Geochemistry of brachiopods: Oxygen and carbon isotopic records of Paleozoic oceans

0016-7037/86/%3.00+ .I0

Geochimrca 1 Cosmochimica Acta Vol. 50, pp. 1679-1696 0 Pergamon Journals Ltd. 1986. Printed in U.S.A.

Geochemistry of brachiopods: Oxygen and carbon isotopic records of Paleozoic oceans* JANVEIZER’,PETERFRITZ*and BRIANJONES~ ‘Derry Laboratory, Department of Geology, University of Ottawa, Ottawa, Ontario, KlN 6N5, Canada ‘Department of Earth Sciences, University of Waterloo, Waterloo, Ontario, N2L 3G1, Canada ‘Department of Geology, University of Alberta, Edmonton, Alberta, T6G 2E3, Canada (Received October 11, 1985; acceptedin revisedform May 9, 1986) Abstract-Combined trace element and isotope studies of 319 brachiopods, covering the Ordovician to Permian time span, show that 6r3C and 6180 in well preserved specimens varied during the Paleozoic. The overall 6r3Csecular trend is in accord with the previously published observations, but its details are obscured by vital isotopic fractionation effects at generic level. Nonetheless, the results suggest that the negative correlation between marine B’3Ccarbonate and G”S,,,rr~~deteriorates at time scales of ~10~ years, due to the long residence time, and thus slow response, of SOf- in the ocean. For oxygen isotopes, all Devonian and older specimens have aI80 of =Z-4L, while the well preserved Permian samples have near-present day b180 of about - 1% (PDB). This isotopic dichotomy is probably not due to post-depositional phenomena, salinity,

or biogenic fractionation effects.This leaves open the perennial arguments for a change in ‘8O/‘6Oof sea water versus warmer ancient oceans. The present data are difficult to explain solely by the temperature alternative. The coincidence of the proposed shift in aI80 with the large Late Paleozoic changes in marine 87Sr/*6Sr,‘3C/‘ZC,?S/“S, and “sea level stands” argues for a tectonic cause and for a change in ‘8O/‘6O of sea water, although such explanation is difficult to reconcile with global balance considerations and with isotopic patterns observed in alteration products of ancient basalts and ophiolites. Whatever the precise cause, or combination of causes, the implications for tectonism and/or paleoclimatology are of first order

significance. INTRODUCTION

THE TREND OF decreasing 6’*0 values of marine chemical sediments with increasing age has been well documented. This trend has been observed for carbonates (BAERTSCHI, 1957; CLAYTON and DEGENS, 1959; DEGENSand EPSTEIN,1962; KEITH and WEBER, 1964; WEBER, 1965a,b; DONTSOVA et al., 1972; BAUSCH and HOEFS, 1972; PERRY and TAN, 1972; SCHIDLOWSKIet al., 1975, 1983; VEIZER and HOEFS, 1976), cherts (DEGENS and EPSTEIN, 1962; PERRY, 1967; PERRY and TAN, 1972; KNAUTH and EPSTEIN, 1976; KNAUTH and LOWE, 1978), phosphorites (LONGINELLIand NUTI, 1968; SHEMESHef al., 1983; Luz et al., 1984), and glauconites (FRIEDRICHSEN,1984; KEPPENSand O’NEIL, 1985). The perennial question of the cause(s) of this isotopic shift is, however, still vigorously debated. Suggested causes include: (1) postdepositional equilibration with I80 depleted meteoric waters, the degree of equilibration increasing with increasing geologic age (DEGENS and EPSTEIN, 1962; KILLINGLEY,1983; and many of the above quoted authors), (2) higher temperatures of the earlier oceans, with the Archean and early Proterozoic water temperatures being in the range of &6O”C (KNAUTH and EPSTEIN, 1976; KOLODNYand EPSTEIN, 1976; KNAUTH and LOWE, 1978; SHEMESHet al., 1983; Luz et al., 1984), and (3) change in the oxygen isotopic composition of sea water over geologic time (WEBER, 1965a,b;

* Publication 05-86 of the Ottawa-Carleton Centre for Geoscience Studies.

PERRY, 1967; DONTSOVA, 1970; DONTSOVA et al., 1972; FRITZ, 1971; PERRY and TAN, 1972; PERRY et al., 1978). The trend towards lower 6’*0 values of chemical sediments, particularly carbonates, with increasing age is minor during the Cenozoic and Mesozoic, but is significant during the Paleozoic and the Precambrian (c$ VEIZER and HOEFS, 1976). For the Phanerozoic, there are indications of a major step in the 6180 age curve during the Late Paleozoic times. If post-depositional alteration were the only cause of the observed secular trend, this could be of some importance for understanding the diagenetic evolution of (bio)chemical sediments, but would not require any modifications in our present understanding of the hydrologic cycle. Development of appropriate criteria (e.g. LOWENSTAM, 196 1; VEIZER and FRITZ, 1976; VEIZER, 1983a) can aid in selection of the “least altered” samples and in utilization of such samples for classical paleothermometry (EPSTEIN et al., 195 1, 1953). Yet, it is curious that practically all published analyses of samples older than 300 million years (see the above references)-including texturally, mineralogically and chemically well preserved fossils from tight enclosing rocks-yield high isotopic “paleotemperatures”, suggesting that the isotopic data may reflect a primary rather than a secondary feature. If this is correct, the implications are of considerable geological impact, regardless whether the ultimate cause was high temperature, low 18O/‘6O ratios of seawater, or a combination of the two. If the early oceans were warmer, it would be necessary to identify the cause and to reconcile this with the existing sedimentological and paleoecological records. Alternatively, a change in the 1679

1680

J. Veizer, P. Fritz and B. Jones

oxygen isotopic composition of seawater would demand modification of our understanding of global balances for rock-water interactions. In contrast to the oxygen isotope record, secular oscillations in 613C of marine bicarbonate have been convincingly demonstrated on a variety of time scales (COMBTON, 1960; WEBER,1967; KROOPNEK et al., 1977; SHACKLETON, 1977, 1985; BERGER, 1977; FISCHERand ARTHUR, 1977; SCHOLLEand ARTHUR, 1980; VEIZER et al., 1980; LETOLLE and RENARD, 1980; ARTHUR, 1982; MAGARITZ et al., 1983; RENARD, 1985). In addition, on time scales of 10’ years, the S’3C,,,,,,,t, correlates negatively with 8?&,fale (VEIZER et al., 1980) and these two variables also correlate with 87Sr/86Sr-, and the Phanerozoic “sea level stands” (VEIZER et al., 1982, Table 1; HOLSER, 1984; VEIZER, 1985). To the best of our knowledge, no existing model linking the tectonic factors (*‘Sr/ *?Sr “sea level”) to the evolution of the biosphereatmosphere-hydrosphere-sediment system (613C,634S) expiains simultaneously all the observed signs of correlations (cf: VEIZER, 1985). A workable, geologically realistic, unified theory implied by the above relationships will have to await future developments. Nevertheless, the inverse S-C isotopic relationship is almost certainly a reflection of oxygen transfer, and thus of redox balance, between 0 and C exogenic cycles (GARRELSand PERRY, 1974; GARREL~and LERMAN, 1981; BERNERet al., 1983). The ditrerences in residence times of marine sulfate and bicarbonate (7.9 X lo6 vs 8 X lo4 years; HOLLAND, 1978, Table 5-l), however, mean that the degree of coupling of the S and C cycles in the open oceans should be diminished on time scales of
been routinely utilized for Cenozoic and Mesozoic paleoceanographic studies. For the Paleozoic, brachiopods, because of their abundance, widespread stratigraphic distribution, and low-Mg calcitic mineralogy of their shell may have been a suitable phylum. Virtually all Recent and fossil brachiopods have a multi(usually two) layered valve composed entirely of lowMg calcite (JOHNSON, 195 1; LOWENSTAM,196 I ; WILLIAMS, 1968, 1971; BATHURST, 1975). The only possible exception may have been the family Trimerellidae, where the shell could have been originally aragonitic (JAANUSSON, 1966). According to LOWENSTAM( 196 I), extant articulate brachiopods secrete their shells reasonably close to oxygen isotopic equilibrium with ambient water, but more definite data are required to exclude the possibility of minor biogenic fractionation. For carbon, recent work of WEFER ( 1983) demonstrated that living articulate brachiopods fractionate carbon isotopes at the generic level. This observation complicates the possibility of resolving the 613Csecular variations on 106years time scale, since the knowledge of the degrees of such fractionation for specific genera must be known. Alternatively, a single genus must be utilized for studies of this type. Despite the relative post&positional stability of lowMg calcite, the preservation of the primary isotopic signal should not be d priori assumed. Variable degrees of recrystallization and iilling of primary and secondary void spaces by diagenetic calcites may have occurred. Utilizing optical (microscopic, staining, SEM) and trace element techniques, AL-AASM and VEIZER (1982) showed that the presently discussed Ordovician brachiopods from Anticosti Island (see Appendix 1) contained ~20% diagenetic calcite in their shells. In the present study, we shall employ a similar stable isotope, trace element, and textural approach to evaluate the degree of secondary alteration. The theory and applications of the technique, originally developed for brachiopods as well as belemnites (LOWENSTAIW,1961; VEIZER, 1974; VEIZER and FRITZ, 1976), are summarized in detail in VEIZER (1983a,b). In short, alteration (due to a multitude of mutually reinforcing factors) usually leads to a decrease in Na and Sr and an increase in Mn and Fe contents of the progressively more altered calcites. These trends are frequently accompanied by depletions in “0 and 13C. SAMPLES AND ANALYTICAL TECHNIQtiES A total of 319 Paleozoic fossilshave been separated manually for this study. AH were scanned optically, analyzed for b”O and 613C.and 276 samules were analvzed also for trace (Sr, Na, Mn,’ Fe, Mg) and- major (Ca, insoluble Residue = I.R.) composition. Sample descriptions, locations and analytical data are summarized in Appendix I. Analytical techniques for chemical data and their pm&ions and accuracies have been described in VEIZERet al. ( 1978)and BRANDand VE~ZER (1980). Isotopic analysis of carbon dioxide, produced by reaction of the sample with 100% H,pO, at 5O”C, were carried out on a semi-automatic MM 903 mass-spectrometer with tripple collector system. Anaijkzal reproducibility for individual runs is better than +0.05%x1for both 6”O and 613C.

