oxygen isotope evolution of phanerozoic seawater

Palaeogeography, Palaeoclimatology, Palaeoecology 132 (1997) 159–172

Review Paper

Oxygen isotope evolution of Phanerozoic seawater

Jan Veizer a,b, Peter Bruckschen a, Frank Pawellek a, Andreas Diener a,Olaf G. Podlaha a, Giles A. F. Carden a, Torsten Jasper a, Christoph Korte a,

Harald Strauss a, Karem Azmy b, Davin Ala ba Institut fur Geologie, Ruhr-Universitat, D-44780 Bochum, Germany

b Ottawa-Carleton Geoscience Center, University of Ottawa, Ottawa, Ont. K1N 6N5, Canada

Received 29 March 1995; accepted 19 March 1997

Abstract

A compilation of over 2000 measurements of 18O and 13C on Phanerozoic low-Mg calcite shells, such as brachiopods,belemnites and oysters, delineates secular 18O/16O and 13C/12C variations that are similar to those previously describedfor whole rocks. The trend for the d18O suggests about ~5±2‰ enrichment from the Cambrian to today. In contrast,the d13C rise during the Paleozoic is followed by its decline in the Mesozoic and Cenozoic. Optical (textural ) andchemical criteria suggest that the interior ‘‘secondary’’ layer of the brachiopod shells, the material that carries thesesignals, is well preserved in many samples and the extracted secular isotopic trends are therefore a primary feature ofthe geologic record. The similarity of the d18O/d13C isotope patterns in ancient and modern brachiopods also supportssuch an interpretation. In our view, the 18O enrichment in progressively younger samples is principally, although notexclusively, a reflection of the evolving 18O/16O composition of seawater. If so, a delineation of this trend mayultimately result in development of a valuable paleoclimatic and paleoceanographic tracer for the Phanerozoic. © 1997Elsevier Science B.V.

Keywords: oxygen; isotopes; seawater; Phanerozoic

1. Introduction retain a record on time scales of up to 109 years,while the vestiges of ‘‘short-lived’’ entities, such asthe hydrosphere and atmosphere, survive for onlyThe dynamics of our planet, as of any natural

system, is dominated by cyclic processes that oper- a maximum of ~104–105 years (e.g., air bubblesin the ice caps). In order to decipher the propertiesate on a hierarchy of temporal and spatial scales.

Evolutionary phenomena and ‘‘events’’ are only of the ancient hydrosphere and atmosphere, wetherefore have to rely on derivative, or proxy,superimposed on this background dynamics. It is

a property of any cyclic system that its past record signals that such ‘‘short-lived’’ systems inscribe onthe ‘‘long-lived’’ ones. Isotopes of oxygen, carbon,is obliterated by a continuous process of

generation/destruction (recycling), while the quan- strontium, sulfur and other elements, measured inmineral phases that precipitated from seawater,titative record is retained only for a single, extant,

cycle (cf. Veizer, 1988). As a consequence, the are among the best proxies available for deci-phering the properties of the ancient ocean–‘‘long-lived’’ populations, such as the solid Earth,

0031-0182/97/$17.00 © 1997 Elsevier Science B.V. All rights reserved.PII S0031-0182 ( 97 ) 00052-7

160 J. Veizer et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 132 (1997) 159–172

atmosphere system; a precondition for understand- geological time scales. It is for this reason that wehope to reopen the discussion of the matter,ing the impact of issues epitomized by the concept

of ‘‘Global Change’’. regardless of its outcome.One of the fundamental issues of geological and

biological evolution on this planet, including thepossible impact of anthropogenic phenomena, is 2. Stability of marine precipitatesits temperature history; a variable for whichoxygen isotopes may serve as an invaluable proxy. It is an unfortunate reality that mineral phases

precipitated from seawater are only metastableThis has been demonstrated already in the pioneer-ing work of Emiliani (1955) that documented for precursors of the components that eventually con-

struct limestones and dolostones, cherts and phos-the first time the existence of glacial–interglacialoscillations in 18O/16O ratios of Quaternary marine phorites. The post-depositional transformation of

