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Oxygen isotopes in iron (III) oxides
Chemical Geology, 85 (1990) 337-344 Elsevier Science Publishers B.V., Amsterdam - - Printed in The Netherlands
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Oxygen isotopes in iron (III) oxides 2. Possible constraints on the depositional environment of a Precambrian quartz-hematite banded iron formation Crayton J. Yapp Department of Geology, University of New Mexico, Albuquerque, NM 87131 (U.S.A.) (Received May 25, 1989; revised and accepted March 8, 1990)
ABSTRACT Yapp, C.J., 1990. Oxygen isotopes in iron (III) oxides, 2. Possible constraints on the depositional environment of a Precambrian quartz-hematite banded iron formation. Chem. Geol., 85: 337-344. Quartz (SiO2) and hematite (Fe203) possess considerably different gram formula weights and different stoichiometric coefficients for their oxygen. These differences mean that the numerical values of SiO2 and Fe203 compositions when reported as weight percent are not the same as the numerical values of X(O), i.e. oxygen in the mineral as a mole percent of the total oxygen in a binary quartz-hematite system. Consequently, end-member quartz and hematite ~lSO-values which are determined from linear extrapolations of~180 vs. weight percent Fe203 data from binary mixtures can be in error by as much as several per mil. Extrapolation to end-member ~180-values in binary quartz-hematite systems should always be performed using mole fraction of oxygen as the compositional variable. The hematite-water fractionation curve of Yapp and the quartz-water curve of Knauth and Epstein were used together with the kinetic isotope exchange model of Criss et al. to constrain the depositional environment of the Late Proterozoic quartz-hematite banded iron formation deposits of Urucum, Brazil. The depositional temperatures permitted by the model assumptions employed here range from 0 ° to ~ 35 °C (or 54°C). The permitted range of depositional water ~180-values is from - 6.60/00 (or - 8.9%0) to 0.0%0. While the range of permitted depositional conditions is rather large, it does encompass the possible cool temperatures and brackish or fresh waters which are indicated by sedimentologic evidence in these deposits.
1. Introduction
Gregory and Criss (1986) and Criss et al. ( 1987 ) have emphasized the importance of~plots to the interpretation of oxygen isotope data from rocks. These authors present mathematical models for the representation of disequilibrium isotopic exchange in ~-~ coordinates. Isotopic exchange rates are modeled with "pseudo" first-order differential equations in which the rate of isotopic change of a particular mineral is proportional to the instantaneous departure of that mineral from isotopic equilibrium with the ambient water. These workers applied their equations to closed water-absent systems, closed water-rock sys0009-2541/90/$03.50
terns (their "closed" systems) and systems open to pervasive fluid flow. In the latter case the fluid flow can be represented as a continuum from no fluid movement to fluid-dominated (or "fluid-buffered" ) systems. Gregory and Criss (1986) and Gregory et al. (1989) discussed the quartz-magnetite ~ ~80 data from a number of Precambrian banded iron formations (BIF) in the context of these disequilibrium exchange models. They concluded that the pattern represented by a large ~180 range for magnetite and a small ~ 8 0 range for coexisting quartz would arise from disequilibrium isotopic exchange in an open system with the additional condition that the rate constant for magnetite-water oxygen iso-
© 1990 Elsevier Science Publishers B.V.
338 tope exchange must be much larger than the rate constant for quartz-water exchange. This characteristic steeply sloped open-system data array for ~18OMt vs. ~18OQtz was exhibited by several BIF. In particular it was evident in those deposits which have apparently experienced only low-grade metamorphism (Gregory et al., 1989). The equations of Criss et al. (1987) have not yet been applied to oxygen isotope data from quartz-hematite-dominated Precambrian BIF. The application of these kinetic equations to such data requires knowledge of the equilibrium oxygen isotope fractionation factor for hematite-water. The low-temperature hematite (goethite)-water fractionation curve presented in Part 1 (Yapp, 1990a in this issue) provides an impetus for examining quartz-hematite oxygen isotope data from hematite-rich Precambrian BIF of low metamorphic grade in the context of the kinetic exchange theory of Gregory and Criss (1986) and Criss et al. (1987 ). Published data from appropriate BIF appear to be limited to the quartz-hematite deposits of Urucum, Brazil (Hoefs et al., 1987). The magnetite-water fractionation curve of Blattner et al. (1983) was employed by Hoefs et al. ( 1987 ) to interpret the Urucum hematite ~180 data. The present work examines the constraints that can be placed on the environment of deposition of the Urucum deposits using the hematite-water curve of Yapp (1990a) in the context of the aforementioned kinetic exchange theory. 2. Published Urucum data
Hoefs et al. (1987) analyzed eight samples of banded hematite-jaspilite from the Urucum deposits. They noted that the samples were composed of only two minerals, quartz and hematite. Of these eight samples, one (their U45) contained grains of detrital quartz which eliminates it as a probable source of authigenic mineral-pair isotope paleotemperature data. These authors pointed out that their samples
C.J.YAPP were of such a fine-grained, intimately mixed nature that it was impossible to obtain mineral separates for oxygen isotope analyses. Consequently, they cut their samples into sections which consisted of either Fe-rich or Fe-poor layers. These layers were analyzed for their Fe203 contents and their ~ ~80-values. Hoefs et al. ( 1987 ) plotted their measured 6~80-values for the different Fe-rich and Fe-poor layers from a sample against the corresponding "percent of Fe203" in these layers. The good linear correlation thus obtained was extrapolated to pure SiO2 and Fe203 in each sample to obtain the apparent O~80-values of those end-members in each sample. Six of their samples were treated in this manner (U16, U19, U27, U32, U38 and U44). Among other things this method requires that the respective ~ tSO-values of end-member quartz and hematite be uniform throughout the volume of space defined by the sample. Linearity of a ~180 vs. composition plot in a two-component system supports the assumption of spatially uniform end-member O180-values within a sample. However, it is not stated in the paper by Hoefs et al. (1987) whether their reported "percent Fe203" represents weight percent Fe203, mole percent Fe203 or oxygen in Fe203 as a mole percent of the total oxygen. The differences between these compositional variables are of importance in the calculation of end-member ~80-values from the 6180 values of quartzhematite mixtures. The proper choice of the composition variable for such end-member calculations is oxygen in the component of interest as a molefraction of the total oxygen in the analyzed aliquot. Thus, for t80<< 160 in a two-component system: ~ t 80mixture = ~ 1802"~- ((~1801--(~1802)X(0)1