Geochemistry of brachiopods The overall reproducibility with respect to the laboratory working standard, expressed in PDB, is -0.15%. The samples studied range in age from Ordovician to Permian, but the bulk of fossils are of Devonian age, because this time interval was considered critical for the aims of the present study. The present sample population encompasses the families Athyrididae, Atrypidae, Brachythyrididae, Buxtoniidae, Chonetidae, Cyrtospiriferidae, Deltyrididae, Dictyoclostidae, Echinoconchidae, Enteletidae, Fardeniidae, Gigantoproductidae, Linoproductidae, Lissatrypidae, Meekellidae, Meristellidae, Nucleospiridae, Pentameridae, Pholidostrophiidae, Plaesiomyidae, Productellidae, Reticulariidae, Spiriferidae, and Trigonirhynchiidae and contains 38 genera. With the exception of 5 samples from Utah, Alaska, Oklahoma and Kansas, all studied specimens originated from various localities in Canada (Appendix 1). The general geology of Canada is summarized in DOUGLAS (1970). OXYGEN AND CARBON ISOTOPIC

RECORD

The histograms of 6’*0 and 613Cstratigraphic variations in Paleozoic brachiopods are given in Fig. 1. For oxygen, the data show a considerable spread of values, but the most ‘80-rich Ordovician, Silurian and Devonian samples are - -4k PDB. The 6’80 for younger samples increases to near-modern values of

Carboniferous kl=37)

I

/-LJ_

J-0

Permian h=29) xLL--l-n

-

Corbontfarous (n=37)

Devonion In:1781

Silurian (n=541

Ordovicim In=20

I

l-l

FIG. 1. Variations with age of 6°C and 6’*0 in Paleozoic brachiopods.

1681

about - 1%” during the Permian, in accord with the previously indicated Late Paleozoic oxygen isotopic shift (cf: VEIZER and HOEFS, 1976). Carbon isotopes show a similar trend towards heavier values, particularly during the Carboniferous and Permian. This is in agreement with the general features of the previously published 6’3C-,t, secular trend (VEIZERet al., 1980; LINDH, 1983; HOLSER,1984). Nevertheless, the spreads of aI80 and 613Cwithin populations (age groups) are similar to, or in excess of, variations between populations. It is therefore essential to evaluate the role of primary and secondary factors on within population variations of ‘8O/‘6O and ‘3C/‘zC. Post-depositional alteration of brachiopod shells

Consideration of the chemical composition of Paleozoic brachiopods shows that their modal values are shifted in the direction expected from post-depositional alteration, although a substantial number of shells still retain primary chemistry (Fig. 2). Factor analysis of the total population (Table 1) identifies four factors controlling the chemical and isotopic composition of brachiopod shells. Factor 1 represents simultaneous loss of Na and Sr from the shells as a result of postdepositional recrystallization. This is principally a consequence of Na and Sr partition coefficients between calcite and ambient water being less than 1 (cf: VEIZER, 1983a). Factor 2 represents mostly precipitation of a later diagenetic Fe, Mn-enriched calcite cement into primary and secondary void spaces. This may also slightly increase the Mg concentrations. These two factors are not a unique feature of the present sample set, but have been ubiquitously recovered and documented in many previous carbonate studies (e.g. BRAND and VEIZER, 1980, 198 1; AL-AASM and VEIZER, 1986). Factor 3 is a sole control of carbon isotopic composition. The overall lack of correlation between trace elements and 613C is probably due to the fact that post-depositional carbon isotopic variations are of lesser magnitude than the primary scatter caused by the previously discussed biogenic fractionations (vital effect) and by secular variations in carbon isotopic composition of marine bicarbonate. The fourth factor, reflecting the complementary antithetic relationship between carbonate and noncarbonate contents of the shells, is of little interest for the present discussion. Oxygen isotopic composition shows only a very weak affinity to factors 1 and 3, indicating that secular variations in isotopic composition of seawater (or in temperature) and post-depositional recrystallization may all be involved, but the relationships are complex. If the isotopic composition (or the average temperature of seawater) were changing systematically during the Paleozoic, the constituent brachiopod age groups would have been characterized by different starting isotopic compositions. Each age group should be therefore evaluated independently. In particular, the Permian and Devonian populations are of critical importance, because the former may represent the begin-

1682

J. Veizer, P. Fritz and B. Jones

.: ..

40

30 % 20

IO :

mmNQ

(102

Ib3 wm

mmMn

1104

1’01

pm-

Sr

IO 3

110’

b

Fe

N= 276

pm Mg Pm. 2. Histograms of chemical camposition of Paleozoic brachiopods. Heavily stippled area represents typical concentrations observed in Recent brachiopods (LOWENSTAM,196 I ; POPPet al., 1986b; LEPZELTER et al, 1983; MORRWN and BRAND, 1984). Theoretically expected concentrations far inorganic calcite in equilibrium with present day sea water (cf the summary in VEIZER,(1983a), as well as ISHIKAWAand ICHIKUNI (1984) and FKJSENBERG and PLUMMER(1985) for sodium) fall within the stippled range.

ning of the near-modem conditions, while the latter would mark the termination of the Early-Middle Paleozoic steady-state (cJ Fig. 1). Factor analyses of Permian brachiopods (Table 2) reveals three dominant factors controlling the chemical and isotopic composition of the shells. Disregarding factor 5 (leaching of Na and Fe from the Insoluble Residue during laboratory digestion of samples), “0 depletion appears to have been controlhxl by post-depositional recrystalhzation of the shell calcite. This also causes the loss of Sr and Na (Factor 1). 13Cdepletion, on the other hand, is related to a subsequent episode of void filling by Mn-rich calcite cement (Factor 2) although the simultaneous decrease in Na and Sr indicates that factors 2 and 1 are not entirely independent. The scattergram of Na (or Sr) versus al80 (Fig. 3) clearly shows the depletion of these variables with increasing degree of diagenetic alteration. At Na values typical for Quaternary marine low-M@,calcites, the S’*O

is as heavy as - 1o/wPDB. Although the primary 6”O scatter, due to temperature variations, could be estimated as about -2.5 ? 1So/w, the bulk of the trend is clearly due to post-depositional alteration, In a similar manner, the scattergram of 613Cversus Na (or Mn, Sr) indicates +4 -t 1.5% PDB as the likely primary values (Fig. 4). In the Devonian population, the trends are less pronounced but nevertheless discernible (Table 3). In this case, in contrast to the Permian, the depletion in “0 is controlled by precipitation of Mn, Fe-rich calcite cement into primary and secondary void spaces (Factor 2). The small gradient in aI80 (or d”C), as opposed to the large Mn spread, is a consequence of chemical attributes of the void filling cement. This cement differed from the low-Mg calcite of the shells by more than two orders of magnitude in Mn content, but only by -35% in 6”O. Consequently, a small to moderate cementation results in a discernible increase in Mn con-

Geochemistry of brachiopods TABLE 1 VARIMAX ROTATED FACTOR ANALYSIS OF ALL PALEOZOIC BRACHIOPOOS FACTOR 2

FACTOR 3

FACTOR 4

Na Mn Fe sr IR

0.017 0.352 0.141 0.026 0.916 -o-s-r5 -0.043 0.648 xiv5

-0.075 -0.140 0.069 0.252 0.006 0.634 _. 0:045

0.725 K33i 0.061 -0.208 -0.006 -0.327 0.130 0.032 0.269

0.089 0.133 -0.581 ra?-5 0.016 -0.095 -0.009 -0.151 0.573

Eigenvalue

1.712

1.291

0.951

0.512

FACTOR 1 6')C 6'80 log ca

109 4

log log log log Log

Percent Of variation explained Interpretation

38.3

28.9

postdepositional recrysta1lization and mineralogical stabilization

Void filling by ferroan calcite cement

21.3 ?