aragonite and high-Mg calcite into calcite or dolo-sediments (e.g., Hays et al., 1976). These havesubsequently been reproduced in the much dis- mite, opal A into quartz, and francolite into fluor-

apatite clearly results in modification of thecussed ice cores (e.g., Jouzel et al., 1993). Anextension of this oxygen isotope record into the inherited isotope signal, usually leading to lighter

d18O and d13C values (e.g., Lohmann, 1988). Ifdimmer geological past is, however, fraught withuncertainties and disagreements. For the Cenozoic, the 18O depletion in old (bio)chemical sediments

were solely a result of post-depositional phen-the consensus still exists that the d18O of pelagicand benthic foraminifera records a progressive omena, the trend would only reflect an advancing

equilibration with meteoric waters at ever increas-cooling of the deep ocean waters (Savin and Yeh,1991), probably reflecting changes to the ‘‘con- ing burial temperatures; the degree of equilibration

being statistically proportional to the age of theveyor belt’’ circulation pattern that controls theenergy balance of the oceans (cf. Broecker and rocks (Degens and Epstein, 1962; Killingley,

1993). Let us therefore first consider the merits ofPeng, 1982). This consensus starts to break up forthe Mesozoic and disintegrates when dealing with this alternative.

Indisputably, post-depositional stabilization ofthe Paleozoic and older time intervals.The general observation that the d18O of pre- rock components usually leads to alteration of the

geochemical signal. The question is only by howMesozoic marine precipitates becomes pro-gressively depleted in 18O with increasing age, much. It is frequently argued that compared to

calcite, quartz and apatite are the more stableestablished already in the pioneering work ofBaertschi (1957) and Clayton and Degens (1959), phases and they should therefore better retain

vestiges of marine isotopic signature (e.g., Knauthis not disputed. Subsequent work convincinglyconfirmed this to be the case for carbonates (e.g., and Epstein, 1976; Kolodny and Epstein, 1976).

Yet, it is not the relative stability of the finalSchidlowski et al., 1975; Veizer and Hoefs, 1976),cherts (e.g., Perry and Tan, 1972; Knauth and product, but the extent and duration of the trans-

formation pathway that are the decisive factorsEpstein, 1976) and phosphorites (e.g., Longinelliand Nuti, 1968). The prime issue arising from controlling the degree of modification of the origi-

nal signal. The transformation of a carbonatethese experimental data is whether the trend is ofprimary or secondary (post-depositional ) origin. protolith into limestone, an Ostwald ripening pro-

cess by dissolution/reprecipitation is, geologicallyA resolution of this fundamental question is aprecondition for any discussion of possible caus- speaking, a relatively rapid venture. It is accom-

plished mostly during early diagenesis, when nei-ative factors, and it is this issue that is at the coreof the ‘‘irreconcilable’’ disagreements in the ther the temperature nor the isotopic composition

of the diagenetic waters are radically different fromresearch community (cf. Land, 1995; Veizer, 1995).Unfortunately, perpetuation of the stalemate may that of seawater (cf. Choquette and James, 1990;

James and Choquette, 1990). In contrast, stabiliza-be depriving us of one of the most powerful toolsfor studies of the terrestrial exogenic system on tion of cherts and phosphorites is a protracted

161J. Veizer et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 132 (1997) 159–172

affair, fully accomplished only at pressures and The alternative technique employed in this paperutilizes a marine phase that is stable in diagenetictemperatures of deep burial (Hesse, 1990a,b;

McArthur, 1985). Once stabilized, the rocks, environments and can therefore retain the originalisotope signal. Such a phase, the low-Mg calcite,regardless of lithology, stay relatively ‘‘tight’’ to

subsequent superimposed effects. Since all these is secreted only via a biological intermediate, suchas the foraminifers, some molluscs (e.g., oysters),rocks passed through a diagenetic stage, they are

all shifted to lighter d18O values, cherts more so belemnites and brachiopods (cf. Morrison andBrand, 1986). The utility of low-Mg calcitic fora-than carbonates. Yet, this diagenetic resetting

