The Sr, C and O isotopic evolution of Neoproterozoic seawater

Ž .Chemical Geology 161 1999 37–57www.elsevier.comrlocaterchemgeo

The Sr, C and O isotopic evolution of Neoproterozoic seawater

Stein B. Jacobsen a,), Alan J. Kaufman b

a Department of Earth and Planetary Sciences, HarÕard UniÕersity, 20 Oxford St., Cambridge, MA 02138, USAb Department of Geology, UniÕersity of Maryland, College Park, MD 20742, USA

Received 21 August 1998; accepted 30 November 1998

Abstract

Sr and C isotopic data obtained on stratigraphic suites of well-preserved marine limestone from Siberia, Namibia,Canada, Svalbard and East Greenland provide a relatively detailed first-order record of isotopic variation in seawater throughthe late Neoproterozoic Era. This data is used to revise the 87Srr86Sr and d

13C curves of this important interval, duringwhich several discrete global ice ages occurred and the first macroscopic animals evolved. Through this time, the lowest87 86 Ž .Srr Sr values ca. 0.7056 characterize the interval between about 750–800 Ma and have been interpreted to reflect amajor hydrothermal event. From 750 to 600 Ma, the Sr isotope values oscillate between highs and lows, ranging between

Ž . 87 860.7063 and 0.7074. Between 600 Ma and the Early Cambrian ca. 535 Ma , Srr Sr values rise sharply from 0.7063 to0.7087. This is thought to reflect enhanced continental input to the oceans associated with a Pan-African continental

Ž . 13collision. This small subset of limestone samples dolomites dominate the Neoproterozoic record shows the d C curve risesfrom values close to 0 prior to 800 Ma to about q6‰ at 750 Ma and about q8‰ for the time between 600 and 730 Ma.

ŽDuring the time between 600 and 542 Ma, the highest values are about q4‰ higher values in each interval are preserved in.little-altered dolomites . Strong positive-to-negative excursions to values of y5‰ are associated with both Vendian

glaciations estimated at about 575 and 590 Ma and with Sturtian glaciations estimated at about 720 and 740 Ma. In strongcontrast, based on our view of least altered samples, there are no distinct changes in 87Srr86Sr across Neoproterozoic glacialintervals. The duration of these global refrigeration events is a subject of considerable debate. However, consideration of Srresidence times based on elemental partitioning, and the relationship between d

13C and 87Srr86Sr variations, suggest thatthese negative carbon isotope excursions would have lasted at least 350,000 years and no more than about one million years,assuming modern diagenetic fluxes of Sr to the oceans and total absence of continental fluxes. q 1999 Elsevier Science B.V.All rights reserved.

Keywords: Isotopic evolution; Neoproterozoic; Seawater

1. Introduction

Major questions relating to the evolution of theEarth’s oceans, atmosphere, climate and sedimentary

) Corresponding author. Tel.: q1-617-495-5233; fax: q1-617-495-8839; e-mail: [email protected]

shell are most likely resolvable through the study ofhigh resolution isotopic variations in seawaterthrough time. The seawater Sr and C isotopic curvesare also important tools for stratigraphic correlation,in particular for intervals like the NeoproterozoicŽ .544–1000 Ma that lack an adequate biostrati-

Ž .graphic framework Kaufman et al., 1997 .

0009-2541r99r$ - see front matter q 1999 Elsevier Science B.V. All rights reserved.Ž .PII: S0009-2541 99 00080-7

( )S.B. Jacobsen, A.J. KaufmanrChemical Geology 161 1999 37–5738

Our understanding of 87Srr86 Sr variation inseawater through Earth history has improved dra-matically over the past two decades. The secular87Srr86Sr and d

13C trends in Phanerozoic seawaterŽare now well-established Burke et al., 1982; Veizer

.et al., this volume . Extending the Sr isotope recordback in time, this laboratory has focused on detailedstudies of stratigraphic suites from Early Cambrian

Ž .and late Neoproterozoic ca. 800–530 Ma succes-sions in Svalbard, East Greenland, Canada, Siberia

Žand Namibia Derry et al., 1989, 1992, 1994; As-.merom et al., 1991; Kaufman et al., 1993, 1996 .

The demonstration that both the Sr and C isotopiccomposition of seawater varied systematicallythrough this interval allows, for the first time, forma-tion-level correlations worldwide. This stratigraphicframework provides us with the potential to resolvefundamental questions of the relative timing of bio-logical, environmental and climatic events in theNeoproterozoic Earth. In particular, chemical stratig-raphy may allow for the correlation of thick glacialblankets, which are preserved in all regions withmajor Neoproterozoic sedimentary basins. Glacialdiamictites of Neoproterozoic age are recognized at

Žtwo broad levels, identified as the Varangian ca.. Ž600 Ma, Knoll and Walter, 1992 and Sturtian ca.

.700–760 Ma, Hoffman et al., 1998a ice ages. Eachof these in turn consists of at least two discrete

Žpulses of glaciation cf. Young, 1995, but see.Kennedy et al., 1998 for a different point of view .

We have developed simple models to interpretŽthese records cf. Jacobsen, 1988; Asmerom et al.,

.1991; Derry et al., 1992; Kaufman et al., 1993 andwill build on these previous efforts here. The Srisotopic balance of the modern oceans is dominatedby the erosional flux of material from continentalsources, with a smaller but significant contributionfrom mantle sources via submarine hydrothermal

Žsystems Goldstein and Jacobsen, 1987; Palmer and. ŽEdmond, 1989 . At present, other fluxes e.g., diage-

.nesis of marine carbonates are insignificant, butmay have been more important in past oceans. Low87Srr86Sr and high initial 143Ndr144 Nd ratios forArchean chemical precipitates both argue for a much

Žstronger mantle input into the Archean oceans Veizerand Compston, 1976; Jacobsen and Pimentel-Klose,

.1988a,b . There are also strong suggestions of majorhydrothermal events during parts of the Neoprotero-

Ž .zoic Veizer et al., 1983; Asmerom et al., 1991 . Onthe other hand, most of the geochemical cycles thatcontrol the ocean system today were probably in

Ž .place by the early Paleozoic cf. Holland, 1984 .Thus, the Proterozoic must have been a period oftransition from early ‘‘mantle-buffered’’ to modernenvironments.

The purpose of this paper is to highlight theŽ .transitional Neoproterozoic interval by i reporting

our present compilation of data on Sr and C isotopevariations in marine carbonates deposited between

Ž .;520–800 Ma, ii evaluating the data in light ofŽ .diagenetic and or metamorphic alteration, and iii

presenting a more elaborate geochemical box modelfor interpreting Sr and C isotopic variations in sea-water, with particular focus on the periods of global

Ž .glaciation cf. Hoffman et al., 1998b .

2. Geology, stratigraphy and samples

The samples used for this study are from sedimen-tary basins which have been described in detail in

Žearlier publications Knoll et al., 1986, 1995;Fairchild and Spiro, 1987; Kaufman et al., 1991,1993, 1996, 1997; Derry et al., 1989, 1992, 1994;Asmerom et al., 1991; Kaufman and Knoll, 1995;Grotzinger et al., 1995; Pelechaty et al., 1996; Hoff-

.man et al., 1998a,b; Saylor et al., 1998 .Samples of all these sedimentary basins were

typically collected at ;10 m intervals, but in somecases finer scale sampling was achieved. These sedi-mentary successions were used for this compilation

Ž .in part because i well-preserved samples of lime-Ž .stone were available, ii they appear to cover, with

overlap between sections, the entire time span fromthe fossiliferous Early Cambrian back to around 800

Ž .Ma, and iii parallel studies of the paleontology andstratigraphy made collections readily available.

3. Ages

The absolute chronology of few Neoproterozoicsedimentary sequences are well determined and bios-tratigraphy for this time period still provides onlybroad constraints. For assigning absolute ages, wehave relied on a few precise ages obtained by U–Pb

( )S.B. Jacobsen, A.J. KaufmanrChemical Geology 161 1999 37–57 39

isotopic analyses of zircons or baddeleyite. Theseminerals are isolated from volcanic layers interbed-

Žded with sedimentary rocks cf. Bowring et al., 1993,.1998a that could be correlated based on recogniz-

able isotopic features or stratigraphic events. Tocalculate the ages of individual samples, ages that

Žwere assigned to key isotopic events see Table 1 in.Hayes et al., 1999, this volume and were interpo-

lated between or extrapolated beyond tie-points us-ing the known stratigraphic height of samples rela-tive to some reference height.