11.5 Total carbonate

tent, but only in marginal alteration of the whole shell 6”O. 13C depletion is a consequence of post-depositional recrystallization (Factor l), the latter also causing the loss of Mg. The scattergram of 6’*0 versus Mn (Fig. 5) shows that the heaviest 6’*0, at Mn values for Quaternary marine low-Mg calcites, is about -4%0 PDB and a conservative estimate for the total primary range is perhaps -5.5 f 1.5%0. This is -3%0 lighter than the estimate for the Permian, regardless whether the total range or the heaviest values are considered. Similarly, the primary range in 613Cvalues appears to have been -0.5 f l.O%a PDB (Fig. 6), that is -4.5%0 lighter than the Permian. Similar evaluations for Silurian and Ordovician populations, although less well constrained due to smaller population sizes, indicate 6”O of - -6.5 f 1.5 and -5.0 f 0.5%0, respectively. For carbon, the likely primary values are 0.0 f 1.5 and + 1.0 +- 0.5%~ Carboniferous samples show a considerable spread and, at this stage, the likely primary values cannot be narrowed to better than - -4.5 t- 1.5%0 for 6180 and perhaps +2.0 + 2.0%0 for 613C.This may be partially a reflection of a considerable secular drift during this time interval (see Fig. 1). SECULAR VARIATIONS

The above evaluation shows, that the primary difference in 6180 between the Permian and younger and the Devonian and older brachiopods is at least 3%0. Complementary results on additional 3700 Paleozoic brachiopods, studied by a comparable trace element and isotope approach (LDWENSTAM, 196 1; BRAND and VEIZER, 1981; BRAND, 1981, 1982, 1983; POPP et al., 1986a,b,c; ADLIS et al., 1985), are in full agreement with the outlined secular trend. This observation is supported also by the fact that several mineral phases, with variable stabilities and thus alteration potentials, show overall aI80 secular trends of comparable magnitudes (cJ: WEBER, 1965a,b; VEIZER and HOEFS, 1976; SHEMESHet al., 1983). Our sampling density does not permit the precise timing of the advocated shift, but it may have been confined to the Devonian-Mississippian

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transition (POPP et al., 1986c). In the subsequent text, we shall explore the consequences of the above isotopic shift, but will not dwell on the finer scale age structure for 6180. Secular variations in 613Cappear to indicate a continuous enrichment in 13Cfrom Silurian to Permian, in accord with the previously published Phanerozoic curves (VEIZER et al., 1980; LINDH, 1983; HOLSER, 1984). This overall trend is also visible in the Devonian period (Fig. 7), with the selected age groups each representing -7 million years duration. However, the scatter of data is considerable and at least 1%0of the total variation can be attributed to differences in biogenie fractionation at the generic level and another l%o represents within genera scatter (Fig. 8). Consequently, brachiopods can be utilized for studies of 6t3C secular variations only at a level of a single genus or even species. Sampling at this level over an extended time interval is, however, difficult. Nevertheless, the reality of the indicated unidirectional Devonian age trend (Fig. 7) is confirmed by the corresponding Eifelian to Frasnian shift in the average 6°C for the genus Atrypa. No overlapping genera were sampled over the Emsian-Eifelian time span. These difficulties discourage studies of 613C secular variations on
TABLE 2 VARIMAX ROTATED FACTOR ANALYSIS OF PERMIAN BRACHIOPOOS

61% 6180 log ca log Hg log Na log Mn log Fe log sr Log IR Eigenvrlue

FACTOR 1

FACTOR 2

-0.062 0.826 m 0.028 0.532 -0.148 0.020 0.520 m

0.828 -7ix5 0.058 -0.552 0.347 -0.644 xr5a7 0.601 m

FACTOR 5

0.960

Percent Of variation explained

16.7

Interpretation

postdepositional recrystallization and mineralogical stabilization

Void filling by sparry calcite cement

Leaching from Insoluble Residue during laboratory digestion of samples

Note that, for convenience, factors are numbered as in Table 1. They ape not listed in the order of their relative importance.

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J. Veizer, P. FritzandIS.Jones

FIG. 3. !&~ttergramof S”O and Na concentrations in Permian brachiopods. See Fig. 2 for explanations.

Salinity

CAUSES OF O’8/O16 SECULAR VARIATIONS Accepting the primary nature of the observed 6’*0 signal, it is essential to search for causal relationships.

It may be suggested that the changing Late Paleozoic S”O signal reflects an evdution in biogenic fmctionation of oxygen isotopes by brachiopods. This, however, is unlikely because the trend is also observed in the weIl preserved skeIetal mater&I of other phyIa, such as rugose corals, crinoids and some moIlucs (BRWD and VEIZER, 1981; BRAND, 1981, 1982, 1983; LEUTLOFF and MEYERS, 1984), in inorganic marine cements (LEHMAN, 1983; JAMEX and CHOQUETTE, 1983), and in bulk rocks (VEIZER and HoEES, 1976). This argues for a cause of more fundamental significance, such as changing salinity, temperature, or oxygen isotopic composition of sea water.

Living and fossil brachiopods are exclusively marine organisms living in normal salinity waters (RUDWICK, 1970, TASCH, 1980).This and their usual association with tabulate corals and crinoids as a Paleozoic offshore community (SEPKOSKIand MILLER, 1985) preclude growth in water of other than normal safnity. Temperature Living brachiop&s are known from all climatic zones, with inarticulates Iiving predominantly in subtropical-tropical and articulates in temperate regions (RUDWICK, 1970,p. 159). Their present-day habitats therefore span the whole G3O”C temperature range of oceanic water masses. The optimal habitats for Paleozoic bmchiopods may have been comparable, but all presently studied specimens grew at low-solitudes, as indicated by paleomagnetic (HABICHT, 1979) and

103

IO’

wm Na RG. 4. Scattergram of S’% and Na coa~nt~tio~s in Permian brachiopocls.See Fig. 2 for explanations.

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Geochemistry of brachiopocls TABLE 3 VARIMAX ROTATED

61% 6180

iogca log W log log log log Log

Na Mn Fe sr IR

FACTOR ANALYSIS

FACTOR 1

FACTOR 2

-0.473 -nzK -0.058 0.553

-0.031 -0.445 -K-r68 0.454 -0.336 0.946 Kin-K-i%-0.020

_E -01082 0.745 o.001

Eigenvalue Percent Of

OF DEVONIAN

1.519

BRACHIOPODS

FACTOR 4 0.014 0.067 -0.436 xiv8 0.071 0.014 -0.208 0.141 0.859

2.287

32.3

48.6

postdeoositional r&rystallization and mineralogical stabilization

Void filling by ferrOan sbarry calcite cement

0.790 19.2

variation explained Interpretation

Note:

Total carbonate

see Table 2.

sedimentological (FRAKES, 1979; ZIEGLER et al., 1977) criteria. The present day tropical-temperate temperature barrier is at winter temperatures of - 15-18’C (HALL, 1964; DODD and STANTON, 198 1, p. 448) and this may or may not have been also the low-temperature limit for the studied Paleozoic communities. In any case, it is not the low, but the high-temperature limit, which is posing problems in subsequent interpretations. Assuming 0% SMOW as the oxygen isotopic composition of the surhcial low-latitude Paleozoic sea water, the calculated temperature of the ambient Permian ocean water would be -26 + 5°C. In contrast, the ambient ocean temperature for Devonian and older seas would have to be at least -35 + 4’C, and possibly more for some marginal Devonian (Fig. 5) and Silurian samples. If “non-glacial” 6’80,.t, of - 1‘%JSMOW (e.g. BERGER, 1979)were accepted, the above temperatures would decrease by -4.5’C. For the present discussion, the glacial effect is considered of subordinate impor-