(1) where ~ ~80 ~, 61802 = ~ 180 of pure component 1 and 2, respectively;
X(O),=n(O),/[n(O)~+n(O)2]
OXYGENISOTOPESIN IRON (Ill) OXIDES,2 i
~
i
I
]
I
i
339 i
]
i
1
1.0 /-/ ///
0.8
///////~ ///////
E 0.6
"T
8 x
0.4
0.2
0.0 I 0.0
i
~ 0.2
E
i 0.4
i
016
I
.I 0 8
i
I 1.0
WHm Fig. 1. The relationship between the mole fraction (X) of total-system oxygen contained in hematite and the weight fraction (W) of hematite in a two-component, quartzhematite system. The correct relationship is indicated by the solid curve. The dashed line is a reference line depicting a = b. Only at the pure end-member compositions does X= W. The correct relationship between X and W represented by the solid curve has implications for the calculation of end-member ci~80-values from measurement of c~' 80-values in mixtures of the end-members (see text ).
and n (O) h n (O) 2 = moles of oxygen in components 1 and 2, respectively. The weight fraction of hematite ( WHm) in a
quartz-hematite system is not equal to X(O ) Hm in eqn. 1. A plot of X(O ) Hm VS. WHm for a two-component quartz-hematite system is found in Fig. 1. As can be seen in Fig. 1, over a limited range of Wnm-values the relationship between X(O ) Hmand Wiamwill approximate a straight line. However, either the intercept of this approximate straight line will be different from zero, or the slope will be different from one, or both. Thus, a plot of c~80 (mixture) vs. WXm for a limited range of Wnm in a twocomponent system will be approximately linear, but it will not yield the correct ~ ~SO-values upon extrapolation to end-member compositions. Furthermore, the end-member c~80values obtained from a linear extrapolation in a plot of ~18Ornixture VS. WHm will depend upon sample selection. As a general illustration of the possible errors that can arise from the use of weight fraction data in the calculation of end-member c~80-values in quartz-hematite mixtures, the "percent Fe:O3" data of Hoefs et al. (1987) were assumed to be weight percent Fe203 and were converted to corresponding X(O)Hm data. The measured ~18Omixture data of Hoefs et al. ( 1987 ) were regressed against the corresponding X(O ) Hm-Values for each of the six samples discussed previously and end-member c~180-
TABLE 1 Comparison of end-member quartz and hematite 6~80-values (in %o vs. V-SMOW) for samples from the Urucum deposits, Brazil, as calculated by use of the presumed weight fraction of the chemical component (case A) and by use of the corresponding mole fraction of total oxygen contained in the component (case B) Sample
Case A hematite
Case B quartz
apparent T
hematite
quartz
(°c) UI6 UI9 U27 U32 U38 U44
+5.4 +3.9 +0.9 + 6.4 + 3.4 +6.1
+25.5 +22.6 +22.6 + 24.8 + 21.4 +26.3
95 120 69 127 135 93
apparent T
(°c) +3.3 +0.7 -2.4 + 4.2 + 0.6 +2.3
+24.6 +21.8 +21.6 + 22.8 + 20.0 +25.3
76 78 41 122 105 54
Raw data used in the calculations are from Hoefs et al. (1987). Apparent temperatures were calculated using eq. 2 in the text.
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C.J. YAPP
calculated using the two sets of compositional data can differ by as much as 3.8%0, while quartz ~180-values differ by as much as 2.00/oo. The preceding discussion and calculation are intended as an illustration of the importance of performing these types of end-member calculations using only X(O )Hm or X(O )otz data. Since it is not obvious what the "percent Fe203" data of Hoefs et al. (1987) actually
values for quartz and hematite were calculated. For each of these six regressions the linear correlation coefficient (r) was better than - 0 . 9 9 . The end-member quartz and hematite 180-values calculated using eqn. 1 are listed in Table I together with end-member ~ 180_val_ ues calculated using the published "percent Fe203" data of Hoefs et al. (1987). It is seen that the end-member hematite ~180-values I
I
I
I
I
I
Open-System Exchange ~ l a O w =0.0 %0
/
Exchange T = 135"C
10
/
Y
Urucum BIF Data From Hoefs
/
et al. (1987)
6 E "1-
0
O0
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/ I
16
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20
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~1
24
KHm > > KQtz i
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28
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32
36
~18Oet z
Fig. 2. A plot of corresponding hematite and quartz 61SO-values from the Late Proterozolc BIF of Urucum (solid triangles). The data are calculated from those published by Hoefs et al. ( 1987 ) and are listed in case A of Table I in the current work. Trajectory A in the diagram represents the water-absent, closed-system isotope exchange trajectory for a system with an XQtz/XHm ratio of ~ 0.53 (see text). Trajectory A has been constrained to pass through the point representing sample U44 to provide a limit on the domain of open-system prograde exchange. The dashed line labeled "Equilibrium Marine Precipitates" defines the locus of points representing the &lSO-values of hematite and quartz precipitated at different temperatures from ocean water with a &tSO-value of 0.0%0. Trajectory A and the marine precipitate line intersect at the point represented by the open square. The dotted line is the extrapolation of trajectory A to the 0 °C equilibrium isotherm. Trajectory B is the "fluid-buffered" open-system exchange trajectory for the case of g r i m > > gQtz, an exchange temperature of 135 °C and water with a di~sO of 0.0%0 (see text ). Trajectory B originates at the open square and represents the limiting open-system exchange trajectory under the model assumptions. All of the data points in the diagram could have evolved from the presumed initial point (open square) within the limits of fluid flow (/z) and relative rate constant magnitudes represented by trajectories A and B. All of the data points could also have evolved (given the assumptions of the model ) from any initial point lying on, or within, the boundaries defining the stippled area (for a minimum prograde exchange T of 135 ° C). This stippled area represents the range of initial temperatures and water ~ l SO.values permitted by the exchange model assumptions (see text). The permitted depositional temperatures range from 0 ° to ~ 35 ° C, while the ~lSO-values of the possible depositional water range from - 6 . 6 to 0.0%o.