U–Pb determinations on volcanic ash interbeddedwith these carbonate-dominated successions provide

Ž .direct ages for i the Nemakit–DaldynrTommotianŽ . Ž .boundary at 534"1 Ma Bowring et al., 1993 , ii

the PrecambrianrCambrian boundary at 543.9"1Ž .Ma Bowring et al., 1993; Grotzinger et al., 1995 ,

Ž . 13iii a late Vendian post-glacial positive d C excur-Žsion slightly older than 548.8"1 Ma Grotzinger et

. Ž .al., 1995 , and iv a Sturtian pre-glacial positive13 Žd C excursion around 758.5"3.5 Ma Hoffman et

.al., 1996 . In addition, a U–Pb age of 723"3 Maon a lava flow that caps the Shaler Group of arctic

Ž .Canada Heaman et al., 1992 provides a minimumage constraint for a strong negative d

13C excursionŽ . 87 86Kaufman and Knoll, 1995 and the lowest Srr Sr

Žvalues recorded in the late Neoproterozoic Asmerom.et al., 1991 . A maximum age constraint for this

arctic succession comes from ca. 1.077 Ga Grenvillezircons preserved in sandstone of the lowermost Rae

Ž .Group Rainbird et al., 1996 . However, if the pro-posed correlation of the lower Shaler succession withthe mixed siliciclastic and carbonate succession in

ŽAustralia containing the Rook Tuff is correct Rain-.bird et al., 1996; Kah et al., in press , the maximum

age of Shaler carbonates is no more than 802"10Ž .Ma Fanning et al., 1986 .

With these absolute age constraints in mind, inte-grated chronostratigraphy based on temporal trendsin both C and Sr isotopes, as well as available bio-and lithostratigraphy allow for detailed worldwide

Žcorrelations cf. Kaufman and Knoll, 1995; Kaufman.et al., 1997 .

Ž .For example, Asmerom et al. 1991 correlate theupper Shaler Group in arctic Canada with the lower-most Akademikerbreen Group in Svalbard based on

87 86 Ž .the similarity of Srr Sr compositions ca. 0.7065Žat the top and base of these groups, respectively see

.Fig. 4 later in the text . Significantly, above thecorrelated horizon in both successions is an intervalof negative d

13C. This pre-723 Ma carbon-isotopicŽevent may mark a global ice age cryochron viz.

.Hoffman et al., 1998b although Sturtian glacialdiamictites have not been recognized in either re-gion. Sturtian diamictites with locally developediron-formation are, however, present in both theMackenzie Mountains of Canada and Namibia, andatop each are cap carbonates with negative d

13C and87 86 Žcomparable Srr Sr values 0.7066–0.7068; Kauf-

.man et al., 1997 . A maximum age for the Canadiandiamictite is provided by U–Pb zircon dates of 778–779 Ma for igneous intrusions in underlying forma-

Ž .tions Heaman et al., 1992 . This is consistent with aU–Pb zircon age date on a granite clast within theiron-formation bearing diamictite, indicating that this

Žice age is no older than 755"18 Ma Klein and.Beukes, 1993; Ross et al., 1995 . In sum, the avail-

able radiometric constraints suggest that this globallycorrelated negative carbon isotope excursion likelyoccurred between 720 and 760 Ma. In order to assignindividual ages to samples, we chose a 740 Ma agefor this event.

Recent sequence and chemostratigraphic studieshave demonstrated that a second and younger di-

Žamictite occurs in the Otavi succession Hoffmann.and Prave, 1996; Hoffman et al., 1998a,b ; signifi-

cantly, both Otavi diamictites are bounded by 13C-enriched carbonates below and 13C-depleted carbon-ates above. The younger diamictite, some 400 mabove the older, and its associated negative carbonisotope excursion is also considered to be SturtianŽ .Hoffman et al., 1998a,b and given an age estimateof 720 Ma. Available radiometric constraints providean upper bound for the whole Otavi Group at around

Ž .600 Ma Stanistreet et al., 1991 .In many Vendian sections, particularly in the

north Atlantic region, there is also lithologic evi-dence for two pulses of Varangian diamictites. Eachof these is associated with a negative carbon isotope

Žexcursion Knoll et al., 1986; Fairchild and Spiro,.1987; Kaufman et al., 1997 . In the scheme pre-

sented here, the younger of the two events is esti-mated around 575 Ma, while the older is estimated

Ž .around 590 Ma cf. Saylor et al., 1998 . Varangianglacial strata in Avalon post-date a volcanic unitthousands of meters below the diamictite with a


ŽU–Pb age of 606q3.7ry2.9 Ma Krogh et al.,.1988 . New U–Pb zircon measurements from other

volcanics beneath the Avalonian diamictite suggestthat the Varanger ice age may be younger than 595

Ž .Ma Bowring et al., 1998b .In the Mackenzie Mountains, evidence for

Varangian glaciation is preserved in the Ice BrookŽ .diamictite Aitken, 1991 ; notably, the overlying capŽ .dolomite Tepee Formation and limestone in the

lower Sheepbed Formation record significantly nega-13 Ž .tive d C values Narbonne et al., 1994 . This event

is correlated with the older Varanger ice age andaccordingly given an estimated age of 590 Ma.Stratigraphically higher in the Sheepbed d

13C risesto q7‰ falls sharply to near 0‰ and then returns tosignificantly positive values near the top of the for-

Ž .mation. Kaufman et al. 1997 provisionally interpretthis strong mid Sheepbed negative shift as a markerfor the younger Varanger glaciation at 575 Ma,although diamictites are not preserved at this level.

A final notable negative carbon isotope excursionis recognized at the Precambrian–Cambrian bound-

Žary Narbonne et al., 1994; Kimura et al., 1997;.Bartley et al., 1998 , but significantly, at this level,

Žthere is no convincing evidence for glaciation but.see Bertrand-Sarfati et al., 1995 . Based on absolute

radiometric determinations from Siberia and Nami-bia, this isotopic event is estimated to have lasted nomore than one million years, between 543 and 544

Ž .Ma Pelechaty et al., 1996 .The isotopic database with age estimates for indi-

vidual samples is available on request from theŽauthors it is also available on the web: http:rr

.geochemistry.harvard.edur .

4. Simple models of fluid–rock interaction

The isotopic compositions of carbonates can besubstantially changed from primary seawater equili-brated values by meteoric diagenesis, dolomitization,and metamorphism. Various models for fluid–rock

Žinteraction e.g., Taylor, 1977; Nabelek, 1987; Ban-.ner and Hanson, 1990; Schrag, this volume have

been used to established parameters for determiningthe degree of alteration of Sr, C and O isotopiccompositions in carbonates.

The response of the concentration of an element iŽ .C to fluid–rock exchange depends on the initiali

Ž w0 .concentration of the element in the fluid C andiŽ r 0.the rock C , on the effective fluid–rock distribu-i

tion coefficient:

C ri

D s 1Ž .i wCi

Žand on the relative proportions of the fluid mass. Ž .fractionsF and rock mass fractions1yF in-

volved in the exchange. The most commonly usedparameter to express the degree of fluid–rock inter-action is the weight ratio of water or fluid to rock:

Ž .hsFr 1yF .ŽThe response of stable isotopic systems such as

13 18 .d sd C and d sd O to fluid–rock exchangeC O

depends on the initial isotopic composition of theŽ w0 . Ž r 0.fluid d and the rock d , as well as on thei i

water–rock fractionation factor:

D s d r yd w . 2Ž .Ž .i i i

Here, D and D may depend on temperature asi i

well as on other factors. Also, the degree of ap-proach to equilibrium is important. For non-equi-librium conditions, the equations given here are stillvalid, however, D and D are effective fractionationi i

factors that may be determined empirically or bysome kinetic model. The simple theory below is forconstant D and D values but can easily be general-i i

ized to variable values.For a closed system, which is the most straight-

forward case, the conservation relation for concentra-Ž .tions i.e., Sr, Mn, O, C is:

C r 0 qhC w 0i irC h s 3Ž . Ž .i 1q hrDŽ .i

where C r and C w are the concentrations of thei i

particular element in the rock and the water, respec-rŽ .tively. For stable isotopes, where d h is the finali

isotopic composition of the rock we have:

C r 0i r 0 w 0d qh d qDŽ .i i iw 0ž /Cird h s . 4Ž . Ž .i r 0Ci

hq w 0ž /Ci

ŽFor radiogenic isotope ratios such as a sSr87 86 .Srr Sr this is simplified and the radiogenic iso-


Ž r .tope ratio a varies with h according to the fol-i

lowing equation:

C r 0i r 0 w 0a qhai iw 0ž /Cira h s . 5Ž . Ž .i r 0Cihq w 0ž /Ci

Since h-values in the range of 1 to 10 are oftenŽobtained for altered limestone cf. Banner and Han-

.son, 1990 the closed system model is not veryrealistic. Therefore, a simple open-system model isneeded. We assume that only an infinitesimal amountof fluid is in the rock at any time. For such condi-tions, the following differential equations govern

Ž .open system exchange: i for concentrations:

r tdC C d Fi iw 0s C y , 6Ž .id t D d ti

Ž .ii for stable isotopic compositions:

r w 0dd C d Fi i w 0 rs d qD yd , 7Ž .i i irž /d t C d ti

Ž .and iii for radiogenic isotopes:

r w 0da C d Fi i w 0 rs a ya . 8Ž .i irž /d t C d ti

The solutions to the above open system differen-tial equations are:

hr w 0 r 0 w 0C h sD C q C yD C exp y 9Ž . Ž .Ž .i i i i i i ž /Di

d r hŽ .i

r 0C hi r 0 w 0d qD exp y1 D qdi i i iw 0 ž /ž / DC iis

r 0C hiqD exp y1iw 0 ž /½ 5ž / DC ii

10Ž .

r 0C hi r 0 w 0a qa D exp y1i i iw 0 ž /ž / DC iira h s .Ž .i r 0C hiqD exp y1iw 0 ž /½ 5ž / DC ii

11Ž .

Fig. 1 shows a calculation of the pattern forlimestone–fluid interaction for d

18 O, 87Srr86Sr andd

13C with varying water to rock ratio for both closedand open system interaction. Both models give rela-tively similar results for C and O, but there is arelatively large difference between the two modelsfor 87Srr86Sr. This is due to the large change in Srconcentration of the limestone during water–rockinteraction. For evaluating 87Srr86Sr, it is essentialto consider the open system model. The calculationsshow that h)1 is required for the d

18 O value to betotally equilibrated with the fluid. For 87Srr86Sr andd

13C, h-values )10 and )100 are required, re-spectively, to equilibrate the rock with the fluid. Fig.2 shows a calculation of the expected pattern for

Fig. 1. Evolution of d18 O, 87Srr86 Sr and d

13C with varying waterto rock ratio for both closed and open system interaction. Lime-stone initial composition: d

18 Os0, 87Srr86 Srs0.706 and d13C

s5, Srs1000 ppm. Initial fluid composition: d18 Ow 0 q D sO

y14, 87Srr86 Srs0.710 and d13C w 0 q D s y12, Sr s50C

ppm, C s1000 ppm. D values: D s1, D s0.54, D s120.Sr O C


Fig. 2. Expected evolution during fluid–rock interaction of 87Srr86 Sr and d13C with varying MnrSr and d

18 O for both closed and opensystems. Values used are the same as those in Fig. 1 except for values for Mn. Initial limestone composition: Mns1 ppm. Initial fluid

Ž .composition: Mns4.1 ppm. D values: D s600 cf. Banner and Hanson, 1990; Derry et al., 1992 .Mn

limestone–fluid interaction for 87Srr86Sr and d13C

with varying MnrSr and d18 O for both closed and

open system interaction. During marine as well asmeteoric diagenesis of a limestone Mn increases andSr decreases, therefore the MnrSr ratio is generallyconsidered a good indicator of the degree of alter-ation. Except for the plot of d

13C vs. MnrSr, highlynon-linear trends resulted. It is clear that even forrelatively low MnrSr of ;1 large shifts in isotopiccomposition may have taken place.

5. Identification of primary isotopic signatures

Many samples of Proterozoic carbonates are al-tered to some degree, and therefore, cannot be usedto determine secular trends in 87Srr86Sr and d

13C of

coeval seawater. Various investigators establishedparameters for determining the degree of alteration

Žof Sr and C isotopes in Precambrian carbonates e.g.,Brand and Veizer, 1980, 1981; Veizer et al., 1983;Derry et al., 1989, 1992; Asmerom et al., 1991;

.Kaufman et al., 1993 . We believe that our earlierisotopic studies show that meaningful results con-straining the chemical and isotopic evolution of sea-

Ž .water can be obtained provided i samples aremicro-drilled in order to obtain the least-altered ma-

Ž .terial, from each, and ii samples are screened foralteration using chemical and isotopic parameters.

The calculations in Fig. 2 are compared withvariably altered samples from Namibia, Australiaand Svalbard in Fig. 3. The trends indicated byarrows are those expected from fluid–rock interac-tion with increasing water–rock ratio. Mixing of the


ŽFig. 3. Comparison of calculations are shown in Fig. 2 with variably altered samples from Namibia, Australia and Svalbard Derry et al.,.1992 . The trends indicated by arrows are those expected from fluid rock interaction with increasing water–rock ratio. Mixing of ‘‘primary’’

and ‘‘diagenetic’’ endmembers along these paths yield roughly linear arrays in the diagrams involving 87Srr86 Sr. Examination of thesesamples show that they consist of such mixtures. Therefore, micro-drilling was used to obtain the least altered materials. These all hadMnrSr-2 and d

18 O)y10.

‘‘primary’’ and ‘‘diagenetic’’ end-members of thesepaths will yield roughly linear paths in the diagramsinvolving 87Srr86Sr. Close examination of thesesamples suggest that they consist of such mixtures.Therefore, micro-drilling was used to obtain the leastaltered materials. These were all found to haveMnrSr-2. In general, samples with MnrSr)2 arealmost always found to have altered Sr isotopiccompositions, while those with lower values yieldedquite consistent results. From Fig. 3, one might evenexpect such samples to have highly altered 87Srr86Srif they followed the calculated water–rock interac-tion trends. We suggest that these represent mixturesof relatively primary and diagenetic grains. If cor-rect, much smaller shifts in the 87Srr86Sr for asample with a MnrSr ratio of -2 would be ex-pected. While it is clearly important to characterizethe degree of alteration of samples using elementaland isotopic parameters, we also check for consis-

tency in isotopic values between closely spaced sam-ples. With respect to construction of the 87Srr86Srtemporal curve, we use only analyses of those lime-stones which have passed all of our petrographic,

Ž . Ž 18 87 86 .elemental MnrSr , and isotopic d O, Rbr SrŽtests cf. Asmerom et al., 1991; Derry et al., 1992;

.Kaufman et al., 1993 . In addition, for the isotoperecords presented below, we only used samples with

Ž y3 .low RbrSr ratios -5=10 and high Sr concen-Ž .trations 150–2500 ppm because samples not meet-

ing these criteria often yield inconsistent results.

6. Neoproterozoic 87Srrrrrr86Sr isotope record

Ž .Veizer and Compston 1976 obtained a seawater87Srr86Sr curve for the whole Precambrian and

Ž .Veizer et al. 1983 later improved on this record forŽ .the Neoproterozoic ;1000–540 Ma . These au-


thors observed a sharp increase in the 87Srr86Sr ratioof marine carbonates between 2.5–2.0 Ga ago that isconsistent with a significant decrease in the post-Archean hydrothermal flux of Sr to seawater. Thisisotopic event may be closely linked to the agglom-eration of an early supercontinent and a concomitantincrease in the flux of more radiogenic Sr from acontinental source. On the other hand, low 87Srr86Srvalues in ca. 900 Ma carbonates from the Adrar ofMauritania were interpreted as the ocean’s responseto a hydrothermal event, most likely associated with

Žbreakup of a Neoproterozoic supercontinent viz..Rodinia .