IO’

tance because both, the “cold” post- and the “warm” preCarboniferous oceans coexisted, at one time or another, with glacial as well as non-glacial climates. Yet, the ~3% 6’*0 difference between Early-Middle and Late Paleozoic brachiopods is a persistent feature. The validity of the above statement is, however, conditional on the assumption that Phanerozoic glaciations were of comparable magnitude to the Cenozoic one. The presently known large scale glaciations have been described from the Early Proterozoic, Late Proterozoic, Late Ordovician-Early Silurian, Late Devonian (Famennian)-Late Permian, and Cenozoic (SCHWARZBACH,1963; FRAKES, 1979; BUDYKO,1977; CROWELL, 1982; CAPUTO and CROWELL, 1985; HAMBREYand HARLAND, 1985), with the terminal Proterozoic one possibly reaching even low latitudes (MCWILLIAMSand MCELHINNY, 1980). Each of these glacial episodes lasted 2 lo-20 million years (FAIRBRIDGE, 1982) and during the Paleozoic Era glaciations appear to have been the norm and not the exceptions (CAPUTO and CROWELL,1985). The usual defense against coexistence of the supposedly “hot” (360°C) Proterozoic oceans with glacial climates is an invoked large diachronism of glacial and non-glacial sediments (e.g. KNAUTH and LOWE, 1978). A similar resort to diachronism is not a viable proposition for the Phanerozoic, where the synchroneity of glacial climates with the “warm” Paleozoic oceans is undeniable. In our view, 35 + 4°C (and more for the Proterozoic) warm low-latitude oceans and simultaneous, possibly even low-latitude (e.g. the terminal Proterozoic), widespread glaciations are difficult to reconcile. Note that even during the ice free times of the Cretaceous, the equatorial surface ocean temperatures did not exceed, and likely were less than, 32°C (FRAKES,1979, chapter 6; SAVIN and YEH, 198 l), although the temperate climates stretched into high latitudes. Furthermore, the brachiopod community with coexisting rugose corals, crinoids, and other phyla was relatively stable during the Phanerozoic (RUD-

2

103

ppm ‘Yvln FIG. 5. Scattergram of 6”O and Mn concentrations in Devonian brachiopods. See Fig. 2 for explanations.

J. Veizer, P. Fritz and B. Jones

1686

IO2

IO3

104

wm Na PIG. 6.8cattergram of Si3Cand Na concentrations in Devonian brachiopods. See Fig. 2 for explanations.

1970; SEPKOW and MILLER, 1985) and the proposition that the same pre-CarboniFerous community should have lived in habitats - 10°C warmer than its later counterpart is paleoecoiogically unpalatable. This is quite apart from the question of whether hving marine communities containing, among others, mollkscs could have thrived, as opposed to bareiy survive, in permanent 35 f 4°C habitats (& READ, 1963). For example, collagen, a body protein present in several metazoan phyla, collapses in thriving reproductive populations at upper limits of these temperatures (RIGBY, 1968; l&By and HAFEY, 1972). Furthermore, BROCK ( !985) concluded that 38’C is the upper limit of tolerance for aquatic vertebrates. As a final point, some meteorologists argue that - 3 1T is an upper WICK,

limit on the temperature of the surface ocean water for atmospheric radiation balances not vastly departing from the present-day conditions (NEWELL and DOPPLICK, 1979). We would iike to emphasize that none of the above arguments is unequivocal (see, for example, VALENTINE, 1985), but their combined weight suggests that the temperature effect alone was not the principal cause of the observed Late Paleozoic S’*O shift of -3%~ Additional support comes from the extention of the isotopic record into the Proterozoic and the Archean. As stated above, the presently known record would demand ~40°C warm oceans in temporal and/or spatial proximity with the well estabhshed Early Proterozoic glaciations (e.g. FRAKES, 1979, p. 38). Further-

FIG. 7. Histogrrugof 6’v vkatiom jn Bmsian (Early Devonian), E&k&n(Middle Devonian) and Frasnian (Late Devoniam Braehiopod!i.Bmpty cimles-genus Atrypu, N1 do&-other genera.

Geochemistry of brachiopods I

I .

I

.

I

.

.

.

.

::.

1687 I

.

.

.

1 Cupolarostrum

:.

.

Anatrypo

:

.

:i”i:~:i””

..:..

:

.

.

.

A trypo

. .

.

Perryspiri fer

FIG.8. 613Cdistribution in various Eifelian (Middle Devonian) brachiopod genera.

more, HOLLAND(1978, p. 205) pointed out that equilibrium temperature for sea water saturated with respect to halite, anhydrite, and gypsum is 18’C, and gypsum precipitates from seawater only at temperatures ~58°C. The gypsum-halite association is therefore indicative of temperatures <25”C (HOLLAND, 1984, p. 427). Evaporites with this mineralogical assemblage, although not necessarily coprecipitated, have been ubiquitous throughout the Phanerozoic (e.g. BRAITCH, 1962) and they have been documented as far back as 1.6 Ga ago (MUIR, 1979). Furthermore, the discovery of gypsum casts in sequences as old as 3.5 Ga precludes at least the existence of the postulated 60-70°C Archean oceans (WALKER et al., 1983). Isotopic composition of sea water The above elimination process leaves the evolution in oxygen isotopic composition of sea water as the only presently known alternative. The buffering of ‘8O/‘6O in sea water by interactions with silicates at high- and low-temperatures (crossover -300 + SO’C) has been discussed previously by CHASE and PERRY (1972), MUEHLENBACHSand CLAYTON (1976), PERRY et al. (1978), LAWRENCEand GIESKES(1981) and GREGORY and TAYLOR (198 1). Overall, a net enhancement of high-temperature interactions will tend to enrich sea water in “0 and vice versa. The observed Late Paleozoic 6”O increase of -3L would therefore demand a substantial enhancement in the rate of high temperature oxygen isotopic exchange (plutonism, deep ocean ridge circulation, etc.) relative to low-temperature processes (e.g. shallow ridge circulation, submarine weathering). This is an attractive proposition, because the Late Paleozoic interval is also a time of very pronounced variations (Fig. 9) in marine 8’Srf86Sr (PETERMANet al., 1970; VEIZER and COMPSTON, 1974; BURKE et al., 1982; POPP et al., 1986b), 613C(VEIZER et al., 1980; LINDH, 1983; HOLSER, 1984), 634S(CLAY-

POOL et al., 1980; HOLSER, 1984), and in “sea level” stands (VAIL and MITCHUM, 1979). These oscillations, at 10’ years frequency, have likely been controlled by global tectonism (VEIZER, 1985) and global tectonism may have modulated also the rate of high/low-temperature exchange of oxygen between silicate crust and sea water. At this stage, we do not have the necessary stratigraphic sampling resolution to decide whether the inferred 6”O shift coincides with a specific Late Paleozoic inflection in the complementary isotopic curves. Similarly, we are not in a position to establish or disclaim the existence of possible second order, less than + 1%o,variations within the overall Paleozoic 6”O secular trend. Despite our preference for this solution, three sets of arguments mitigate against its unquestionable acceptance. Firstly, KNAUTH et al. (1985) and KNAUTH and BEEUNAS(1986) analyzed fluid inclusions from Phanerozoic salts, believed to represent unmodified trapped sea water, and reported near-present day 6”O (but not aD) values as far back as the Silurian. Secondly, the patterns of 6”O distributions in secondary minerals of greenschist facies metabasalts on the modem ocean floor (MUEHLENBACHSand CLAYTON, 1972; STAKES and O’NEIL, 1982), in Cretaceous ophiolitic sections (GREGORY and TAYLOR, 1981), and in the Archean greenstones (BEATYand TAYLOR, 1982; SMITH et al., 1984) are similar. This is consistent with their interpretation as products of alteration by sea water with 6”O of about 00/w(SMOW). Thirdly, MUEHLENBACHS and CLAYTON (1976) and GREGORY and TAYLOR (198 1) proposed that sea water 6”O is buffered at OL (SMOW) by the competing and complementary interactions with silicates of oceanic crust at low and high-temperatures. Accepting these balances, it is impossible to generate the inferred 3%0 shift in 6”O of the Paleozoic sea water. The difficulty is compounded by preliminary indications (POPP et al., 1986~) that the oxygen isotopic shift may have been restricted to the

1688

J. Veizer, P. Fritz and B. Jones

_------------

LEVEL

+3 it-i3

+1 -1 1;

%o

CDM

20

30

TODAY

our present knowledge of the silicate/water exchange reservoir sizes and fluxes. These are open to discussion (PERRY et al., 1978), particularly around subduction zones and in erogenic belts. At this stage, we do not discount therefore the possibility that a change in the oxygen isotopic composition of sea water could have been induced by tectonic phenomena. A resolution or quantification of this question will not only improve our understanding of the hydrologic cycle, but may also provide a tool for understanding the tectonic evolution of the Earth and, alter appropriate corrections, perhaps also the temperature history of its surface.