could only explain the overall downward shift of minifers as recorders of Cenozoic paleoceano-graphic signals has been amply demonstrated inthe trends, but not their continuous 18O depletions

with increasing age. We do not wish to imply that literature (cf. Williams et al., 1988) and does notrequire any specific justification here. Belemnitethe rocks are absolutely stable. We only argue that

the sum of post-stabilization effects accounts for rostra and oysters, on the other hand, maybe somewhat porous (~3–10%: Veizer, 1974;only ‘‘fixed’’ proportion of the overall isotopic

trends. Podlaha, 1995) and this porosity may be filled bysecondary calcite. Their isotope signal may or mayUnfortunately, the uncertainties inherent in esti-

mates of the magnitude of post-depositional iso- not therefore be slightly altered (Veizer and Fritz,1976; Jones, 1992). Nevertheless, these latter fossiltope shift of bulk rocks are such that this approach

can easily be challenged. It is therefore not likely groups are of particular importance for theMesozoic only, the time interval where the d18Oto lead to definitive answers regarding the isotopic

composition of ancient seawater. We must there- of whole rocks did not appear to have beendrastically different from their Cenozoic counter-fore search for alternative approaches. One such

technique was developed by the Michigan group parts (Veizer and Hoefs, 1976). Such considera-tions will therefore be of importance forof Lohmann (e.g., Carpenter et al., 1991) and it is

based on selected components of limestones, such understanding the higher-order structure in thePhanerozoic d18O record, but are not decisive foras the early diagenetic marine cements. A sequence

of d18O and d13C measurements on coeval cement delineation of the overall trend. In contrast to theCenozoic and Mesozoic, the Paleozoic whole-rockgenerations yields a plot that apparently converges

to isotopic composition of the original marine samples show a clear depletion in 18O (Veizer andHoefs, 1976). It is therefore this time interval thatphase. It is not entirely clear why this should be

so, since the cements are today only paramorphic is of critical importance for testing of the d18Osecular trend and the only ubiquitous low-Mgcalcites after aragonite or high-Mg calcite precur-

sors. The likely explanation is that the cements calcite phase present is that of the brachiopods.The question of the structure and mineralogy ofcontain differing proportions of microdomains that

are still preserved as aragonite or high-Mg calcite. the brachiopod shell and its ability to retain itsoriginal marine attributes is thus of primeAlternatively, diagenetic stabilization of microdo-

mains may have been accomplished at variable importance.water/rock ratios. These microdomains then yieldthe projection trends that point to the compositionof the precursor. Whatever the explanation, the 3. Shell structure of brachiopods and their

preservation potentialapproach appears to yield results (cf. Lohmannand Walker, 1989) that are in accord with analternative technique that is described in the subse- Brachiopods are marine organisms with shells

of two valves, usually fixed on the sea-floor byquent text. The Michigan approach may be ofparticular importance for the Precambrian carbon- means of a stalk or pedicle. About 3000 fossil

genera, but only some 100 recent species are knownate rocks that are devoid of skeletal components.It also may serve as a cross-check for the (Clarkson, 1993). The shell of articulate brachio-

pods is usually composed of low-Mg calcite, butPhanerozoic results discussed below.


species of articulate brachiopods have hollowspines that perforate the entire shell.

Under a standard optical microscope, thestudied brachiopod shells appear to have been verywell preserved (Fig. 2a–c). However, cathodolum-inescence studies revealed some shell domains withMn-activated red-yellow-orange luminescence sim-ilar to that observed in the matrix (Fig. 2d–e).Mn-rich solutions apparently penetrated the shellsalong the calcite fibres as well as along the crosscut-ting microcracks (cf. also Veizer et al., 1986).