Based on the screened database discussed above,we estimate initial 87Srr86Sr ratios in seawater fromour best-preserved samples. This compilation signifi-cantly improves the resolution of the 87Srr86Sr record

Žfor the late Neoproterozoic to Early Cambrian Fig..4 . The dotted vertical lines in this figure show the

absolute or estimated ages of the Cambrian–Pre-

Ž . Ž .cambrian boundary C–Pc , the upper V1 and lowerŽ . Ž .V2 Vendian glaciations, the upper S1 and lowerŽ .S2 Sturtian glaciations, and a possible early Stur-

Ž .tian glaciation S3 .Specifically, the Shaler Group data indicate

smooth and systematic variation in 87Srr86Sr be-Ž .tween ca. 800–750 Ma range: 0.7068–0.7056 , with

the lowest 87Srr86Sr value of 0.7056 at ca. 775 Ma,Žinterpreted as a major hydrothermal event Asmerom

.et al., 1991 . In that publication, it was suggestedthat the Shaler event might correlate with a similarlow Sr isotopic event in the Adrar of Mauritania.However, Rb–Sr age determinations on glauconitesfrom the Mauritanian sediments indicate an agearound 900 Ma, suggesting that hydrothermal domi-nation of ocean water Sr continued back through the

Žearly Neoproterozoic Kaufman et al., 1998; Bartley.et al., unpublished data .In fact, similarly low

87Srr86Sr values characterize limestone samples de-posited across the Neoproterozoic–Mesoproterozoic

87 86 ŽFig. 4. Temporal variations of Srr Sr in Neoproterozoic carbonates Derry et al., 1989; Derry et al., 1992; Asmerom et al., 1991;.Kaufman et al., 1993; Derry et al., 1994; Kaufman et al., 1996 from Siberia, Namibia, Svalbard and Canada. Dotted vertical lines represent

the estimated ages of key boundaries or the termination of glacial events as discussed in the text: C–PcsCambrian–Precambrian boundary,V1 and V2supper and lower Vendian glaciations, S1 and S2s upper and lower Sturtian glaciations, S3sa possible early Sturtianglaciation.


boundary in the Turukhansk Uplift region of north-Žeastern Siberia Gorokhov et al., 1995; Bartley et al.,. 87 86unpublished data . From 730 to 600 Ma, Srr Sr

values fluctuate around 0.707, rising from near0.7068 to as high as 0.7075 through the interval ofSturtian glaciation. Well-preserved and high Sr caplimestone atop the older Sturtian diamictites world-wide have 87Srr86Sr between 0.7066 and 0.7068,while similarly preserved samples from carbonatesabove the second Sturtian diamictite have values

Žnear 0.7074 Kaufman et al., 1997; Misi and Veizer,.1998; Kaufman, unpublished data . Above the

younger Sturtian glacial, 87Srr86Sr declines in ourdata set to values as low as 0.7068 prior to the first

Ž .Varangian diamictite. Fairchild et al. in press , ana-lyzed additional samples from immediately below

Žthe older Varanger diamictite in East Greenland cf.. 87 86Derry et al., 1989 and suggest that Srr Sr values

may have dropped to as low as 0.7063 before that iceage. Across the first Varangian ice age, there isapparently little change in seawater 87Srr86Sr, with a

Žbest-preserved intertillite value of 0.7068 Kaufman.et al., 1997 . However, across the second Varanger

event there is a marked rise in 87Srr86Sr with valuesin well-preserved high Sr cap carbonates near 0.7081Ž .Kaufman et al., 1993; Saylor et al., 1998 . Thesehigh values, which continues through a plateau near0.7085 to the Precambrian–Cambrian boundary, andthen rise up to 0.7087 in Nemakit–Daldyn beds of

Ž .the Early Cambrian Kaufman et al., 1996 . Thestrong Vendian rise in 87Srr86Sr is thought to reflectenhanced continental input to the oceans associatedwith a Pan-African continental collision.

In sum, based on the available data from well-pre-served limestone samples, it appears that with theexception of the younger Varanger ice age, the Neo-proterozoic glaciations had little effect on the Srisotope composition of the oceans. With respect tochronostratigraphic correlation, Sr isotope composi-tions alone allow us to clearly identify the older

Ž .Sturtian ca. 0.7066–0.7068 and younger VarangerŽ .ca. 0.7081 cap carbonates. It is admittedly moredifficult to discriminate between the younger Stur-

Ž . Ž .tian ca. 0.7074 and older Varanger ca. 0.7068Ž .post-glacial caps, a point that Kennedy et al. 1998

marshall to claim that there is only one Sturtian andone Varanger event. These authors highlight the factthat in any single Neoproterozoic succession there is

clear lithostratigraphic evidence for only two levelsof glacial diamictite. In the absence of lithologicevidence, however, strong positive-to-negative strati-graphic trends in d

13C abundances of marine carbon-ates may be regarded as biogeochemical markers of

Ž .glaciation Kaufman et al., 1997 . We contend thatthe absolute magnitude of least-altered 87Srr86Sr

Žvalues, which are regionally consistent cf. Saylor et. 13al., 1998 , in the C-depleted cap carbonates support

the view that at least four ice ages insulted theNeoproterozoic Earth. This conclusion is further sup-ported by available biostratigraphy and d

13C che-mostratigraphic trends, which indicate that values of)q10‰ characterize carbonates deposited betweenrecognized Varanger and Sturtian diamictites, thusproviding an isotopic divide between Neoproterozoicglacial epochs. To date, similarly high d

13C valueshave not been recorded in open marine carbonates

Žfrom Varanger and younger successions Smith et.al., 1994, 1996 .

6.1. Comparison with other studies of Neoprotero-zoic Sr isotope Õariations

Sr isotopic compositions of late Neoproterozoiccarbonates from Australia, Mongolia, Siberia, Oman,South Africa and Brazil have also been published inthe literature but are not included in the presentcompilation. In some cases these successions can bechronostratigraphically matched with our results, butin others, poor age constraints or lack of stratigraphicvariations make detailed correlation problematic.

In Australia, Sr isotopes have been measured oncarbonates in the Bitter Springs, Skillogalee and

Ž .Brighton formations Veizer and Compston, 1976 .Diagenetic tests have not been run on these samples,so it is difficult to evaluate whether these representprimary compositions or diagenetic artifacts. Re-gional lithostratigraphic correlations indicate that allthese units are younger than the 802"10 Ma Rook

Ž .Tuff Fanning et al., 1986 . According to the mostrecent correlation of Neoproterozoic strata in Aus-

Ž .tralia and Canada Rainbird et al., 1996 , the BitterSprings and Skillogalee formations are broadlyequivalent to the Shaler Group, while the Brighton iscoeval with the Keele Formation in the MackenzieMountains. Strontium isotope compositions of lime-stone and dolomite samples from the Bitter Springs


Formation range from 0.7068 to 0.7116. The lowest87Srr86Sr value from this unit does approximateuppermost Shaler Group values shown in our compi-

Ž .lation Asmerom et al., 1991 , but considering theirrelatively high Rb content most of these analysesappear to reflect secondary compositions. Similarly,all samples from the Skillogalee Formation appearaltered, recording values between 0.7091 to 0.7134.Limestone samples from the Brighton Formationrange from 0.7075 to 0.7092; significantly, the low-est values are only slightly higher than those recordedin the proposed equivalent Keele FormationŽ .Narbonne et al., 1994 used in our compilation.

Ž .Brasier et al. 1996 report a large number of87Srr86 Sr measurements of variably preserveddolomitic limestone and dolomite samples from theNeoproterozoic and Lower Cambrian Tsagaan Oloomand Bayan Gol formations in Mongolia. Well-pre-served limestone samples directly overlying a Stur-

Ž .tian-aged diamictite Maikhan Uul Formation showa distinct rise from 0.7067 to 0.7072 within 50 m ofsection. The low values at the base of this post-gla-cial carbonate are an exact match for limestonesfrom Sturtian cap carbonates in the Mackenzie

Ž .Mountains and Namibia Kaufman et al., 1997 , bothplotted in our compilation. The upward trend in Srisotope compositions through this relatively thin in-terval is also consistent with variations recorded inNamibia and Svalbard through a thicker stratigraphic

Žyet broadly equivalent interval cf. Derry et al.,.1989 . Similarly, the rapid lower Vendian rise in

87Srr86Sr from values near 0.7070 to 0.7085, docu-Ž .mented by Kaufman et al. 1993 with integrated

data from several basins, is preserved in the Mongo-lian succession. Close comparison of 87Srr86Sr val-ues from the Lower Cambrian interval in Mongolia

Ž .reported by Brasier et al. 1996 and in SiberiaŽ .Kaufman et al., 1996 further support the trendcompiled here.