-

1 -? :

CONCLUSIONS

Combined trace element and isotopic study of Paleozoic brachiopod shells leads to the following conclusions:

1

-c J0

-2 &

-4

-6 / 6

4

x10*

0

2

years

BP

FIG. 9. Secular isotopic age curves for sea water as &rived f+om(a)sulfbr&otopaainevaporites&f&,(b)carbonisotapes in catbonates, (c) strontium isotopes in carbonates and fodsil apatites, and (d) the tentative oxygen isotopes in carbonates, principally brachiopods. For sources of these curves see the text and NOL%R (1984).

Devonian-~ transition and thus could have been rapid. As was the case with temperatures, none of the above arguments is unequivocal. The fluid inclusion studies need conknation andexplanation ofthe disparity between oxygen and hydrogen isotopic compositions. The plausible alteration argument entails a plethora of assumptions, such as the degree of sea water partioipation in alteration processes, the degree of silicate/water equilibration or water/rock ratio, the starting isotopic composition of parental material, metamorphic and crys&zation temperatures, etc. This is particularly the case for the pre-Cretaceous studies. Finally, the tightness of the balance argument is conditional on

1) Paleozoic brachiopods fractionated carbon isotopes at the generic level; 2) In accord with previous studies, Paleozoic brachiopods reflect a general enrichment in 13C in the course of the Paleozoic, with major enrichment in the Permian; 3) The previously described negative correlation of marine 634S,tir.tcand 6’3Ccarbonete (VEIZER et al., 1980) is valid only at time resolutions of 10’ years; 4) The 6”O record of brachiopod shells shows _ 3k depletion in ‘*O during the Late Paleozoic, with Devonian and older samples having S’s0 of G -4% PDB; 5) The proposed oxygen isotopic shift may be a consequence of changing ‘aO/‘6O of the coeval sea water, of warmer (35 f 4°C) Early and Wd-Paleozoic oceans, or both. Our preference is for the first alternative, but both solutions pose problems for complementary geological observations. 6) The acceptance of the Paleozoic brachiopod data as representative of secular variations for 6”O leads to consequences of first order significance, regardless of the actual cause(s) of the oxygen isotopic shift. If the data indeed reflect a changing ‘sO/‘6O composition of sea water, they may indicate the existence of a global tectonic event(s) which caused a net increase of hightemperature oxygen exchange between sea water and silicate crust. If they reflect paleotemperatures, the coexistence of persistent Paleozoic glaciations with 35 -t 4°C low-latitude oceans requires reconsideration of the existing paleoclimatic models. If neither proposition proves satisfactory, alternative solutions must be explored, such as, for example, the possible consequences of the expansion of vascular plants and animals on land in the Devonian. Aduwwle&emnts-This work hasbeen supported financially by the Natural Sciences and Enginqiag Re-search Council of Canada. We acknowledge tech&al support by Robert Drimmie and John Loop; d&&g by Edward I-learn and A. Steele; drafting and preparation of hi Appendix by W. F. Schmiedek typiw by Julie Hayes; &nation of Ordovician samples and oftraoe element data for these samples by I. S.

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VEIZERJ., LEMIEUXJ., JONESB., GIBLINGR. M. and SAVELLE J. (1978) Paleosalinity and dolomitization of a Lower Paleozoic carbonate sequence, Somerset and Prince of Wales Islands, Arctic Canada. Can. .I. Earth Sci. 15, 1448- 146 1. VEIZERJ., HOLSERW. T. and W~rx;us C. K. (1980) Correlation of “C/“C and %/“S secular variations. Geochim. Cosmochim. Acta 44,579-587.

VEIZERJ., COMP~TONW., HOEF~J. and NIELSENH. (1982) Mantle buffering of the early oceans. Naturwissenschafien 69, 173-180.

WALKER J. C. G., KLEIN C., SCHIDLOWSKIM., SCHOPF J. W., STEVENSOND. J. and WALTERM. R. (1983) Environmental evolution of the Archean-Early Proterozoic Earth. In Earth’s Earliest Biosohere: Its Origin and Evolution (ed. J. W. SCHOPF),pp. i60-290.P&&ton Press. WEBERJ. N. (1965a) The 0’8/0’6 ratio in ancient oceans. Geokhimiju 6,674-680 (in Russian). WEBERJ. N. (1965b) Evolution of the ocean and the origin of fine-grained dolomites. Nature 207, 930-933. WEBER J. N. (1967) Possible changes in the isotopic composition of the oceanic and atmospheric carbon reservoir over geologic time. Geochim. Cosmochim. Acta 31,23432351.

WEFER G. (1983) Die Verteilung stabiler Sauerstoff- und Kohlenstoff-Isotope in Kalkschallen mariner GrganismenGrundlage einer isotopischen Palokologie. Habilitationsschrift, Universitat Kiel, p. 15 I. WILLIAMSA. (1968) A history of skeletal secretion among articulate brachiopods. Lethaia 1,268-287. WILLIAMSA. ( 197 1) Scanning electron microscopy of the calcareous skeleton of fossil and living brachiopods. In Scanning Electron Microscopy: Systematic and Evolutionary Applications, (ed. V. H. HEYWOOD),pp. 37-66. Academic

Press. ZIEGLERA. M., SCOTESEC. R., MCKERROWW. S., JOHNSON M. E. and BAMBACHR. K. (1979) Paleozoic paleogeography. Ann. Rev. Earth Planet. Sci. 7, 473-502.

1692

J. Veizer, P. Fritz and B. Jones

APPEND1

X

1:

SAMPLE NO.

Sample

descriptions

by geologic

SPECIES

age and locality.

c

PDB 13

PERMIAN Assistance 410 411 Assistance 421 Takhandit

Formation, Marie Bay, Melville Island, Dictyoclostus sp. 1.9 Dictyoclostus sp. 2.0 Formation, Sabine Peninsula, Melville Liostella sp. 5.0 south side of Yukon River, Formation, :,"f Linopthoductus, sE;T 3.8 Acheson waters ed west Spirifer sp. 4.1 Dry Canyon, Wasatch Mtns., Utah 422 Pustula subhorrida 1.5 Kandik, Alaskd 414 Echinoconchus sp. 3.3 Linoproductus sp. 1.4 416

Mg to Sr on I.R. free basis. PPM NA

PPM MN

PPM FE

PPM SR

I.R.

1709 1705 NUT

1368 1180

480 440

1292 1406

929 836

3.80 2.96

2089

1804

54

1150

91 0

15.21

1927

1403

15

4903

959

83.57

-4.6

3709

863

5

2313

848

19.50

-7.9

4936

342

69

1020

338

48.81

0

PDB

PPM

18

MG

NUT -4.2 -5.3 Island, -3.8 Yukon

-5.5

B

-4.1 -4.7

5989

1637

129

6957

984

27.9i

2.4 -2.5 0.9

-1.0 -3.2 -2.6

3657 3537 2222

1196 880 1186

5 55 12

427 2182 125

1008 759 834

7.08 8.52 13.66

3.3

-7.6

-0.3 -0.9 -3.6

-7.6 -2.9 -5.1

1264 2208 2168

344 734 303

115 155 192

464 271 240

401 532 375

10.99 1.07 0.54

0.4

-7.8

‘326

163

47

326

6 5 t,

2.2 1.3 3.0 1.8 2.2 3.3 0.7 2.1

-2.7 -1.4 -1.2 -2.6 -1.2 -1.6 -10.6 -2.9

2286 2893 1638 2365 2842 708 1308 2566

1675 1280 2012 1091 1205 683 118 1162

41 17 5 42 16 5 26 27

820 1394 928 433 1022 075 69 2569

945 899 1290 952 2357 677 352 121s

50.31 19.60 12.01 24.90 32.68 4.96 0.29 25.23

3.5

-10.1

1579

558

10

468

9/T

57.94

4.5 5.2

-6.8 -5.3

491 736

1234 1127

5 5

129 1247

1405 1305

1.33 10.40

lineatus neoinflatus

0.4 4.2

-6.1 -4.4

9148 2245

276 1231

182 53

641 1441

4 37 837

0.52 54.11

%%%%tion, Beaver River, NUT 310 Choristites esplanadensis Banff Poriii iation, Wapiti Lake, BC