OXYGENISOTOPESIN IRON(III) OXIDES,2
represent, I will adopt the working assumption that their "percent Fe203" values are the appropriate units and yield valid end-member quartz and hematite ~ 180_values" These values (case A, Table I) are employed in subsequent discussions, but the model results for both cases A and B are presented. It should be mentioned that uncertainties in the calculated end-member ~ 180_values can be as high as _+1%0 due to scatter in the measured data, even if the appropriate compositional variable has been used. Fig. 2 contains a plot of end-member 318OHm VS. ~18OQtz (case A, Table I) for the six samples from the Urucum deposits analyzed by Hoefs et al. (1987 ). An interesting feature of the data array in Fig. 2 is the relatively large range in quartz ~180-values (+21.4 to +26.3%0) as well as hematite ~180-values ( + 0.9 to + 6.4%o). These relatively large d180 ranges for both quartz and hematite contrast with the small ~ 180 ranges for quartz and large 3180 ranges for magnetite noted by Gregory et al. ( 1989 ) for many quartz-magnetite-rich BIF of low metamorphic grade.
3. Model of isotopic exchange The presence of quartz-filled fractures of possible diagenetic origin in the Urucum deposits (Hoefs et al., 1987 ) suggests that water was moving through the deposits and that isotopic exchange could be modeled as an open system. The Urucum deposits may have been buried to depths of only ~ 2 km (Hoefs et al., 1987 ), which indicates relatively low temperatures of exchange. The work of Gregory and Criss (1986) and Gregory et al. (1989) emphasized the resistance of quartz to low-temperature retrograde oxygen isotope exchange. Furthermore, the results of Becker and Clayton (1976) and Yapp (1990b) indicate that hematite oxygen isotope exchange is not likely to occur in low-temperature retrograde processes. This suggests that the quartz and hematite of the Urucum deposits may have ac-
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quired their 180/160 values during a prograde process such as burial diagenesis. It is assumed that the rate constant for prograde hematite-water oxygen isotope exchange (KRm) is much greater than that for prograde quartz-water exchange (KQtz). There is at present no experimental evidence to support this assumption, but the assumption is made to facilitate discussion and is in keeping with the apparent fact that KMt is much greater than KQtz in quartz-magnetite BIF of low metamorphic grade (Gregory and Criss, 1986; Gregory et al., 1989 ). Oxygen isotope disequilibrium in the Urucum deposits is suggested by the observation that samples from the highest stratigraphic levels (shallowest depth of burial) have the highest apparent temperature of formation (see U38 in Table I and also Hoefs et al., 1987). The extreme conditions of a "fluid-buffered" open system ( ~ ) and a water-absent closed system (Criss et al., 1987 ) will be employed in an attempt to set limits on the initial depositional environment of the Urucum BIF deposits. It is assumed that the observed quartz and hematite ~ 180_values were acquired during diagenetic disequilibrium isotope exchange. In addition it is assumed that the waters of the depositional environment were the input waters for diagenetic exchange, that all authigenic quartz and hematite of the deposit had sedimentary precursors which formed under identical equilibrium depositional conditions and that a probable maximum ~180-value for the presumed depositional water was 0.0%o. i.e. seawater (e.g., Epstein and Mayeda, 1953). The mathematical details of the model employed here can be found in Criss et al. ( 1987 ). If the quartz-water fractionation curve of Knauth and Epstein (1976) is combined with the hematite-water curve of Yapp (1990a), the following equation results: 1000 In OI~Q_H~---1.46" 10 6 T - 2 + 9.0
(2)
w h e r e OgQ_H-~ ( 180/160)Qtz / ( 180/160)Hm; and
T is in kelvins. The highest apparent tempera-
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ture calculated for the samples of case A in Table I is given by U38 (135°C) using eq. 2. It will be assumed that this represents an estimate of the minimum temperature of isotopic exchange during diagenesis. Higher presumed exchange temperatures will not affect the conclusions of this study. Use of an alternative quartz-water fractionation curve, such as that of Clayton et al. (1972), in conjunction with the hematite-water curve of Yapp (1990a) would result in the calculation of systematically higher apparent temperatures for the samples of Table I. Although there is likely to be considerable compositional heterogeneity in the Urucum BIF deposits on the scale of hand samples, the lack of published average modal abundance data on the six samples of Table I requires that the limiting water-absent, closed-system exchange trajectory be approximated by the average compositional data reported for the composite BIF deposits. The average Fe content appears to be ~ 5 4 wt.% (Hoefs et al., 1987 ). This value for a system composed only of quartz and hematite corresponds to an XQtz/ nHm ratio of ~0.53, where X represents the oxygen in the indicated mineral as a mole fraction of the total system oxygen. The negative value of the ratio OfXQtz/SHm in closed-system prograde exchange corresponds to the slope of the exchange trajectory in ~Hm vs. ~Qtz coordinates [see Criss et al. (1987) for mathematical details]. The significance of this exchange trajectory in this discussion is that it represents an upper boundary condition on the proposed field of open-system prograde exchange. In particular the closed-system exchange trajectory represent the limiting conditions of p = 0 and X w = 0 in the open-system model, where/z is the normalized flow rate of incoming water and Xw is the oxygen in water as a mole fraction of the total oxygen in the system at any instant (see Criss et al., 1987). Sample U44 of case A in Table I corresponds to a limiting isotopic coordinate for water-absent closed-system prograde exchange in this dis-