Thick carbonate successions in the southern UralMountains along the Lena River have been the focusof recent carbon and strontium isotopic studies aswell. The age of these sediments, however, are onlybroadly constrained by paleonotologic and lithos-tratigraphic correlation coupled with uncertain RbrSrisochrons based on glauconite minerals. Gorokhov et

Ž .al. 1995 present data from carbonates of the Zhuyaand Yudoma groups, interpreted as Neoproterozoic

in age, as well as from the Macha and Tolbachanformations, which, based on abundant Atdabaniansmall shelly fossils, are Lower Cambrian in age.While all these samples are limestone with lowMnrSr, most had elevated Rb or low Sr concentra-tions and would have been rejected with the diage-netic screens discussed above. These Neoproterozoicunits preserve 87Srr86Sr values between 0.7081 and0.7094; the lowest values are consistent with Ven-dian carbonates worldwide, but not with older Riph-ean rocks as suggested by some Russian stratigra-

Ž .phers cited in Gorokhov et al., 1995 . In anotherregion of the southern Urals along the Sim and Ufarivers, Sr isotope compositions of limestone samplesfrom the Inzer Formation have recently been re-

Ž .ported by Kuznetsov et al. 1997 . A Pb–Pb date of836"27 Ma on limestone in the lower half of thissuccession provides a maximal depositional age con-

Ž .straint Ovchinnikova et al., 1995 , but an upper-bound for this unit is poorly defined by radiometricmeans. However, the uniquely low and consistent87Srr86 Sr values from well-preserved limestone

Žsamples from the Inzer ranging between 0.7053 and.0.7056 allow for direct comparison with Shaler

Group analyses plotted in Fig. 4. A recent lowresolution study of carbon isotope variations in the

Ž .Inzer Formation Podkovyrov et al., 1995 also sug-gests a broad similarity between the Russian andCanadian successions. On the other hand, availableconstraints do not preclude the possibility that Inzersediments might be slightly older than the Shalersuccession. If correct, this observation further sup-ports the view that the hydrothermal domination ofca. 750 Ma seawater Sr was not an isolated event,but continued back through the early Neoproterozoicand beyond.

In Oman, C and Sr isotopes have been measuredfor Vendian carbonates from the Huqf Group overly-ing diamictites in the basal Abu Mahara Formationbelieved to be equivalent to Varanger glacial stratain the north Atlantic region. Volcanic rocks underly-ing the diamictites in Oman have been dated by

ŽRb–Sr techniques at 554"10 Ma Dubreuilh et al.,.1992 , but the systematics have not been published.

Given the inconsistency between this age and abso-lute U–Pb zircon constraints on Varangian glaciationŽ .Bowring et al., 1998a,b; Krogh et al., 1988 we

Žconsider that the Rb–Sr age is reset cf. Kaufman


. 13and Knoll, 1995 . Close comparison of d C secularŽ .trends in the Huqf Group Burns and Matter, 1993

with radiometrically well-constrained curves inŽNamibia Kaufman et al., 1991; Grotzinger et al.,

.1995; Saylor et al., 1998 indicate that Buah andolder strata in Oman must predate 549"1 Ma. Srisotope compositions of samples from this succes-

87 Žsion show a strong secular increase in Sr ca..0.7082 to 0.7094; Burns et al., 1994 , which is

similar, but offset to higher 87Srr86Sr values likelydue to the analysis of dolomite, to the early Vendianpost-Varanger rise in 87Srr86Sr indicated in ourcompilation.

Ž .Veizer et al. 1983 also reported Sr isotope abun-dances for late Neoproterozoic carbonates in theKango and Malmebury formations in South Africa.

ŽThe age of these units is poorly constrained Knoll etal., 1986 suggest an age of 750"150 Ma for this

. 87 86unit . However, the Srr Sr values of these lime-stones are broadly consistent with pre-Varanger suc-

Ž .cessions in Svalbard Derry et al., 1989, 1992 , theŽ .Mackenzie Mountains Narbonne et al., 1994 and

Ž .Namibia Kaufman et al., 1997 .Ž .Finally, Misi and Veizer 1998 recently reported

87Srr86Sr compositions of limestone samples of theBambui and Una groups above a Sturtian diamictitein Brazil. Lacking reliable radiometric age con-straints, these authors used Sr isotope chemostratig-

Ž .raphy to suggest an age of 600 to 670 Ma for theseunits on the Sao Francisco craton. Well-preservedlimestones with high Sr abundances yielded valuesbetween 0.7074 and 0.7078, with an upsection trendto lower values. Like the Kango and Malmesbury

Žformations mentioned above, the strontium and car-.bon; see Iyer et al., 1995 isotope values are broadly

87 86 Ž 13 .comparable with Srr Sr and d C measure-ments in the pre-Varanger successions presented here.

In sum, data from other late Neoproterozoic basinsconfirm the first-order trends of our compilationbased on our analyses of samples from Svalbard andEast Greenland, Siberia, Namibia and Canada. Fu-ture revisions based on better absolute age con-

Ž .straints cf. Bowring et al., 1998a,b for Neoprotero-zoic sedimentary succession, as well as high densitysampling will undoubtedly improve the resolution ofSr isotope chemostratigraphic schemes. We suggest,however, that this scheme should never be used inisolation, but integrated with all other available

stratigraphic tools in order to best interpret Neopro-terozoic Earth history.

7. Neoproterozoic d13C and d18 O isotope records

Ž .Schidlowski et al. 1975 obtained a Precambriand

13C record, which suggested that Proterozoic car-bonates lacked strong temporal variation. However,more recent work has demonstrated that from about850 Ma ago until the end of the Proterozoic Eon,d

13C fluctuated with frequencies comparable toPhanerozoic variations but with far greater magni-

Ž .tudes reviewed in Kaufman and Knoll, 1995 .The d

13C and d18 O values for the same carbonate

samples used to construct the 87Srr86Sr curve in Fig.4 are shown in Figs. 5 and 6, respectively. The solidcurve shown in Fig. 5 is based on additional d

13CŽ .data see Hayes et al., 1999, this volume that would

not pass the screening tests used for the Sr isotopiccurve. Symbols and dotted vertical lines are the sameas in Fig. 4. The d

13C curve shown in Fig. 5provides a relatively detailed record of isotopic vari-ation in seawater through the Neoproterozoic andEarly Cambrian, and permits direct correlation of theSr and C isotopic changes for this time period.

The d13C curve rises from values close to 0‰

prior to 800 Ma to about q6‰ at 750 Ma and aboutq8‰ for the time between 600 and 730 Ma. Duringthe time between 600 and 542 Ma, however, thehighest values are about q4‰. Brief negative excur-sions to values of y5‰ are associated with Vendianglaciations at about 575, 590 Ma and with Sturtianglaciations about 720 Ma and 740 Ma. In contrast, asshown in Fig. 4, there are no distinct changes in87Srr86Sr associated with glacial intervals. Duringthe late Vendian, d

13C values fall from q4 to y4‰at the Precambrian–Cambrian boundary and thenfluctuate between y2 and q2‰ in the lower Cam-brian.

d18 O values for the same samples are most likely

of all the isotopic values to be altered from primaryvalues as discussed above. Almost all samples have

18 Ž .d O values below y5‰ Fig. 6 . Thus, the highest,and least likely to be altered, values are similar to

Žearly Paleozoic values reported by Veizer et al. this.volume . However, we cannot rule out that the sam-

ples have all been altered from values similar to


Fig. 5. Temporal variations in d13C of carbonates for the same samples used for Sr isotopes in Fig. 4. Symbols and dotted vertical lines are

Ž .the same as in Fig. 4. Solid curve is based on additional data reviewed in Kaufman and Knoll, 1995; Kaufman et al., 1997 .