1.5

-4.9

2507

690

157

35

69x

30.74

-1.2

-7.1

1.3

-3.8

3502

1620

564

2067

1116

4.90

-1.5

-6.7

6168

462

5

32

1154

11.0:

-0.2

-5.6

1902

788

27

217

944

38.39

-4.2

3493

1066

16

337

945

6.56

Ogilvie 423 Peace

Dictyoclostus Mtns., Jungle Dictyoclostus River. La Biche

uralicus Creek, NUT sp. River. Alberta

Spirifcr sp. Peel Plateau. Mt. Effie. NUT 427 In&t. terebratulid Peel Plateau, Peel River, NWT 408 Dictyoclostus sp. lclostus sp. er sp. zastus sp. 5tus sp. zetulid rhynconellid 410 Indet. 426 Dictyoclostus sp. Sheep Creek! NWT 404 Spirjfer sp. West of 406 407 localitv

-------. 400

401

Bear Creek, Smirifer s urrifer unknown

NWT so. sp.

Linooroductus bictl foclostus

cf. cf.

Mt.

Head Format 314 Gioantoeroductus brezeriarus Pekisko Formation, Jo&fish, Yukon 329 Birifer sp. Pekieko Fotmatlon. Liard River. Nahanni 309 Spirifer’centronitus Pekisko Formation, Monoghan Creek 306 sirifer centronatus Pekisko Formation, Moose Dome area, BC 302 Spirifer ainnewankensis Pekisko Formation, Mt. Rundle, Alberta 320 Erachythyris sp. Pekisko Formatioir, Nahanni Range, NUT 311 Spirifer cf. striatus Pekisko Formation. Waoiti Lake. EC

Rundle 325

Shunda

330

0.10

Syrinsothyris sp. Formation, Slate Creek, Rhipidomella sp.

2.7 Butte.

NUT 216

-4.1

1535

271

33

643

1198

4.20

3.0

-4.6

4121

1077

10

626

803

0.32

1.7

-4.2

4711

1099

11

130

1082

0.82

3.2

-3.6

7230

316

14

164

921

0.58

-5.3

1701

1041

16

290

1124

67.94

-2.6 2.9

-5.8

3327

603

23

108

1237

0.25

2.1

-5.6

1517

721

10

117

1269

12.28

BC 0.6

-6.2

2.3 2.5

-5.0 -4.1

1577 971

1307 966

26 5

602 166

951015

23.40 18.31

1.5 1.7 0.6

-5.0 -4.3 -6.5

1243 828 3144

519 527 462

13 16 54

139 15 190

1219 750 485

25.66 31.86 38.64

1.6

-5.6

1464

207

26

187

867

18.44

Formation

Spirifer

rowleyi

Geochemistry ofbrachiopods

1693

SAHPLB PDB PDB PPM PPM PPM PPM PPM % NO. SPECIES c 13 NA MN FE SR I.R. 0 18 MG -------------------------------------------------------------_--------------_ .-..-----_---_-_-----_ Wetumka Formation, near Ada, Oklahoma 431 Mesolobus mesolobus 0.7 -4.8 2239 180 940 949 10.95 5067 Beaver Dam, BC 327 Martinia rostrata -1.4 -8.6 4050 813 85 399 947 2.86 Canyon Creeks-BBC 307 Syrrnqothyris cf. hannibalensis 5.7 44 -3.1 1707 134 684 1505 4.87 Captain Creek Island, Ottawa, Kansas 433 Enteletes hemiplicata 2.0 -7.4 2153 193 778 2590 1338 0.33 Moose Lake, BC 321 Spiriter esplanadensis -0.6 -6.7 474 1072 1469 3900 152 1.15 326 Spirifer sp. 1.5 1126 -5.9 2990 584 89 1276 1.16 Mt. Greenock, Alberta 301 Cleiothyridina lata 1.1 7417 1187 35 714 1430 -5.3 3.98 Mt. Hannington, BC 322 Spirifer sp. 2.1 -3.4 1518 203 29 701 151 10.10 -5.7 1.8 2570 295 57 835 667 54.17 Nai!ini R~afs&%"adensiS Spirifer aff. rowleyi 1.5 -4.6 865 420 12 907 69 16.78 Ogilvie Mtns., Jungle Creek, NWT 430 Indet. brachiopod 4.3 -7.5 3189 1369 388 5394 1244 29.77 432 Indet. brachiopod 4.5 -5.6 1616 769 10 713 545 31.34 Pekisko Lake Nahanni Ridge, NWT 316 Clerothyridina sp. 2.1 -5.8 5599 359 38 490 589 33.56 Slate Creek BC 303 Spirifer minnewankensis 1.8 -6.4 5320 260 16 10 552 0.15 Wapiti Lake, BC 319 Punctospirifer sp. 2.0 -6.8 2408 661 96 352 1236 0.20 Wapiti Pass, EC 304 Setiqerites setigerus 1.0 -5.9 5658 349 60 1294 473 43.74 308 Maqinata sp. 1.2 -5.9 2216 632 54 1134 562 58.02 y.DEVONIAN Hay Hay River, NWT ..-River _. AtrYpa _.Formation, “N-IA cosmeta 1.4 -5.9 1010 272 86 429 425 19.61 HR-1B Atrypa cosmeta 0.8 -6.2 1421 667 151 914 700 5.36 HR-1C Atrypa cosmeta 0.9 -5.6 1512 538 185 923 659 5.37 HR-2A Atrypa cosmeta 1.0 -6.3 1955 430 215 1434 618 4.09 HR-28 Atrypa cosmeta 1.0 -6.4 1571 441 172 865 628 5.94 HR-3 Atrypa cosmeta 0.4 742 -6.3 2662 291 2238 738 6.48 M.DEVONIAN Bird Fiord Formation, Bathurst Island, NWT I 2022 358 194 1577 0.7 -5.8 733 1.63 8 0.7 2258 323 -6.1 184 599 550 1.65 9 1330 273 359 2604 0.5 -6.5 455 0.88 10 2693 -6.3 192 208 396 984 0.8 1.27 11 -1.6 1443 297 161 2603 433 3.99 12 ;*: -6.4 4499 1436 325 1077 1002 2.29 13 0:1 -6.2 3053 562 577 286 1093 5.54 14 Atrypa sp. 1.0 -6.3 1391 317 186 1222 513 1.89 Cupularostrum sp. 15 -1.9 -8.9 2698 138 1190 5811 722 0.17 16 AnatrYoa (Variatrypa) arctica -5.5 0.4 1692 390 155 1246 611 0.30 17 -6.0 1527 323 0.7 173 1597 437 1.61 18 2723 0.9 -6.7 622 200 790 937 3.87 19 1864 279 0.7 -6.1 248 2703 448 2.26 20 Cupularostrum repetitor 1.6 -6.5 21 2770 1.80 -5.9 5260 1377 328 929 0.3 7764 -1.5 -7.4 1703 76 1250 487 0.36 :: 1078 138 2346 2.08 -5.0 363 509 0.7 -6.9 1.3 :"5 1161 1.23 4514 2344 106 872 0.1 -5.7 26 353 1263 1437 617 76 12.67 0.6 -5.6 27 2021 767 1.73 2348 288 288 -6.1 0.5 28 -0.2 -6.6 29 2070 1069 1807 253 258 9.35 -6.6 0.8 2476 30 -7.4 3204 528 181 866 6.34 1.7 31 720 2143 298 106 656 1.73 -6.4 -;*: 32 -6.7 33 -6.4 l:o 34 913 1147 -5.7 3049 814 347 8.52 0.2 2132 35 2866 576 274 4.59 -6.8 798 427 3419 36 ::: 1420 363 603 1.49 -6.4 37 -8.6 2717 38 1408 253 997 3.37 -6.6 Z 5138 1581 39 2174 282 249 1006 1.42 -6.3 -;*; 40 204 2169 -7.4 2663 949 670 5.64 41 958 5.19 0:1 -7.2 5113 1464 402 3550 42 -6.2 -0.3 43 -0.4 1507 278 1151 1087 -6.0 3927 0.44 44 Atrypa arctica -7.8 Schiso horia sulcata 2114 833 1296 2.20 4460 137 -6.9 :z &Spinatrypa) borealis -7.1 2457 498 241 371 899 21.27 47 Cupularostrum repetitor 127 545 4351 503 4.58 -7.7 2681

J. Veizer, P. Fritz and B. Jones

1694

SAMPLE NO.