C.J. YAPP
cussion. The equation of the chosen closedsystem prograde exchange trajectory through U44 is: 18OHm= -- 0.53~ 18OQtz+ 20.0
(3)
This equation is depicted as trajectory A in Fig. 2. The assumption that this is prograde diagenetic exchange allows the placement of constraints on the initial point of origin of trajectory A. The lower limit on the depositional temperature of the chemical sediment precursors to the quartz and hematite of U44 is presumed to be 0 ° C. Thus, extrapolation (dotted line) of trajectory A to the 0°C isotherm of Fig. 2 places a lower limit on the ~L80-value of the presumed depositional waters. This calculated 180-value for the initial water is -6.6%o. The dashed line in Fig. 2 labeled "Equilibrium Marine Precipitates" defines the locus of isotopic compositions of quartz and hematite in equilibrium at different temperatures with waters whose ~80-value is 0.0%o (i.e. presumably ocean water). It is assumed in the prograde exchange model presented here that the initial depositional waters did not have ~'80-values more positive than 0.0%o. Consequently, the intersection of closed-system trajectory A with the curve defined by equilibrium marine precipitates permits the calculation of a possible maximum temperature of deposition in the original chemical sediment. This intersection is indicated by the open square in Fig. 2 and corresponds to a depositional temperature of ~ 35 ° C. The other extreme in the continuum of opensystem isotopic exchange is represented by the limiting condition o f p ~ o o (fluid-buffered exchange) with the aforementioned assumption that grim >> gQt z. Open-system fluid-buffered isotope exchange with fluid whose fi180-value is 0.0%o at an exchange temperature of 135 °C is shown in Fig. 2 as trajectory B. The initial point for trajectory B is the open box in Fig. 2 representing the intersection of closed-system trajectory A with the locus of equilibrium ma-
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OXYGEN ISOTOPES IN IRON (III) OXIDES, 2
rine precipitation. Note that all open-system exchange trajectories for which 0 > Kqtz will lie within the boundaries defined by trajectories A and B and the isotherm chosen as the temperature of exchange [see Criss et al. (1987) for mathematical details]. It is seen that all six of the data points for case A of Table I lie within this exchange domain or on its boundaries. This implies that the isotopic values of the initial point of trajectories A and B of Fig. 2 are representative of plausible original depositional conditions (i.e. ocean water at ~ 35°C). However, the point represented by the open square in Fig. 2 is not a unique initial value for exchange trajectories which define domains within which the data points of Fig. 2 will lie under the assumptions of this model. In fact, any initial point which lies on, or within, the boundaries of the stippled area in Fig. 2 will yield limiting closedand open-system exchange trajectories which encompass all of the data of Fig. 2 for a prograde exchange temperature not lower than 135°C. Thus, in the context of of this model the temperature of the original environment of deposition might have been anywhere between ~ 0 ° and ~ 35 ° C and the d t80 of the ambient water could have been between ~ - 6 . 6 and ~0.0%o. This proposed possible range of initial environmental conditions is only as credible as the assumptions which were employed to apply the exchange model. At this time there are no objective criteria which can be used to validate the assumptions. In addition there are the previously discussed uncertainties associated with the choices of appropriate calculated quartz and hematite d~80-values. For example, if the d~80-values of case B in Table I had been used in the exchange model, a maximum possible depositional temperature of 54 °C would have been indicated for precipitation o f the chemical sediments from ocean water with a d~80value of 0.0%o. The minimum possible calculated d~80-value of the environmental water would be ~ -8.90/00 at a depositional temper-
ature of 0 ° C using the data of case B of Table I. While the calculated ranges of depositional temperatures and water ~180-values for the Urucum quartz-hematite BIF are rather large, it is important to note that cooler temperatures and brackish or fresh waters in the environment of deposition are permitted by the exchange model. Walde et al. ( 1981 ) discuss the stratigraphy of the Jacadigo Group which contains the Urucum quartz-hematite deposits in the Band' Alta Formation. They note that the arkoses and conglomerates which underlie and are vertically transitional to the quartz-hematite deposits probably indicate a continental depositional environment. Furthermore, Walde et al. ( 1981 ) point out the existence of possible dropstones of glacial origin in the Band' Alta Formation. This sedimentologic evidence for the existence of contemporaneous glaciation and perhaps a nonmarine depositional environment is consistent with the cooler temperatures and freshwater components permitted by the stippled domain of initial values shown in Fig. 2. The exchange model applied to the Urucum data does not address the question of isotope exchange mechanisms. It assumes continuous isotopic exchange at a single temperature and does not consider the isotopic effects of mineralogic phase changes during diagenesis. Knowledge of the relative timing of such transitions for silica and ferric oxides during diagenesis is lacking, but may be of fundamental importance to the ultimate interpretation of 5t80 data arrays such as those of the Urucum deposits. 4. Conclusions
The use of oxygen mole fractions is of fundamental importance when calculating the 5 ~80-values of end-member components from the measured compositional and fi 180-values of binary quartz-hematite mixtures. Failure to use oxygen mole fractions can result in errors