18 Ž .Fig. 6. Temporal d O isotopic variations relative to the PDB standard in the same samples of Neoproterozoic carbonates as shown in Fig.Ž .5 and 6 . Symbols and dotted vertical lines are the same as in Fig. 4.


Ž 18 .those of modern carbonates d O;0‰ . While thePhanerozoic record suggest such shifts can be realŽ .Veizer et al., this volume , at this time the samecannot be stated with the same confidence for thelate Neoproterozoic.

8. Box model for Sr and C cycles

A box model for evaluating 87Srr86Sr and d13C

variations in Neoproterozoic carbonates is shown inFig. 7. This is a more elaborate version of a modelfor the Sr and C cycles that we have used earlierŽGoldstein and Jacobsen, 1987; Jacobsen, 1988; As-merom et al., 1991; Derry et al., 1992; Kaufman et

.al., 1993 . The main difference is for the C cycle.Previously, we used a model that primarily involvedthe crustal rock reservoirs, with the ocean as a rapidmixer. This was sufficient for evaluating variationsin C-burial rates in sediments. We are now interestedin applying this model to shorter term variations dueto Neoproteozoic glacial intervals. Therefore, themore elaborate model shown in Fig. 7 was adopted.This new model considers all the same reservoirsand fluxes for both the Sr and the C cycles. Tocouple the Sr and C cycles in the earlier model, itwas simply assumed that the erosion rate for Sr is

Žproportional to that of organic C Derry et al., 1992;.Kaufman et al., 1993 . The new model involves

Ž .exchanges between five reservoirs: seawater SW ,Ž .depleted mantle DM including oceanic crust, and

Ž .three continental crustal CONT reservoirs

Fig. 7. Cartoon of box model of the Sr and C cycle used forevaluating 87Srr86 Sr and d

13C variations in Neoproterozoic car-bonates.

ŽcarbonatessCARB, organic carbon in sedimentss.ORG and silicate rockssSIL .

The starting point is to consider the balance equa-Ž SW .tion for the total amount N of a chemicaliŽ .element or a stable isotope i in seawater. The rate

of change of this amount is the difference betweenŽ .the sum of all the input fluxes from reservoirs j to

Ž i . Ž i .seawater J and the output fluxes J .j – SW SW – j

d N SWi i is J y J . 12Ž .Ý Ýj – SW SW – jd t j j

Combining this equation for 87Sr and 86Sr, itfollows that the balance equation for the ratio of

j Ž87 86 . j Žthese isotopes a s Srr Sr is cf. Brass,Sr.1976 :

da SWSrSW Sr j SWN s J a ya 13Ž .Ž .ÝSr j – SW Sr Srž /d t j

since there is no significant isotopic fractionationbetween seawater and its outputs.

For the carbon isotope balance, d sd13C, thereC

are, however, additional terms due to carbon isotopeŽ j SW .fractionation D sd yd between seawaterj – SW C C

and its output fluxes. These are all small except forthe kinetic isotope fractionation associated withphotosynthesis. Thus, for simplicity, we ignore all ofthese effects except for the large isotope fractiona-tion effect in organic matter buried in sedimentsŽ ORG SWD sd yd ;y25; see Hayes et al.,ORG – SW C C

.1999, this volume . It follows that the carbon isotopebalance equation is:

dd SWCSW C j SWN s J d ydŽ .ÝC j – SW C Cž /d t j

yJ C D 14Ž .SW – ORG ORG – SW

We now introduce the following into the above Srand C balance equations:Ž .i the mass fraction of Sr in seawater arrivingfrom source reservoir j: xSr sJ Sr rÝ J Sr ,j – SW j – SW j j – SWŽ .ii the mass fraction of C from source reservoir j:xC sJ C rÝ J C ,j – SW j – SW j j – SWŽ .iii the residence times of C and Sr in seawatert SW sN SWrÝ J C and t SW sN SWrÝ J Sr ,C C j j – SW Sr Sr j j – SW


andŽ .iv the mass fraction of C buried as organiccarbon relative to the total input rate of C to theocean: yC sJ C rÝ J C sJ C tSW – ORG SW – ORG j j – SW SW – ORGSWrN SW.C C

They can then be written in the following form:

da SWSrSW SW Sr jt qa s x a 15Ž .ÝSr Sr j – SW Srž /d t j

dd SWCSW SW C jt qd s x dÝC C j – SW Cž /d t j

yyC D . 16Ž .SW – ORG ORG – SW

To further simplify these equations, we note thatfor the mass fractions of the continental input to theoceans we have for Sr and C:

xSr a CO N T sxSr a SILCONT – SW Sr SIL – SW Sr

qxSr a CARBCARB – SW Sr

qxSr a ORG 17Ž .ORG – SW Sr

xC d CONT sxC d CARBCONT – SW C CARB – SW C

qxC d ORG. 18Ž .ORG – SW C

We also have the following relationship betweenŽcontinental fluxes and hydrothermal depleted man-

.tle fluxes:

xSr s 1yxSr 19Ž .Ž .DM CONT

d DM sxC d DM qxC d CONT. 20Ž .C DM – SW C CONT – SW C

The last equation is approximately valid sinceCONT DM Žestimates for d and d are both ;y5.5 cf.C C

.Derry et al., 1992 .Using all these relationships for mass fractions,

we can simplify to:

da SWSrSW SWt qaSr Srž /d t

sa DM qxSr a CONT ya DM 21Ž .Ž .Sr CONT – SW Sr Sr

dd SWCSW SW DM Ct qd sd yy D .C C C SW – ORG ORG – SWž /d t

22Ž .

The key parameters are xSr for Sr andCONT – SW

yC for C and these are the only importantSW – ORG

parameters for times sufficiently short that a CONTSr

does not vary significantly. Thus, given records ofd SW and a SW, both xSr and yC mayC Sr CONT – SW SW – ORG

be evaluated. For a complete evaluation of the five-box model, a set of such equations is needed for eachreservoir. To solve the complete system is beyondthe scope of this paper, however. In Section 9, wewill use these simple results for the seawater reser-voir to evaluate isotopic changes associated with theinterval of late Neoproterozoic glaciations. This newmore elaborate model, however, does not lead tomajor changes in our earlier interpretations of the

ŽNeoproterozoic Sr and C isotope records Jacobsen,1988; Asmerom et al., 1991; Derry et al., 1992;

.Kaufman et al., 1993 .The relationship to our previous models are as

follows: based on 87Srr86Sr of seawater, we evalu-Ž Sr .ated variations in the river water flux JCONT – SW

Ž .the erosional flux of Sr from the continents to theŽ Sr .hydrothermal flux J of Sr through oceanDM – SW

ridges. This has the following simple relationship toxSr , in Sr isotope balance equation above:CONT – SW

J Sr xSrCONT – SW CONT – SW

s . 23Ž .Sr Srž /J 1yxDM – SW CONT – SW

Similarly, we evaluated the burial rate of organicŽ C .C J , which shows a simple relationship toSW – ORG

Žoverall erosion rates estimated from the Sr isotopic. 13record and secular variations in d C. This parame-

ter is simply related to yC in the C isotopeSW – ORG

balance equation above:

yC N SWSW – ORG CCJ s . 24Ž .SW – ORG SWtC

Such considerations of the 87Srr86Sr record ofseawater showed that since the global dissolved fluxof Sr in rivers is proportional to the global erosionrate, the J Sr value constrained this way is aCONT – SW

proxy of erosion rates through time. We note thatthis does not necessarily imply that 87Srr86Sr inseawater is a proxy for erosion rates. Most likely, the87Srr86Sr of seawater is a function of both varyingerosion rates as well as changes in the 87Srr86Sr ofthe river flux. Model results indicated three distinctepisodes of high global continental erosion rates dueto uplift caused by continental collision at these


Ž .times: ;0 Ga Himalayan–Tibetan collision , ;0.4Ž .Ga Caledonian–Appalachian collision and ;0.6Ž .Ga Pan-African collision ago. This analysis sug-

gested that there were two times of very rapid in-87 86 Ž .crease in Srr Sr 0 and 0.6 Ga , which appear to

be caused by high erosion rates. The third eventŽ .;0.4 Ga is, however, not associated with particu-larly high 87Srr86Sr. Thus, there is, in general, alimited correlation between the Sr isotope curve andhigh vs. low erosional fluxes. The idea that 87Srr86Srpredominantly reflects variations in erosional fluxesfrom the continents thus appears incorrect. The pro-cesses operating during the Vendian and Cambrianperiods resulted in the largest change observed in87Srr86Sr of seawater at any time during the Earth

Ž .history Kaufman et al., 1993 . The resultingJ Sr curve with a peak at ;585 Ma suggest-CONT – SW

ing that part, but not all, of the change is the result ofa change in 87Srr86Sr of the continental flux. Incontrast, the Cenozoic change of J Sr whichCONT – SW

yields a peak at ;5 Ma, appears to be associatedwith very little change in the Sr isotopic compositionof the input to the oceans.