PPM SR

630 162

5702 1269

816 1371

3.54 1.97

575 1108 319 579

344 529 511 418

702 1353 5313 4047

881 1153 917 1841

6.26 1.25 2.08 2.06

2396 1082 2193

348 2140 347

204 26 169

1489 94 701

758 1684 2944

0.25 20.48 1.14

4016

419

187

1934

1246

0.65

-5.4 -6.0 -5.5 -5.0 -5.0 -5.1 -6.6

1884 6390 2028 3942 1469

283 2344 277 405 396

192 200 156

1197 1984 2202 141 989

5bO 957 585 1067 651

I .2) I, 0.32 0.97 2.94 2.95

2464 1300 1452 1193 1129

348 349 428 430 373

222 221 200 110 123

2613 3786 2708 2301 2584

586

2.40 1.23 0.33 i.?4 l.'?i

1998

304

145

1943

2609 1583 2470 1866 3599 2437 1605 3957 1900

316 395 494 465 916 590 1544 422 348

181 85 119 236 10 143 8": 180

2537 506 1328 986 1240 1770 416 538 52

212; 804 8 3? T?!: 681 763 1054 4 5'7 90:'

2507 4416

235 1077

252 453

397 2651

5R0 62R

3768 1478

397 479

101 79

34 36

1645 3441 2267

328 1358 552

18 33 105

170 297 1105

678 9 67

744

0.39 0.73 1.25

0.8 0.6 0.0 -0.5

-5.9 -6.1 -6.3 -5.4 -6.4 -6.4 -4.9 -6.4 -7.2 -6.1 -5.8 -6.5 -6.6 -5.9 -6.4 -5.1 -7.3 -5.4 -4.6 -6.0 -6.9 -5.7 -6.3 -5.5 -6.0 -6.0 -6.6 -6.3 -6.6 -5.1 -5.5 -6.9 -6.7

3573 993 1996 5237

603 396 2139 1758

220 36 80 72

2004 179 628 821

/17 697 121? 1060

4.34 0.79 0.52 0.2;

-0.2 0.6 0.6 0.2 --0.2 -0.2 -0.3 -0.8 -0.5 -0.6 -0.2 -0.1

-4.0 -5.6 -6.2 -6.4 -6.2 -5.8 -5.5 -5.9 -6.4 -6.6 -5.8 -4.7

1822

324

35

937

535

0 , 65

1470 1757 1269 1703

485 521 1824 2024

87 165 23 27

1303 415 102 43

1330 1378

0.14 0.01 8.70

2007

2588

38

238

1414

0.5R

945 1892

1823 1476

13 62

42 632

1215 1134

3.76 0.00

0.3 0.8 0.2 0.6 0.2

-6.1 -4.9 -4.4 -5.5 -6.0 -6.2 -6.0 -5.5 -5.9

2057 1118 1456 1647

368 183 2128 425

37 10 14 109

383 20 34 322

?3 1 577 1206 65h

2.25 lR.07 25.66 12.13

1717

757

269

2246

831

0.60

3701

4486

70

1200

10 I ’

-0.7

Bird Fiord Formation. Devon Island, NWT Atrypa sp. 110 111 Schizophoria sulcata 112 Atrypa sp., Schizophoria sulcata 120 127 Atrypa sp. Spinulicosta sp. 128 Cupularostrum sp. A 129 Bird Fi ord Formation. Ellesmere Island, Atrypa sp. 1 2 Atfypa sp. Spinatrypa (Spinatrypa) borealis 3 4 5 Cupularostrum repetitor 6 AtrYDa Sp. 63 Schioophoria sulcata 64 Spinulicosta sp. e64" Cupularostrum sp. AtryDa sp. B :: Atrypa sp. B Atrypa sp. C 104 Atrypa sp. 105 106 Atryua sp. Atrvpa sp. 107 108 Atrypa sp. B 109 mnzPsverdrupi 116 117 Atrypa sp., Schizophoria sulcata 118 Atrypa Sp. B 123 Schizophoria sulcata 125 Atrypa sp. B 126 Cupularostrum sp. 134 Atrypa SP. C 135 Schizo horia sulcata 136 mpinatrypa) borealis 137 142 AtrYDa Sp: B 143 144 Atrypa sp. B Perryspirifer scheii 145 146 Schizoph aria ' sulcata Bird Fiord Formation, Norfolk Island, NWT 119 Elythyna sverdrupi. 121 122 124 130 Atrypa sp. B larrenella SD.

llcata 140 _ Warrenella SD_ -141 Atrylp sp. B tj.&.DEVONIAN Bird Fiord Formation, Devon Island, NWT 66 Atrypa sp. B 67 Ivdclinia qrinnellensis 60 Warrenella sp. Atiy" sp. Spinulicosta sp.

0.7 -0.1 0.6 -0.4 0.5 0.6 1.1

0.3 0.3 0.4 0.2 U.6

0.0

0.1 1.6 0.3 0.1 0.4 0.2 0.2 0.5 2.5 0.9

0.3 0.1 0.3 0.2 0.5 0.4 0.8 -i/i

SiEpPlulcata CuDUlaroStrUm

Sp.

B

0.8

PPM

PPM

NA

MN

:

PPM FE

PDB c 13

SPECIES

PDB 0 18

PPM MG

-6.1 -6.7 -7.6 -6.7 -7.4 -6.4 -7.6 -6.2 -6.9 -5.4 -6.3 -7.9 -6.0 -10.7

4667 5563

845

3506 2194 1612 3152

279

8":

476 405 540 493

1078 67’1

704 123

i

.R.

4.12 2.78 1.72 1 .;: 1.36 6.6.7 0.16 0.51

1i , : b 11.4’;

4.27 1.5il 13.3B

2.51

i.63

Geochemistry of brachiopods

1695

-------------------------------------------------~------------~-_-----------~------------~-~----~ SAMPLE NO.

SPECIES

PDB c 13

PDB 0 18

Bird Fiord Formation, Ellesmere Island, NWT 69 Spinatrypa sp. 0.0 -5.8 70 0.9 -5.3 71 -0.2 -5.7 72 -0.9 -6.5 75 -7.1 1.5 76 Warrenella sp. -5.0 _;*: Nucleospira sp. -6.1 0:s -5.8 1.1 -6.2 0.7 -6.8 89 Perryspirifer scheii -0.7 -5.6 90 Warrenella sp. -0.1 -5.0 Atrypa sp. B 0.2 -4.7 Atrypa sp. -6.6 1.2 -0.4 -8.3 -0.8 -7.5 95 Elythyna sverdrupi 0.5 -5.8 Atrypa sp. B -0.2 -7.3 ;; Atrypa sp. 0.2 -6.7 98 Atrypa sp. -0.7 -7.8 100 Atrypa sp. C -0.6 -6.4 101 Elythyna sverdrupi -0.7 -6.7 102 0.1 -6.1 103 0.2 -6.2 Bird Fiord Formation, Norfolk Island. NWT 73 Spinatrypa sp. -0.2 -5.3 74 Schisophoria sp. -0.5 -5.5 L.DEVONIAN Bird Fiord Formation, Ellesmere Island, NWT 113 Elythyna sverdrupi 0.4 -5.6 114 Elythyna sverdrupi -6.1 115 Elythyna sverdrupi -0":: -5.2 Eids Formation, Ellesmere Island, NWT Eoschuchertella n. sp. EO-1 -0.4 -6.4 EO-2 Eoschuchertella n. sp. -0.3 -4.6 EO-3 Eoschuchertella n. sp. -0.1 -4.5 Eoschuchertella n. sp. -0.1 EO-4 -5.0 EO-5 -0.1 -4.6 EO-6 -0.1 -4.9 EO-7 -0.2 -4.5 EO-8 -0.4 -4.5 0.3 EO-9 -4.1 EO-10 -5.2 EO-11 -0.3 -5.6 EO-12 -0.4 -5.0 PER-l 0.2 -4.8 PER-2 -1.2 -13.4 PER-3 -5.1 0.3 PER-4 0.1 -6.0 PER-5 0.4 -4.5 PER-6 0.2 -4.7 PH-1 -0.6 -10.7 PH-2 -0.6 -6.6 PH-3 -0.4 -6.8 PH-4 -1.3 -9.1 PH-5 -0.1 -5.8 PH-6 -0.5 -5.6 PH-7 -5.8 PH-8 -::: -7.3 PH-9 -0.3 -5.9 SILURIAN Barlow Inlet Formation, Cornwallis Island. NWT 0.9 -6.3 -6.4 0.8 1.2 -6.4 -6.4 1.7 0.4 -6.8 -4.9 ;:: -5.6 -6.0 0.8 -6.0 0.3 -6.4 2.0 1.4 -6.1 -5.7 ::: -5.5 0.9 -6.2 0.i -5.6 3.1 -6.5 Island, NWT A 3.8 -6.7 1.2 -6.7

o;o

PPM MG

PPM NA

PPM MN

PPM FE

PPM SR

% I.R.