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of up to several per mil in calculated endmember 6 ~SO-values. The application of the kinetic isotopic exchange model of Gregory and Criss ( 1986 ) and Criss et al. ( 1987 ) to the quartz-hematite 6180 data published by Hoefs et al. ( 1987 ) for the Late Proterozoic BIF deposits of Urucum, Brazil, provides broad limits on the original depositional environment of these chemical sediments. Use of the hematite-water fractionation curve of Yapp (1990a) and the quartz-water curve of Knauth and Epstein ( 1976 ) in the model calculations yields a permitted depositional temperature range between 0 ° and ~ 35 °C (or perhaps 54°C). The indicated range of initial water 6180-values is between ~ - 6 . 6 ( o r - 8 . 9 ) and 0.0%o. The cooler temperatures and brackish or fresh waters in the environment of deposition which are permitted by the calculated model conditions are consistent with sedimentologic evidence for contemporaneous glaciation and a possible nonmarine depositional environment for these deposits. Gregory et al. (1989) were not sanguine about the possibility of deducing information on depositional environments from oxygen isotope data of Precambrian quartz-magnetite BIF. While not as pessimistic as Gregory et al. (1989), the conclusion to be drawn from the current treatment of the Urucum data is that the depositional conditions of largely unmetamorphosed Precambrian quartz-hematite BIF can, at present, only be constrained within broad limits using only oxygen isotope data. Acknowledgements The ultimate treatment of, and conclusions concerning, the data from the Urucum deposits were considerably influenced by the thoughtful review of Robert Criss, although he may not necessarily agree with the conclusions
C.J. YAPP
themselves. This research was supported by NSF grant EAR-8719070.
References Becker, R.H. and Clayton, R.N., 1976. Oxygen isotope study of a Precambrian banded iron-formation, Hamersley Range, Western Australia. Geochim. Cosmochim. Acta, 40:1153-1165. Blattner, P., Braithwater, W.R. and Glover, R.B., 1983. New evidence on magnetite oxygen isotope geothermometers at 175 °and 112 °C in Wairakei steam piplelines (New Zealand). Isot. Geosci., 1: 195-204. Clayton, R.N., O'Neil, J.R. and Mayeda, T.K., 1972. Oxygen isotope exchange between quartz and water. J. Geophys. Res., 77: 3057-3067. Criss, R.E., Gregory, R.T. and Taylor, Jr., H.P., 1987. Kinetic theory of oxygen isotopic exchange between minerals and water, Geochim. Cosmochim. Acta, 51: 10991108. Epstein, S. and Mayeda, T.K., 1953. Variation of 180 content of waters from natural sources. Geochim. Cosmochim. Acta, 4: 213-224. Gregory, R.T. and Criss, R.E., 1986. Isotopic exchange in open and closed system. In: Stable Isotopes in High Temperature Geological Processes. Mineral. Soc. Am., Rev. Mineral., 16: 91-127. Gregory, R.T., Criss, R.E. and Taylor, Jr., H.P., 1989. Oxygen isotope exchange kinetics of mineral pairs in closed and open systems: applications to problems of hydrothermal alteration of igneous rocks and Precambrian iron formations. Chem. Geol., 75: 1-42. Hoefs, J., Muller, G., Schuster, K.A. and Walde, D., 1987. The Fe-Mn ore deposits of Urucum, Brazil: an oxygen isotope study. In: N. Clauer and S. Chaudhuri (Editors), Isotopes in the Sedimentary Cycle. Chem. Geol. ( Isot. Geosci. Sect. ), 65:311-319 (special issue ). Knauth, L.P. and Epstein, S., 1976. Hydrogen and oxygen isotope ratios in nodular and bedded cherts. Geochim. Cosmochim. Acta, 40:1095-1108. Walde, D.H.G., Gierth, E., Clausthal-Zellerfled, and Leonardos O.H. 1981 Stratigraphy and mineralogy of the manganese ores of Urucum, Mato Grosso, Brazil. Geol. Rundsch., 70: 1077-1085. Yapp, 1990a. Oxygen isotopes in iron (III) oxides, 1. Mineral-water fractionation factors. Chem. Geol., 85: 329-335 (this issue). Yapp, C.J., 1990b. Oxygen isotope effects associated with the solid state a-FeOOH to a-Fe203 phase transformation. Geochim. Cosmochim. Acta, 54: 229-236.
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Oxygen isotopes in iron (III) oxides 2. Possible constraints on the depositional environment of a Precambrian quartz-hematite banded iron formation Crayton J. Yapp Department of Geology, University of New Mexico, Albuquerque, NM 87131 (U.S.A.) (Received May 25, 1989; revised and accepted March 8, 1990)
ABSTRACT Yapp, C.J., 1990. Oxygen isotopes in iron (III) oxides, 2. Possible constraints on the depositional environment of a Precambrian quartz-hematite banded iron formation. Chem. Geol., 85: 337-344. Quartz (SiO2) and hematite (Fe203) possess considerably different gram formula weights and different stoichiometric coefficients for their oxygen. These differences mean that the numerical values of SiO2 and Fe203 compositions when reported as weight percent are not the same as the numerical values of X(O), i.e. oxygen in the mineral as a mole percent of the total oxygen in a binary quartz-hematite system. Consequently, end-member quartz and hematite ~lSO-values which are determined from linear extrapolations of~180 vs. weight percent Fe203 data from binary mixtures can be in error by as much as several per mil. Extrapolation to end-member ~180-values in binary quartz-hematite systems should always be performed using mole fraction of oxygen as the compositional variable. The hematite-water fractionation curve of Yapp and the quartz-water curve of Knauth and Epstein were used together with the kinetic isotope exchange model of Criss et al. to constrain the depositional environment of the Late Proterozoic quartz-hematite banded iron formation deposits of Urucum, Brazil. The depositional temperatures permitted by the model assumptions employed here range from 0 ° to ~ 35 °C (or 54°C). The permitted range of depositional water ~180-values is from - 6.60/00 (or - 8.9%0) to 0.0%0. While the range of permitted depositional conditions is rather large, it does encompass the possible cool temperatures and brackish or fresh waters which are indicated by sedimentologic evidence in these deposits.