Fig. 1. Schematic shell structure of brachiopods. Nevertheless, these luminescent ‘‘streaks’’, particu-larly within the interior secondary layer, usuallyform only subordinate portions of the shell,some inarticulate forms also have a phosphaticwith the bulk remaining non-luminescent.shell. Recent brachiopod shells contain 0.5–7Luminescence, when present, is mostly confined tomole% MgCO3 (Lowenstam, 1961; Morrisonthe outer primary layer, where it follows theand Brand, 1986) and ancient counterparts weredamage to the fine fibres caused by chemical,likely of comparable mineralogy as is evident fromphysical or biological (boring) agents. ThisMg-poor shell fragments enclosed in authigenicdamage to the extracellularly secreted primaryquartz (Richter, 1972) or coexisting withlayer can commence on living brachiopods andaragonitic rugose corals among the Carboniferousextend, or be confined, to their post-mortem stage.Kendrick fauna (Brand, 1981). The inorganic partDrilling of brachiopod shells usually crosses thisof the shell for articulate brachiopods consists ofprimary layer and may result in sample contamina-two layers, the outer ‘‘primary’’ and the innertion by secondary calcite and/or matrix (cf. Diener‘‘secondary’’. The primary layer is finely granular,et al., 1996). Similarly, contamination problemshaving distinct lineation that is oriented perpendic-might arise when using collection material fromular to the shell surface. In contrast, the secondarymuseums. In fresh samples, a portion of the entirelayer consists of elongated calcitic fibresprimary layer frequently adheres to the enclosing(MacKinnon, 1974). The brachiopod shells are ofrock, while the secondary layer spalls along itsthree types (Fig. 1). The impunctate and pseudo-contact with the primary layer. Collection speci-punctate shells have no perforations and frequentlymens, on the other hand, are usually weatheredthe latter ones also have irregularly developedout fossils that retain both layers (Fig. 2h) andsecondary layers. The endopunctate shells, on thehave therefore a higher probability of contamina-other hand, are perforated — in 0.05—0.1-mmtion. Our preparation technique is based on smash-intervals — by ‘‘channels’’ that are oriented per-ing 1–2 blocks into 2–3-mm-size pieces, followedpendicular to the shell surface (Fuchtbauer, 1988).by cleaning in 250 ml of distilled water, decanta-These channels can be filled by secondary mineralstion of the fine particulates, and drying for 3 h atand such samples are thus susceptible to post-

depositional contamination. In addition, some 50°C. The cleanest shell splinters of the secondary

Fig. 2. Optical, cathodoluminescence and SEM pictures of brachiopod shells. Examples of textural preservation of brachiopod shells.a–c. Cross-sections of brachiopod shells under optical microscope. Explanation: p=primary layer; s=secondary layer.d–f. The same samples under cathodoluminescence. Note the Mn-activated orange luminescence, typical of altered parts, that is

concentrated in the primary layers and at the contact of the primary and secondary layer. The latter, except for micro-fractures,remains non-luminescent.

g. SEM image of slightly bent calcite lamellae from the secondary layer of a brachiopod shell.h. SEM image of fine-grained primary layer. The sample was weathered from a marly layer and contains the attached secondary

layer underneath.


layer, clearly recognizable by their structure of fine brachiopod shells is comparable to that ofCenozoic foraminifera. The relative thickness offibres, are then hand-picked under binocular

microscope. All splinters with recognizable micro- the brachiopod shell, if compared to (pelagic)foraminifera, may be an additional advantage infissures are rejected. The selected shell fragments

are ultrasonically cleansed several times in distilled terms of its greater resistance to alteration and interms of the ease of sampling for experimentalwater until its cloudiness disappears. These

samples are subsequently utilized for chemical and studies. This is not to claim that every brachiopodshell is well preserved, and all available opticalSEM studies.

The optical criteria show that the secondary and chemical criteria must be utilized to select thebest samples. Nonetheless, an ad hoc belief thatlayers of many Paleozoic brachiopods are well

(Fig. 2g) to exceptionally well (fig. 5 in Qing and tacitly assumes that all Paleozoic brachiopods,because they have ‘‘depleted’’ d18O, are invariablyVeizer, 1994; fig. 2 in Wadleigh and Veizer, 1992)

preserved. This is confirmed also by their chemis- altered is clearly untenable and internally inconsis-tent (Veizer, 1995). In our view, the onus lies withtry, with trace elements such as Sr, Fe and Mn,

distributed quite uniformly at near equilibrium the proponents of this alternative to prove that allthe above observations and techniques are invalidvalues (Fig. 3; Bruckschen et al., 1995). These

optical and chemical criteria document that the and, if so, explain why then the same criteria areregarded as valid for the Cenozoic foraminifera.preservation of the secondary layer of Paleozoic

Taking into account the above commentary, weconsider the isotopic record based on low-Mgcalcitic shell components to be, in its major part,a primary feature of the geologic record.