1737 1111

812 526

189

122

4048 262

804 724

2.65 28.26

7399 2630 2767

1243 338 1099

456 164 51

8975 1360 672

900 575 918

6.51 3.37 3.34

1766 1346 4344 1809 943

395 418 523 2097 1592

101 88 227 54 11

302 1295 4512 248 177

606 576 707 1309 1144

3.03 0.76 4.11 24.91 16.45

3852 2746 2135 2951 1639 2922 1477

296 1664 1927 507 583 661 468

224 203 118 1:: 165 58

1890 533 568 821 503 2452 517

556 1118 1095 650 730 739 580

7.88 14.32 0.17 0.36 0.41 6.09 0.93

2449 2594 1212

1325 481 520

88 159 84

505 1212 492

1106 638 889

1.96 3.92 0.48

4514

1532

57

1152

1015

0.33

2263 1628 1515

793 471 1291

46 53 50

210 227 351

1119 683 1133

1.06 0.30 5.52

2732 3690 3588 4076 3502 3291 4784 3082 4588 3348 3319 4186 1477 4079 1736 744 1570 1229 3309 3315 4998 3625 4705 5523 4852 2545 4914

716 1478 1239 1089 1752 641 1034 1863 1083 973 1584 1555 1729 134 1954 1121 1548 1254 767 1004 1257 991 1276 1652 1424 869 1372

97 56 128 148 59 110 185 64 152 88 101 85 27 165 35 22 29 28 309 161 292 217 209 195 250 162 275

314 214 454 575 312 219 405 273 624 225 288 449 255 416 154 46 108 176 1982 1391 3051 716 2621 1862 2832 1087 3148

1010 1296 1063 1369 1251 1014 1168 1384 1203 1083 1294 1306 1052 785 1219 1545 1022 773 1126 895 1095 1871 989 1095 1073 789 1105

23.37 7.78 11.22 8.02 2.00 27.74 2.02 1.56 10.14 13.79 2.36 4.06 5.31 7.49 1.28 24.78 6.01 19.72 6.83 17.48 3.69 8.22 6.76 8.22 4.05 25.79 3.78

6269 7257 6475 6627 19626 3468

164 184 223 239 136 267

:; 98

601 417 475 395 1187 2010

542 852 704 667 587 669

3.26 4.22 7.12 4.89 6.26 8.47

6127

145

108

1948

705

3.04

6245 11202 3185 8385 9844

143 159 136 181 607

60 % 75 209

621 643 493 700 2567

965 669 741 897 850

3.76 10.98 2.21 8.66 6.93

2576 6464

350 1300

149 12

521 116

902 1736

2.09 0.01

% 57

1696

J. Veizer, P. Fritz and B. Jones

________________________---__---___.----___---____-------------. SAMPLE NO.

PDB c 13

SPECIES

PDB 0 18

PPM MG

PPM PPM NA MN _____-_---.

PPM PPM FE SR ______-._

8 I.R. ._-.

Douro Formation, Ellesmere Island, NWT

d,

204 206 207 209

Atrypoidea netserki Gypidula sp. Protathyris praecursor Atrypoidea netserki

B 220 Atrypoidea phoca 224 Protathyris praecursor 225 Atrypoidea foxi f. A Douro Formation, Somerset Island, N'WT

0.1

-6.9

1030

154

28

-0.2

-7.3

2995 11229 3649 3996 4119 3753 7623 4121 30191 6279 3773 3484 3232

221 310 299 349 955 849 278 477 140 129 220 1425 335

60 46 5 49 39 38 28 27 52 133 33 20 22

2334 2791 3685 2599 5248 8472 1864 5627 6821

78 77 96 115 345 403 215 312 477

2849 3478 3842 3357 3564 10929 10760 12113 3978

-4.7 -7.4 -4.4

2586 3972

-5.3 -4.8 -4.9 -4.7 -4.6 -4.9 -4.8 -5.1 -5.5 -4.8 -5.0 -4.8 -4.5 -4.8 -5.0 -4.8 -5.0 -4.9 -4.8 -5.1 -5.0

4906 4006 4390 4859 4304 4491 3649 3981 4665 4046 4294 3941 4223 3549 4825 3720 4662 4850 5219 4468 4030

45

792

D.11

11;; 376 234 639 144 492 126 2252 1001 183 145 406

454 660 803 801 1105 1043 558 858 15: 371 5611 138% in!

0.66 13.96 0.30 I;.4 0 0.76 6.55 1. 3s 0.98 ?.I>5 ".l! 0.62 0.44 '1 .q?

Vi 89 116 27 102 78 17 105 59

557 489 686 96 1835 3287 291 1903 2115

65!, 691 869 43; 694 554 667 530 72'3

is.:u 15.00 Z./i' 25.95 13.98 22.41 2.96 37 54 14:90

462 285 922 380 324 2348 195 238 865

54 17 25 34 61 66 29 27 65

1053 474 20 523 1104 3949 869 779 719

1079

160

1587

245

1205 1294 1321 1425 907 660 585 1028 839 591 665 616 1139 828 1355 759 1309 1760 1719 1665 1634

136 155 183 140 130 108 166 216 115 137 102 130 150 144 144 121 133 145 176 126 113

NWT

00:: -2.2 -0.2 -1.0 -1.1 --1.6 -3.1 -0.8

1.2 ii.2 0.3 0.2 0.4 0.2 0.7 0.6 0.8 1.0

0.5 0.6 0.3 0":: -0.2 0.6 0.5 -0.4 0.7 -0.7 -0.1 Ellis Bay Formation, Anticosti IslaInd, uebec 215 Hindella umbonata 0.7 218 -0.3 Hindella primstana 221 Hindella sp. -1.0 ORDOVICIAN Ellis Bay Formation, Anticosti IslaInd, uebec x121 Dinorthis anticostiensis 0.7 X3Zl Dinorthis anticostiensis 0.8 X4Zl Dinorthis anticostiensis 1.1 0.9 X621 Dinorthis anticostiensis 221 Dinorthis anticosticnsis 0.9 321 Dinorthis anticostiensis 0.8 421 Dinorthis anticostiensis 1.0 5XZl Dinorthis anticostiensis 0.8 521 0.5 ijinorthis enticostiensis 621 Dinorthis anticostiensis 0.7 722 Dinorthis anticostiensis 0.9 822 Dinorthis anticostiensis 1.1 9XZl Dinorthis anticostiensis 0.9 922 Dinorthis anticostiensis 1.1 lOXZ2 Dinorthis anticostiensis 1.2 1022 Dinorthis anticostiensis 1.0 1421 Dinorthis anticostiensis 0.9 227 Dinocthis anticostiensis 0.9 228 Dinorthis anticostiensls 0.8 229 Dinorthis anticostiensis 0.7 230 DinorthIs anticostiensis 0.9

-6.6 -6.8 -6.7 -7.0 -6.7 -7.2 -6.8 -6.6 -7.2 -7.0 -7.8 -7.5 -5.7 -6.2 -5.8 -6.2 -6.6 -6.4 -5.9 -6.4 -6.7 -6.3 -6.6 -7.0 -6.7 -7.1 -7.3 -7.2 -7.5 -6.9 -7.0

Hi6 750 536 923 754 154s 594 729 147"

i.bl 0.5'1 1.93 3.50 ,3.32 5.75 29.11 2.53 3.21

1197

105b

0.92

2025

115;

3.46

1090 807 752 1101 660 1163 787 936 996 708 665 994 848 674 706 592 736 953 1103 403 1485

1258 1294 1270 1835 1049 1234 1056 1060 1069 1084 1125 1190 1128 101.1 1247 1093 1319 1279 1261 1117 1131

4.44 3.26 1.56 7.26 2,02 7.64 3.40 2.72 4.66 3.54 2.14 4.20 3.22 2.72 2.92 i.22 2.14 l,P'I 2.04 il.17 1'. 30