1. Introduction
Gregory and Criss (1986) and Criss et al. ( 1987 ) have emphasized the importance of~plots to the interpretation of oxygen isotope data from rocks. These authors present mathematical models for the representation of disequilibrium isotopic exchange in ~-~ coordinates. Isotopic exchange rates are modeled with "pseudo" first-order differential equations in which the rate of isotopic change of a particular mineral is proportional to the instantaneous departure of that mineral from isotopic equilibrium with the ambient water. These workers applied their equations to closed water-absent systems, closed water-rock sys0009-2541/90/$03.50
terns (their "closed" systems) and systems open to pervasive fluid flow. In the latter case the fluid flow can be represented as a continuum from no fluid movement to fluid-dominated (or "fluid-buffered" ) systems. Gregory and Criss (1986) and Gregory et al. (1989) discussed the quartz-magnetite ~ ~80 data from a number of Precambrian banded iron formations (BIF) in the context of these disequilibrium exchange models. They concluded that the pattern represented by a large ~180 range for magnetite and a small ~ 8 0 range for coexisting quartz would arise from disequilibrium isotopic exchange in an open system with the additional condition that the rate constant for magnetite-water oxygen iso-
© 1990 Elsevier Science Publishers B.V.
338 tope exchange must be much larger than the rate constant for quartz-water exchange. This characteristic steeply sloped open-system data array for ~18OMt vs. ~18OQtz was exhibited by several BIF. In particular it was evident in those deposits which have apparently experienced only low-grade metamorphism (Gregory et al., 1989). The equations of Criss et al. (1987) have not yet been applied to oxygen isotope data from quartz-hematite-dominated Precambrian BIF. The application of these kinetic equations to such data requires knowledge of the equilibrium oxygen isotope fractionation factor for hematite-water. The low-temperature hematite (goethite)-water fractionation curve presented in Part 1 (Yapp, 1990a in this issue) provides an impetus for examining quartz-hematite oxygen isotope data from hematite-rich Precambrian BIF of low metamorphic grade in the context of the kinetic exchange theory of Gregory and Criss (1986) and Criss et al. (1987 ). Published data from appropriate BIF appear to be limited to the quartz-hematite deposits of Urucum, Brazil (Hoefs et al., 1987). The magnetite-water fractionation curve of Blattner et al. (1983) was employed by Hoefs et al. ( 1987 ) to interpret the Urucum hematite ~180 data. The present work examines the constraints that can be placed on the environment of deposition of the Urucum deposits using the hematite-water curve of Yapp (1990a) in the context of the aforementioned kinetic exchange theory. 2. Published Urucum data
Hoefs et al. (1987) analyzed eight samples of banded hematite-jaspilite from the Urucum deposits. They noted that the samples were composed of only two minerals, quartz and hematite. Of these eight samples, one (their U45) contained grains of detrital quartz which eliminates it as a probable source of authigenic mineral-pair isotope paleotemperature data. These authors pointed out that their samples
C.J.YAPP were of such a fine-grained, intimately mixed nature that it was impossible to obtain mineral separates for oxygen isotope analyses. Consequently, they cut their samples into sections which consisted of either Fe-rich or Fe-poor layers. These layers were analyzed for their Fe203 contents and their ~ ~80-values. Hoefs et al. ( 1987 ) plotted their measured 6~80-values for the different Fe-rich and Fe-poor layers from a sample against the corresponding "percent of Fe203" in these layers. The good linear correlation thus obtained was extrapolated to pure SiO2 and Fe203 in each sample to obtain the apparent O~80-values of those end-members in each sample. Six of their samples were treated in this manner (U16, U19, U27, U32, U38 and U44). Among other things this method requires that the respective ~ tSO-values of end-member quartz and hematite be uniform throughout the volume of space defined by the sample. Linearity of a ~180 vs. composition plot in a two-component system supports the assumption of spatially uniform end-member O180-values within a sample. However, it is not stated in the paper by Hoefs et al. (1987) whether their reported "percent Fe203" represents weight percent Fe203, mole percent Fe203 or oxygen in Fe203 as a mole percent of the total oxygen. The differences between these compositional variables are of importance in the calculation of end-member ~80-values from the 6180 values of quartzhematite mixtures. The proper choice of the composition variable for such end-member calculations is oxygen in the component of interest as a molefraction of the total oxygen in the analyzed aliquot. Thus, for t80<< 160 in a two-component system: ~ t 80mixture = ~ 1802"~- ((~1801--(~1802)X(0)1
(1) where ~ ~80 ~, 61802 = ~ 180 of pure component 1 and 2, respectively;
X(O),=n(O),/[n(O)~+n(O)2]
OXYGENISOTOPESIN IRON (Ill) OXIDES,2 i
~
i
I
]
I
i
339 i
]
i
1
1.0 /-/ ///
0.8
///////~ ///////
E 0.6
"T
8 x
0.4
0.2
0.0 I 0.0
i
~ 0.2
E
i 0.4
i
016
I
.I 0 8
i
I 1.0
WHm Fig. 1. The relationship between the mole fraction (X) of total-system oxygen contained in hematite and the weight fraction (W) of hematite in a two-component, quartzhematite system. The correct relationship is indicated by the solid curve. The dashed line is a reference line depicting a = b. Only at the pure end-member compositions does X= W. The correct relationship between X and W represented by the solid curve has implications for the calculation of end-member ci~80-values from measurement of c~' 80-values in mixtures of the end-members (see text ).