4. Isotopic composition of Phanerozoic low-Mgcalcitic components

The existing d18O and d13C measurements onlow-Mg calcitic skeletons of Phanerozoic brachio-pods and Mesozoic belemnites and oysters aresummarized in Figs. 4 and 5, respectively. Thesefigures are based on 2435 (Fig. 4) or 2314 (Fig. 5)measurements that include 460 as yet unpublishedresults from our database. The clear feature of thiscompilation is a depletion of ~5±2‰ in 18Othroughout the Phanerozoic (Fig. 4). The overalltrend, however, is not a simple linear feature.Instead, it appears to be a band of ~4–6‰ width,with indications of second-order oscillations super-imposed on the first-order trend. A similar bandis evident also in the d13C database, except that inthis case the possible second-order oscillations aresuperimposed on a hump that peaks in the late

Fig. 3. PIXE line scan through a Lower Carboniferous brachio- Paleozoic (Fig. 5). This hump mimics thepod shell. Step size 100 nm. The trace-element distribution mir- Phanerozoic trend documented by Veizer et al.rors preservational differences between pristine shell material

(1980) and Lindh (1983) from the whole-rock(low Mn, Fe; high Sr) and the diagenetically altered rock matrixdata.and cement. Detection limits for Mn are 20–30 ppm. From

Bruckschen et al. (1995). Some of the observed scatter, particularly the


Fig. 4. Oxygen isotope composition of Phanerozoic low-Mg calcitic fossil shells (N=2435). The data for modern brachiopods arefrom Carpenter and Lohmann (1995). Aragonitic macrofossils after Hudson and Anderson (1989). Brachiopods, belemnites andoysters from Compston (1960), Lowenstam (1961), Veizer and Compston (1974), Veizer and Hoefs (1976), Popp et al. (1986a,b),Veizer et al. (1986), Adlis et al. (1988), Brand (1989a,b,c), Carpenter and Lohmann (1989), Delaney et al. (1989), Rush and Chafetz(1990), Gao and Land (1991), Grossman et al. (1991, 1993), Middleton et al. (1991), Grusczynsky et al. (1992), Jones (1992),Wadleigh and Veizer (1992), Brand and Legrand-Blain (1993), Gao (1993), Lavoie (1993), Brenchley et al. (1994, 1995), Grossman(1994), Qing and Veizer (1994), and Ludvigson et al. (1997). 476 samples are our unpublished measurements as of 1994. Thesummary of all experimental data of Ottawa and Bochum groups, up to 1992, is available from the senior author as a non-refereedcompilation edited by Veizer (1994). Please note that the figure, in contrast to the one published by Veizer (1995), contains also thedata from Brenchley et al. (1994, 1995) and Ludvigson et al. (1997) and more of our unpublished measurements.

Fig. 5. Carbon isotope composition of Phanerozoic low-Mg calcitic fossil shells (N=2314). Sources of data as in Fig. 4.

outliers with the highly depleted d18O and d13C subjected to the above stringent selection process.Furthermore, some samples are bulk shells rathervalues (Figs. 4 and 5), may be due to post-deposi-

tional overprinting, since not all samples were than the secondary layers only. At this stage, we


shall concentrate on the main band, leaving the and/or rapid precipitation of calcium carbonate(Carpenter and Lohmann, 1995). Metabolic activ-issue of outliers to be resolved by future studies.

The main trend scatter of ~4–6‰ is comparable ity, particularly during secretion of the primarylayer, appears to be therefore the dominant con-to that observed in modern brachiopods (Fig. 4).

