci- water temperature and δ18O. Quaternary Science ety of London, 1947:562–581. Reviews, 24:853–868. UREY, H. C. 1948. Oxygen isotopes in nature and in WALLMANN, K. 2001. The geological water cycle the laboratory. Science, 108:489–496. and the evolution of marine δ18O values. Geo- UREY, H. C., H. LOWENSTHAM, S. EPSTEIN, AND chimica et Cosmochimica Acta, 65:2469–2485. C. R. MCKINNEY. 1951. Measurement of paleo- WANG, L. 2000. Isotopic signals in two morphotypes temperatures of the Upper Cretaceous of England, o f Globigerinoides ruber (white) from the South Denmark and the south-eastern United States. Bul- China Sea: Implications for monsoon climate letin of the Geological Society of America, change during the last glacial cycle. Palaeogeog- 62:399–426. raphy, Palaeoclimatology, Palaeoecology, VEIZER, J., D. ALA, K. AZMY, P. BRUCKSCHEN, 161:381–394. D. BUHL, F. BRUHN, G. A. F. CARDEN, A. DI- WEFER, G., AND W. H. BERGER. 1991. Isotope pale- ENER, S. EBNETH, Y. GODDERIS, T. JASPER, ontology: Growth and composition of extant cal- C. KORTE, F. PAWELLEK, O. G. PODLAHA, careous species. Marine Geology, 100:207–248.
37 THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18
WESTERHOLD, T., T. BICKERT, AND U. ROHL. tions of the Royal Society A, 365:1829–1842. 1995. Middle to late Miocene oxygen isotope stra- ZACHOS, J. C., J. R. BREZA, AND S. W. WISE. 1992. tigraphy of ODP site 1085 (SE Atlantic): new con- Early Oligocene ice sheet expansion on Antarc- straints on Miocene climate variability and sea- tica: Stable isotope and sedimentological evidence level fluctuations. Palaeogeography, Palaeoclima- from Kerguelen Plateau, southern Indian Ocean. tology, Palaeoecology, 217:205–222. Geology, 20:569–573. WILLIAMS, D. F., A. W. H. BÉ, AND R. G. FAIR- ZACHOS, J. C., G. R. DICKENS, AND R. E. ZEEBE. BANKS. 1981. Seasonal stable isotope variations 2008. An early Cenozoic perspective on green- in living planktonic foraminifera from Bermuda house warming and carbon cycle dynamics. Na- plankton tows. Palaeogeography, Palaeoclimatol- ture, 451:279–283. ogy, Palaeoecology, 33:71–102. ZACHOS, J. C., M. PAGANI, L. SLOAN, E. THO- WILLIAMS, M., A. M. HAYWOOD, S. P. TAYLOR, P. MAS, AND K. BILLUPS. 2001. Trends, rhythms, J. VALDES, B. W. SELLWOOD, AND C.-D. HIL- and aberrations in global climate 65 Ma to present. LENBRAND. 2004. Evaluating the efficacy of Science, 292:683–693. planktonic forminifer calcite δ18O data for sea ZACHOS, J. C., S. SCHOUTEN, S. BOHATY, T. surface temperature reconstruction for the Late QUATTLEBAUM, A. SLUIJS, H. BRINKHUIS, Miocene. Geobios, 38:843–863. S. J. GIBBS, AND T. J. BRALOWER. 2006. Ex- WILLIAMS, M., A. M. HAYWOOD, M. VAUTRAV- treme warming of mid-latitude coastal ocean dur- ERS, B. W. SELLWOOD, C.-D. HILLBRAND, I. ing the Paleocene–Eocene Thermal maximum: P. A. WILKINSON, AND C. G. MILLER. 2007. Inferences from TEX86 and isotopic data. Geol- Relative effect of taphonomy on calcification tem- ogy, 34:737–740. perature estimates from fossil planktonic foramini- ZACHOS, J. C., L. D. STOTT, AND K. C. LOHMANN. fera. Geobios, 40:861–874. 1994. Evolution of early Cenozoic marine tem- WILSON, D. S., S. S. R. JAMIESON, P. J. BARRETT, peratures. Paleoceanography, 9:353–387. G. LEITCHENKOV, K. GÖHL, AND R. D. LAR- ZEEBE, R. E. 1999. An explanation of the effect of TER. 2012. Antarctic topography at the Eoce- seawater carbonate concentration on foraminiferal ne–Oligocene boundary. Palaeogeography, Pa- oxygen isotopes. Geochemica et Cosmochimica laeoclimatology, Palaeoecology, 335–336:24–34. Acta, 63:2001–2007. WILSON, P. A., AND R. D. NORRIS. 2001. Warm ZEEBE, R. E. 2001. Seawater pH and isotopic paleo- tropical ocean surface and global anoxia during temperatures of Cretaceous oceans. Palaeogeogra- the mid-Cretaceous period. Nature, 412:425–429. phy, Palaeoclimatology, Palaeoecology, 170:49– WILSON, P. A., R. D. NORRIS, AND M. J. COOPER. 57. 2002. Testing the Cretaceous greenhouse hypothe- ZEEBE, R. E. 2012. History of seawater carbonate sis using glassy foraminiferal calcite from the core chemistry, atmospheric CO2, and ocean acidifica- of the Turnoian tropics on Demerara Rise. Geol- tion. Annual Reviews of Earth and Planetary Sci- ogy, 30:607–610. ences, 40:141–165. WILSON, P. A., AND B. N. OPDYKE. 1996. Equatorial ZEEBE, R. E., J. BIJMA, B. HÖNISCH, A. SANYAL, sea-surface temperatures for the Maastrichtian H. J. SPERO, AND D. A. WOLF-GLADROW. revealed through remarkable preservation of meta- 2008. Vital effects and beyond: a modelling per- stable carbonate. Geology, 24:555–558. spective on developing paleoceanographical proxy WILSON-FINELLI, A., G. T. CHANDLER, AND H. J. relationships in foraminifera, p. 45–58. In W. E. N. SPERO. 1998. Stable isotope behavior in paleo- AUSTIN, AND R. H. JAMES (eds.). Biogeochemical ceanographically important benthic foraminifera; Controls on Paleoceanographic Proxies. Geologi- results from microcosm culture experiments. Jour- cal Society of London, Special Publications, No. nal of Foraminiferal Research, 28:312–320. 303. WINTERER, E. L. 2000. Scientific ocean drilling, ZIVERI, P., S. THOMAS, I. PROBERT, M. GEISEN, from AMSOC to COMPOST, p. 117–127. In 50 AND G. LANGER. 2012. A universal carbonate ion years of Ocean Discovery, National Academy effect on stable oxygen isotope ratios in unicellu- Press, Washington, DC. lar planktonic calcifying organisms. Biogeosci- ZACHOS, J. C., M. A. ARTHUR, T. J. BRALOWER, ences, 9:1025–1032. AND H. J. SPERO. 2002. Palaeoclimatology (Communication arising): tropical temperatures in greenhouse episodes. Nature, 419:897–898. ZACHOS, J. C., S. M. BOHATY, C. M. JOHN, H. MCCARREN, D. C. KELLY, AND T. NIELSEN. 2007. The Palaeocene–Eocene carbon isotope ex- cursion: constraints from individual shell plank- tonic foraminifer records. Philosophical Transac-
38