The balance equations for both C and Sr isotopesprovide the basis for obtaining changes in erosionŽ Sr . Ž C .J and organic C burial rates J asCONT – SW SW – ORG

a function of time based on the Sr and C isotopicrecords of seawater. The calculated rate of organic Cburial depend on the product of the erosion rate andŽ SW DM .d yd values in carbonates. Thus, it is possi-C C

13 Ž .ble to have relatively low d C values q1 to q2‰in marine carbonates during periods of high organiccarbon burial if the erosion rate is very high. Itfollows that the d SW is not a proxy for organicC

carbon burial rates while J C estimated fromSW – ORG

these equations is more likely to reflect real globalvariations in organic carbon burial rates. A curve forthe burial rate of organic C constrained by the C andSr isotopic variations shows a prominent peak at;575 Ma. In the latest Proterozoic, these higherosion rates, likely coupled with high organic pro-ductivity, contributed to a significant increase in theburial rate of organic C. The importance of thisresult is that the evolution of atmospheric O is, in2

part, controlled by J C . The long-term accumu-SW – ORG

lation of O in the atmosphere is primarily due to2

inputs related to the reduction of C, Fe or S in theexogenic cycle, and to burial of these in sediments.

For each mole of C buried in sediments one mole ofO is released to the atmosphere. The Vendian peak2

in organic C burial coupled with likely lower fluxesof reducing hydrothermal fluids gave rise to a largeincrease in O in the atmosphere after the younger2

Varanger glaciation. The peak in organic C burial atthis time is sufficient to generate most of the O in2

Ž C .the present atmosphere. Several peaks in JSW – ORG

inferred for the Phanerozoic need not cause largechanges in atmospheric O levels as they may be2

balanced by the S cycle.

9. Implications for Neoproterozoic ice ages

As shown in Figs. 4 and 5, there are at least fourNeoproterozoic glacial intervals that appear to becorrelative between sedimentary successions world-wide, all with characteristic positive-to-negative C

Ž .isotope excursions Kaufman et al., 1997 . In anŽ .earlier publication Kaufman et al., 1993 , we com-

pared these records with those of the Cenozoic toinvestigate if there is a general relationship betweenglobal tectonics and glaciations. We record a spec-

87 86 Žtacularly large change in Srr Sr larger than noted.for the Cenozoic of marine carbonates from low

Žpre-Vendian values ca. 0.7068; but Fairchild et al.Ž . .in press suggest this may be as low as 0.7063 to

Ž .high Middle Cambrian values 0.7090 . Changes inboth the Vendian and Cenozoic were attributed tohigh erosion rates associated with continent–conti-

Žnent collisions Pan-African and Himalayan–Tibe-. 87 86tan . However, for the Cenozoic, the large Srr Sr

rise predates the most intense continental glaciationswhile in the Neoproterozoic the largest rise appearsto post-date the last Varanger ice age. This suggestedthat uplift-driven models proposed to explain Ceno-zoic climatic change cannot account for Neoprotero-zoic ice ages. This may not be surprising since theNeoproterozoic glaciations have been suggested to

Žbe global in extent Harland, 1964; Kirschvink, 1992;.Hoffman et al., 1998a,b , now referred to as ‘Snow-

Ž .ball Earth Glaciations’. Harland 1964 recognized aworldwide distribution of tillites and associatedlithologies beneath strata bearing Ediacaran fossils,and considered that these rocks were deposited dur-ing an ice age that was much more severe than thePleistocene ice age. While a variety of models have


been discussed to account for the C isotope excur-Žsions associated with these glaciations cf. Knoll et

al., 1986; Kennedy, 1996; Kaufman et al., 1997;.Hoffman et al., 1998a,b , it is clear that knowing the

duration of these events would contribute substan-tially to our understanding of both the causes of theglaciations as well as their geochemical conse-quences.

The model discussed in the Section 8 can be usedto evaluate the timing of the isotopic excursionsassociated with glacial intervals given certain clearassumptions. The solutions of the mass balance equa-tions for C and Sr isotopes are:

SW DM C SWd t s d yy D q d tŽ . Ž .Ž .C C SW – ORG ORG – SW C 0

DM Cy d yy DŽ .C SW – ORG ORG – SW

=ty t0

exp y 25Ž .SWž /tC

a SW t s a DM qxSr a CONT ya DMŽ . Ž .Ž .Sr Sr CONT Sr Sr

SW DM Srq a t y a qxŽ . ŽSr 0 Sr CONT

=ty t0CONT DMa ya exp yŽ . .Sr Sr SWž /t Sr

26Ž .

for t) t and xSr , yC , a CONT, a DM , d DM0 CONT SW – ORG Sr Sr C

and D all constant.ORG – SW

The key factors in evaluating the response of sucha system to climatic forcing are estimates of theresidence times of C and Sr in seawater. The present

SW SW Žvalues are t s0.1 Ma and t s2.7 Ma cf.C Sr

Palmer and Edmond, 1989; Kump, 1991; Richter and.Turekian, 1993 , respectively. The Sr residence in

seawater can by mass balance be expressed as aŽ 24function of the mass of seawater M s1.41=10SW

˙. Žg , the rate of carbonate deposition M ;CARB˙ y1Ž . .M rM s0.00142 Ma , the hydrothermalCARB SW

˙ ˙ y1Ž . .water flux M ; M rM s0.1 Ma and theHW HW SW

effective enrichment of Sr in carbonate sedimentsŽ .relative to seawater D :Sr

y1˙ ˙M MHW CARBSWt s qD . 27Ž .Sr Srž / ž /M MSW SW

This yields an effective D value of 190 for theSr

present ocean. This compares with D values ofSr

;1000 and ;300–1000, respectively, for abioticand biotic aragonite, and D values of ;140 andSr

Ž;300, respectively, for abiotic and biotic calcite cf.Veizer, 1983; Carpenter et al., 1991; Carpenter and

.Lohmann, 1992 . Thus, the limits of possible SrŽresidence times in the oceans are 0.7 Ma for DSr

. Ž .s;1000 and 3.3 Ma for D s;140 . There isSr˙no indication that M was different in the Neo-CARB

Žproterozoic compared to the present cf. Derry et al.,˙.1989 and M was most likely higher compared toH W

today, suggesting a somewhat shorter residence timeŽ .for Sr in Neoproterozoic oceans see Fig. 8 . In

previous work, we used a Sr residence time of ;4Ž .Ma cf. Goldstein and Jacobsen, 1987 , which was a

partial residence time with respect to the river waterinput to the oceans. The total residence time isnecessarily shorter. In the following calculations weused 4 Ma for the Sr residence time since thisappears to be an upper limit.