and n (O) h n (O) 2 = moles of oxygen in components 1 and 2, respectively. The weight fraction of hematite ( WHm) in a
quartz-hematite system is not equal to X(O ) Hm in eqn. 1. A plot of X(O ) Hm VS. WHm for a two-component quartz-hematite system is found in Fig. 1. As can be seen in Fig. 1, over a limited range of Wnm-values the relationship between X(O ) Hmand Wiamwill approximate a straight line. However, either the intercept of this approximate straight line will be different from zero, or the slope will be different from one, or both. Thus, a plot of c~80 (mixture) vs. WXm for a limited range of Wnm in a twocomponent system will be approximately linear, but it will not yield the correct ~ ~SO-values upon extrapolation to end-member compositions. Furthermore, the end-member c~80values obtained from a linear extrapolation in a plot of ~18Ornixture VS. WHm will depend upon sample selection. As a general illustration of the possible errors that can arise from the use of weight fraction data in the calculation of end-member c~80-values in quartz-hematite mixtures, the "percent Fe:O3" data of Hoefs et al. (1987) were assumed to be weight percent Fe203 and were converted to corresponding X(O)Hm data. The measured ~18Omixture data of Hoefs et al. ( 1987 ) were regressed against the corresponding X(O ) Hm-Values for each of the six samples discussed previously and end-member c~180-
TABLE 1 Comparison of end-member quartz and hematite 6~80-values (in %o vs. V-SMOW) for samples from the Urucum deposits, Brazil, as calculated by use of the presumed weight fraction of the chemical component (case A) and by use of the corresponding mole fraction of total oxygen contained in the component (case B) Sample
Case A hematite
Case B quartz
apparent T
hematite
quartz
(°c) UI6 UI9 U27 U32 U38 U44
+5.4 +3.9 +0.9 + 6.4 + 3.4 +6.1
+25.5 +22.6 +22.6 + 24.8 + 21.4 +26.3
95 120 69 127 135 93
apparent T
(°c) +3.3 +0.7 -2.4 + 4.2 + 0.6 +2.3
+24.6 +21.8 +21.6 + 22.8 + 20.0 +25.3
76 78 41 122 105 54
Raw data used in the calculations are from Hoefs et al. (1987). Apparent temperatures were calculated using eq. 2 in the text.
340
C.J. YAPP
calculated using the two sets of compositional data can differ by as much as 3.8%0, while quartz ~180-values differ by as much as 2.00/oo. The preceding discussion and calculation are intended as an illustration of the importance of performing these types of end-member calculations using only X(O )Hm or X(O )otz data. Since it is not obvious what the "percent Fe203" data of Hoefs et al. (1987) actually
values for quartz and hematite were calculated. For each of these six regressions the linear correlation coefficient (r) was better than - 0 . 9 9 . The end-member quartz and hematite 180-values calculated using eqn. 1 are listed in Table I together with end-member ~ 180_val_ ues calculated using the published "percent Fe203" data of Hoefs et al. (1987). It is seen that the end-member hematite ~180-values I
I
I
I
I
I
Open-System Exchange ~ l a O w =0.0 %0
/
Exchange T = 135"C
10
/
Y
Urucum BIF Data From Hoefs
/
et al. (1987)
6 E "1-
0
O0
(,c 2 /
-2 -
/ I
16
I
I
20
I
I
/
y
~1
24
KHm > > KQtz i
I
28
I
I
h
32
36
~18Oet z
Fig. 2. A plot of corresponding hematite and quartz 61SO-values from the Late Proterozolc BIF of Urucum (solid triangles). The data are calculated from those published by Hoefs et al. ( 1987 ) and are listed in case A of Table I in the current work. Trajectory A in the diagram represents the water-absent, closed-system isotope exchange trajectory for a system with an XQtz/XHm ratio of ~ 0.53 (see text). Trajectory A has been constrained to pass through the point representing sample U44 to provide a limit on the domain of open-system prograde exchange. The dashed line labeled "Equilibrium Marine Precipitates" defines the locus of points representing the &lSO-values of hematite and quartz precipitated at different temperatures from ocean water with a &tSO-value of 0.0%0. Trajectory A and the marine precipitate line intersect at the point represented by the open square. The dotted line is the extrapolation of trajectory A to the 0 °C equilibrium isotherm. Trajectory B is the "fluid-buffered" open-system exchange trajectory for the case of g r i m > > gQtz, an exchange temperature of 135 °C and water with a di~sO of 0.0%0 (see text ). Trajectory B originates at the open square and represents the limiting open-system exchange trajectory under the model assumptions. All of the data points in the diagram could have evolved from the presumed initial point (open square) within the limits of fluid flow (/z) and relative rate constant magnitudes represented by trajectories A and B. All of the data points could also have evolved (given the assumptions of the model ) from any initial point lying on, or within, the boundaries defining the stippled area (for a minimum prograde exchange T of 135 ° C). This stippled area represents the range of initial temperatures and water ~ l SO.values permitted by the exchange model assumptions (see text). The permitted depositional temperatures range from 0 ° to ~ 35 ° C, while the ~lSO-values of the possible depositional water range from - 6 . 6 to 0.0%o.