This is a reflection of variability in the d18O of trolling variable that causes the observed 18O and13C depletions, with environmental factors (e.g.,ambient water, in temperature and salinity of

brachiopod habitats (cf. Bates, 1994), and in depth, salinity) possibly playing a contributoryrole. The Paleozoic counterparts yield an almostthe later discussed vital fractionation effects

(Carpenter and Lohmann, 1995). Modern brachio- identical picture, except that the pattern is shiftedto heavier d13C and lighter d18O values (Fig. 7).pods, in contrast to most of their Paleozoic coun-

terparts, inhabit not only warm shallow seas, but In ancient carbonates, diagonal d18O/d13C patternshave usually been attributed to post-depositionalalso the temperate and deeper water habitats

(Clarkson, 1993). The expected range of their alteration, because similar trends can be generatedby mixing of juxtaposed microscopic domains ofd18O values should therefore be considerably larger

than that of their fossil counterparts. Yet, if only solid phases with two end-member compositions(e.g., primary and diagenetic calcite) or by mixingwarm water brachiopods are taken into account,

the range of modern d18O values contracts only of fluids (e.g., sea and CO2-charged meteoricwaters) in the course of precipitation of diageneticmarginally, by less than 1‰.

The overall variability and covariance of d18O calcite (cf. Lohmann, 1988; James and Choquette,1990). We do accept that some 18O- and 13C-and d13C in modern brachiopods is well illustrated

in Fig. 6. This covariance appears to be principally depleted samples, particularly those that werenot screened by the techniques described in thea result of kinetic fractionation of oxygen and

carbon isotopes during hydroxylation of CO2 earlier sections of this paper, may reflect somepost-depositional resetting of the isotope signal,due for example to inclusion of a diagenetic solidphase (e.g., the altered primary layer) into themeasured aliquot. Nonetheless, the bulk of the

Fig. 6. Scatter plot of d18O vs. d13C for recent brachiopods.Based on data in Carpenter and Lohmann (1995). Full circles=secondary layer; empty circles=primary layer. The two interior

Fig. 7. Scatter plot of d18O vs. d13C for Paleozoic brachiopodsfields, enclosed by the heavy solid and the dotted line, are forthe modern temperate brachiopods from Tasmania and from (cf. Figs. 4 and 5). The fields for recent brachiopods from Fig. 6.

The dashed line represents a linear least-squares fit regressionother non-tropical brachiopods, respectively (fig. 3 in Rao,1994). for the ancient brachiopod samples.


samples were properly screened, are well preserved, multiple genera of brachiopods for a given strati-graphic unit gave comparable d18O values (Veizerand still display the covariant trend, as do their

modern counterparts. The latter are extant species, et al., 1986; Brand, 1989b,c; Bates and Brand,1991; Grossman et al., 1991; Wadleigh and Veizer,i.e. shells from living or recently dead specimens,

and the observed d18O/d13C covariance cannot be 1992; Qing and Veizer, 1994), suggesting that anyacalcite–water biological disequilibrium fractionation,therefore ascribed to post-depositional phen-

omena. We conclude, therefore, that the observed if present, played only a subordinate role in thegeneration of the Phanerozoic signal. Futuresecular trends for both, oxygen and carbon iso-

topes, are primary features of the geologic record. efforts may, in analogy to foraminifera, demon-strate species- or genera-specific acalcite–water, but ifThere is a general consensus that the d13C trend

reflects the isotopic evolution of the Phanerozoic so, this will be pertinent only to elucidation ofa higher-order structure within the overallseawater and here we propose that this is the case

also for the d18O. Modern brachiopods, par- Phanerozoic band. As of this date, the existingdata for modern as well as fossil specimens areticularly those from low-latitude shelf habitats,

precipitate their calcite in approximate iso- consistent with the proposition that the brachio-pods precipitated the secondary layer of their shelltopic equilibrium with ambient water (Fig. 8)