The most severe effect one can imagine from aglobal glaciation is that all erosional transport to the

Ž Sr .ocean is effectively halted x s0 and that pri-CONT

Fig. 8. Relationship between the residence time of Sr in the˙Ž . Ž .oceans t , the rate of hydrothermal circulation M rMSr HW SW

and the effective global enrichment factor of Sr in carbonateŽsediments relative to seawater D ; the effective modern value isSr

.;190 . Curves are shown for the present rate of hydrothermalŽ y1 .cycling ;0.1 Ma ; Palmer and Edmond, 1989 and for refer-

Ž y1 .ence a value three times higher than the present flux 0.33 Ma .Also shown is the effect of partitioning of Sr between seawaterand biotic vs. abiotic carbonates; their values or ranges are shownfor both aragonites and calcites. If abiotic aragonites were thedominant precipitate in Neoproterozoic oceans, the residence timeof Sr is estimated from this figure to be ;0.7 Ma.


mary productivity in the ocean goes to zero, resultingŽ Cin no organic carbon burial in sediments ySW – ORG

. Srs0 . For a global glaciation with x s0 andCONT

yC s0 starting at t and ending at t , we haveSW – ORG 0 1

ty t0SW DM SW DMd t sd q d t yd exp yŽ . Ž .C C C 0 C SWž /tC

28Ž .

ty t0SW DM SW DMa t sa q a t ya exp yŽ . Ž .Sr Sr Sr 0 Sr SWž /t Sr

29Ž .DM SWŽ . DMwhere d sy5.5 and d t s8 and a sC C 0 Sr

SWŽ . Sr C0.7025 and a t s0.707. If x and ySr 0 CONT SW – ORG

return to normal values subsequent to the glaciationthen for t) t the C and Sr isotopic evolution will1

be:

d SW t sd SW tŽ . Ž .C C 0

ty t1SW SWq d t yd t exp yŽ . Ž .C 1 C 0 SWtC

30Ž .

a SW t sa SW tŽ . Ž .Sr Sr 0

ty t1SW SWq a t ya t exp y .Ž . Ž .Sr 1 Sr 0 SWt Sr

31Ž .

The response of 87Srr86Sr and d13C to such a

global glacial event is seen in Fig. 9. In this simplecalculation, full snowball conditions are assumed atts0, such that the high rates of carbon burial,interpreted from the extreme d

13C values, and river-ine inputs are shut off instantaneously. Given therelatively short residence time of carbon, d

13C ofseawater would approach depleted mantle valueswithin 300,000 years. This represents an absoluteminimum age constraint for the time between thehighly positive to highly negative d

13C values, as-suming only variations in burial of organic matter. Ifthe organic carbon burial rate decreased gradually,the length of time to reach mantle values wouldnecessarily increase. In the Otavi Group of Namibia,

Ž .Hoffman et al. 1998a,b show that this positive-to-negative excursion occurs over ca. 20 m in shallowwater carbonates deposited immediately beneath

Fig. 9. Response of 87Srr86 Sr and d13C to a glacial event that is

three times the C residence time in seawater. This calculationassumes that the maximal organic C burial flux, interpreted from

13 Ž .d C value ca. q8 of pre-glacial carbonate, stops and startsinstantaneously at ts0 and ts3=105 years, respectively. Overthis interval 87Srr86 Sr of seawater would experience little changedue the much longer residence time of this element. The model isused to calculate whole ocean changes. d

13C in carbonates primar-ily record the surface ocean reservoir and can potentially respond

Ž .much faster to perturbations than shown in a above.

glacial diamictite. According to our calculations, itwould take another 50,000 years for the shallow

13 Žocean to return to d C values near 0‰ notably, thisoccurs through 500 m of platform carbonates in the

.Otavi Group of Namibia; Hoffman et al., 1998a,b ,assuming the organic carbon burial flux were sud-

5 Ž .denly re-established at ts3=10 years Fig. 9 . Ofcourse, the negative d

13C of the ocean could con-tinue indefinitely, until the end of global glaciationand renewal of carbon burial fluxes. Any similarnegative carbon isotope event that occurred in less

Ž .than 350,000 years cf. Bowring et al., 1998a would


necessarily require the addition of a highly 13C-de-Žpleted source of carbon to seawater e.g., methane

.hydrates; Dickens et al., 1997 .Over the same 300,000 year interval, the expected

change in Sr isotopes, due to the sudden shutoff ofthe continental flux to seawater, is shown to be

Ž .insignificant Fig. 9 . This is consistent with theobserved lack of Sr isotope variations across most of

Ž .the Neoproterozoic ice ages Fig. 4 . However, overthe multi-million year time scale suggested by Hoff-

Ž .man et al. 1998a,b , the Sr isotopic composition ofthe ocean would necessarily approach depleted man-

Ž .tle values Fig. 10 . The observed record is sparse,but generally shows either no change or a slight risein 87Srr86Sr across the Sturtian and older Varangerice ages. Because of the scatter in the Sr isotopecurve we cannot preclude that 87Srr86Sr may de-crease by as much as ;0.001 during one of the

Žglacial events but the decrease is probably much.less and in some cases it appears to increase . As

Fig. 10. Response of 87Srr86 Sr to the length of the glacialinterval. For the proposed ca. 10 Ma length of the glacial intervalŽ . 87 86Hoffman et al., 1998a,b , the Srr Sr value would drop from0.707 to below 0.703 assuming total cessation of continental

Ž .inputs. Including a diagenetic flux of Sr similar to today 1=

then 87Srr86 Sr would drop to about 0.7038 after 10 Ma. CurvesŽ .are also shown for diagenetic fluxes that are three 3= and ten

Ž . 87 8610= times the present values, resulting in Srr Sr of 0.7047and 0.706 after 10 Ma, respectively. The shaded region indicates

Žthe amount of Sr isotopic change ca. 0.001 if only the hydrother-.mal flux is important expected after a 1 Ma interval. Unless it can

be shown that diagenetic fluxes were an order of magnitudegreater during snowball glaciations, we conclude that the Neopro-teorozoic glaciations lasted less than about 1 Ma each and that anymodel to explain these glaciations needs to account for thisconstraint.

shown in Fig. 10, this yields an upper limit of 1 Mato the duration of the glacial events. If instead theglacial events lasted as long as 10 Ma, as suggested

Ž . 87 86by Hoffman et al. 1998a,b , the Srr Sr value ofseawater should drop to a value close to the depletedmantle value of less than 0.703.

Ž .Hoffman et al. 1998b suggested as resolution tothis problem that the diagenetic flux was high due togreater acidity of the oceans during snowball Earthconditions. This acidity would come from the buildupof abundant hydrothermal CO in the deep ocean.2

For example, were the diagenetic flux to be in-Žcreased by a factor of 3 over the present flux cf.

. 87 86Richter and Turekian, 1993 seawater Srr SrŽ .would approach 0.7047 after 10 Ma see Fig. 10 .

For carbonate dissolution to buffer the seawater87Srr86Sr value, it would be necessary to increasethe diagenetic flux by at least an order of magnitude

Žcompared to the present value see 10= curve in.Fig. 10 . However, unless it can be shown that the

ŽNeoproterozoic was more acidic contrary to the.view from Pleistocene oceans; Sanyal et al., 1997 ,

the absence of a strong drop in 87Srr86Sr acrossthese ice ages must be explained by the fact that they

Žwere either very short lived minimally 300,000.years , or the post-glacial oceans were otherwise

buffered. Thus, unless it can be shown that suchlarge diagenetic fluxes operated during snowballglaciations, we conclude that the Neoproteorozoicglaciations only lasted about 0.3–1 Ma each and thatany model to explain these glaciations needs toaccount for this constraint.

10. Conclusions

Sr and C isotopic data obtained on samples ofmarine carbonates provide a relatively detailed recordof isotopic variation in seawater through the Neopro-terozoic, and allow direct correlation of these iso-topic changes for this time period. This data set wasused to revise the 87Srr86Sr and d

13C curves ofNeoproterozoic seawater.

The frequency and magnitude of variation in theSr isotope record appear comparable in the Nepro-terozoic to typical Phanerozoic values, however, therecord also includes the largest rise in 87Srr86Sr as


well as the lowest value observed for the past 800Ma. The d

13C curve shows both much higher valuesthan those observed in the Phanerozoic as well asrapid up to 15‰ changes associated with glacialevents.

The relationship between d13C and 87Srr86Sr

variations associated with Neoproterozoic glacial in-tervals strongly suggest that these lasted at least300,000 years but less than about 1 Ma if they areassociated with global ‘‘Snowball’’ glaciations.

Acknowledgements

This research was funded by NSF grant EAR94-18445 to SBJ and NSF grants 96-30928 and97-14070 to AJK. We thank M. Brasier and I.Fairchild for their helpful reviews.

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