(Lowenstam, 1961). Brachiopods from latitudes in approximate isotopic equilibrium with ambientwater.higher than 40°N/S may or may not record disequi-

librium in excess of 1‰, particularly as a result ofthe earlier discussed ‘‘vital’’ effect. Still, consideringthe uncertainties as to the ambient conditions at 5. Causes of the observed secular trendthe time of calcification, and the fact that thesehigh-latitude types are not the analogues to the Acceptance of the Phanerozoic d18O trend as a

primary feature of the geologic record raises thePaleozoic counterparts, the observation is notdirectly relevant to our issue. For the Paleozoic, question of possible causes. Four alternatives have

been advanced as possible explanations and dis-cussed in many publications that cannot bereviewed here. We shall list only the authors ofthe original idea, as far as we could ascertain them,and some representative subsequent publications.The proposed explanations were the following:(1) the fractionation factor acalcite–water may have

increased in the course of geologic history. Anexample could be a succession of fossil phylawith differing a (vital effect; McConnaughey,1989a,b), that supplant each other in thegeological record;

(2) the average temperature of the ocean waterdeclined in the course of geologic history(Knauth and Epstein, 1976; Kolodny andEpstein, 1976; Karhu and Epstein, 1986);

(3) the ocean water became progressively enrichedin 18O over geological time ( Weber, 1955a,b;

Fig. 8. Calculated upper an lower equilibrium values and mea- Perry et al., 1978; Walker and Lohmann,sured d18O for the shells of recent brachiopods. Modified from 1989); andCarpenter and Lohmann (1995). In addition, Rao (1994)

(4) the early ocean, in contrast to its modernclaimed that temperate recent brachiopods also precipitate theircounterpart, was salinity rather than thermallylow-Mg calcite shell in isotopic equilibrium with ambient

seawater. stratified. In this case, it is hypothesized that


Fig. 9. The required d18O of seawater during the Phanerozoic assuming a given temperature (17–38°C ) and accepting the averagetrend in Fig. 3 as representative for coeval calcite. At this stage, it is not clear whether the average or the haviest brachiopod samplesare the best approximations to the ‘‘equilibrium’’ values. Future detailed studies of secondary layers of brachiopod shells, centeredon specific locations, may help to resolve the issue. For the same locality and habitat, the average may turn out to be the bestapproximation for a clump of data points. On the other hand, if the data show a definite covariance of d18O/d13C values, notattributable to alteration, the heaviest measurements may yield the best approximation.

its deep waters were generated from brines arguments that rule out changes in temperature ofseawater (alternative 2), or in the salinity structureformed by extensive evaporation at low lati-

tudes (Wilde and Berry, 1984), leaving the of the ocean (alternative 4), as the principal andsole causes. We are therefore inclined to pursuewaters above the halocline (thermocline),

including the shelf seas, depleted in 18O (Brass the possibility of changing the 18O/16O of seawater(alternative 3), despite the objections raised by theet al., 1982; Railsback, 1990; Grusczynsky

et al., 1992). ophiolite studies (e.g., Gregory and Taylor, 1981;Lecuyer and Fourcade, 1991) and by model consid-

The fact that the Phanerozoic d18O trend is erations (e.g., Muehlenbachs and Clayton, 1976;Gregory, 1991). As for the latter, we would likeevident regardless whether the whole rocks, abiotic

marine cements, several phyla or only the brachio- to point out that alternative models ( Walker andLohmann, 1989) do permit a unidirectional ratepods are considered, argues against the first of the

above outlined alternatives. The pros and cons of of change of ~1‰/108 years indicated by theaverage slope in the Phanerozoic d18O trendthe alternatives 2 to 4 have been discussed in detail

in Veizer et al. (1986), Brand (1989b,c), Qing and (Figs. 4 and 9). On the other hand, the existenceof a higher-order structure, such as the one deline-Veizer (1994), and Corfield (1995). Due to space

constraints, and because we have nothing decisive ated by the Cenozoic aragonitic components(Fig. 4) or by the Carboniferous brachiopodsto add to this discussion, we shall not repeat it

here. At this stage, the main point we would like (Bruckschen and Veizer, 1997-this issue), requiresalternative explanations. This is because theto make is that we regard the observed overall

trend to be a primary feature of the geologic required rate of change in d18O appears to be toofast for any existing models that are based onrecord, whatever its causes. The trend itself may

contain components of all the above factors. balances of rock–water interactions. The databasethat is inevitable for testing such propositions is,Nevertheless, we consider as compelling geological


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