OXYGENISOTOPESIN IRON(III) OXIDES,2
represent, I will adopt the working assumption that their "percent Fe203" values are the appropriate units and yield valid end-member quartz and hematite ~ 180_values" These values (case A, Table I) are employed in subsequent discussions, but the model results for both cases A and B are presented. It should be mentioned that uncertainties in the calculated end-member ~ 180_values can be as high as _+1%0 due to scatter in the measured data, even if the appropriate compositional variable has been used. Fig. 2 contains a plot of end-member 318OHm VS. ~18OQtz (case A, Table I) for the six samples from the Urucum deposits analyzed by Hoefs et al. (1987 ). An interesting feature of the data array in Fig. 2 is the relatively large range in quartz ~180-values (+21.4 to +26.3%0) as well as hematite ~180-values ( + 0.9 to + 6.4%o). These relatively large d180 ranges for both quartz and hematite contrast with the small ~ 180 ranges for quartz and large 3180 ranges for magnetite noted by Gregory et al. ( 1989 ) for many quartz-magnetite-rich BIF of low metamorphic grade.
3. Model of isotopic exchange The presence of quartz-filled fractures of possible diagenetic origin in the Urucum deposits (Hoefs et al., 1987 ) suggests that water was moving through the deposits and that isotopic exchange could be modeled as an open system. The Urucum deposits may have been buried to depths of only ~ 2 km (Hoefs et al., 1987 ), which indicates relatively low temperatures of exchange. The work of Gregory and Criss (1986) and Gregory et al. (1989) emphasized the resistance of quartz to low-temperature retrograde oxygen isotope exchange. Furthermore, the results of Becker and Clayton (1976) and Yapp (1990b) indicate that hematite oxygen isotope exchange is not likely to occur in low-temperature retrograde processes. This suggests that the quartz and hematite of the Urucum deposits may have ac-
341
quired their 180/160 values during a prograde process such as burial diagenesis. It is assumed that the rate constant for prograde hematite-water oxygen isotope exchange (KRm) is much greater than that for prograde quartz-water exchange (KQtz). There is at present no experimental evidence to support this assumption, but the assumption is made to facilitate discussion and is in keeping with the apparent fact that KMt is much greater than KQtz in quartz-magnetite BIF of low metamorphic grade (Gregory and Criss, 1986; Gregory et al., 1989 ). Oxygen isotope disequilibrium in the Urucum deposits is suggested by the observation that samples from the highest stratigraphic levels (shallowest depth of burial) have the highest apparent temperature of formation (see U38 in Table I and also Hoefs et al., 1987). The extreme conditions of a "fluid-buffered" open system ( ~ ) and a water-absent closed system (Criss et al., 1987 ) will be employed in an attempt to set limits on the initial depositional environment of the Urucum BIF deposits. It is assumed that the observed quartz and hematite ~ 180_values were acquired during diagenetic disequilibrium isotope exchange. In addition it is assumed that the waters of the depositional environment were the input waters for diagenetic exchange, that all authigenic quartz and hematite of the deposit had sedimentary precursors which formed under identical equilibrium depositional conditions and that a probable maximum ~180-value for the presumed depositional water was 0.0%o. i.e. seawater (e.g., Epstein and Mayeda, 1953). The mathematical details of the model employed here can be found in Criss et al. ( 1987 ). If the quartz-water fractionation curve of Knauth and Epstein (1976) is combined with the hematite-water curve of Yapp (1990a), the following equation results: 1000 In OI~Q_H~---1.46" 10 6 T - 2 + 9.0
(2)
w h e r e OgQ_H-~ ( 180/160)Qtz / ( 180/160)Hm; and
T is in kelvins. The highest apparent tempera-
342
ture calculated for the samples of case A in Table I is given by U38 (135°C) using eq. 2. It will be assumed that this represents an estimate of the minimum temperature of isotopic exchange during diagenesis. Higher presumed exchange temperatures will not affect the conclusions of this study. Use of an alternative quartz-water fractionation curve, such as that of Clayton et al. (1972), in conjunction with the hematite-water curve of Yapp (1990a) would result in the calculation of systematically higher apparent temperatures for the samples of Table I. Although there is likely to be considerable compositional heterogeneity in the Urucum BIF deposits on the scale of hand samples, the lack of published average modal abundance data on the six samples of Table I requires that the limiting water-absent, closed-system exchange trajectory be approximated by the average compositional data reported for the composite BIF deposits. The average Fe content appears to be ~ 5 4 wt.% (Hoefs et al., 1987 ). This value for a system composed only of quartz and hematite corresponds to an XQtz/ nHm ratio of ~0.53, where X represents the oxygen in the indicated mineral as a mole fraction of the total system oxygen. The negative value of the ratio OfXQtz/SHm in closed-system prograde exchange corresponds to the slope of the exchange trajectory in ~Hm vs. ~Qtz coordinates [see Criss et al. (1987) for mathematical details]. The significance of this exchange trajectory in this discussion is that it represents an upper boundary condition on the proposed field of open-system prograde exchange. In particular the closed-system exchange trajectory represent the limiting conditions of p = 0 and X w = 0 in the open-system model, where/z is the normalized flow rate of incoming water and Xw is the oxygen in water as a mole fraction of the total oxygen in the system at any instant (see Criss et al., 1987). Sample U44 of case A in Table I corresponds to a limiting isotopic coordinate for water-absent closed-system prograde exchange in this dis-
C.J. YAPP
cussion. The equation of the chosen closedsystem prograde exchange trajectory through U44 is: 18OHm= -- 0.53~ 18OQtz+ 20.0
(3)
This equation is depicted as trajectory A in Fig. 2. The assumption that this is prograde diagenetic exchange allows the placement of constraints on the initial point of origin of trajectory A. The lower limit on the depositional temperature of the chemical sediment precursors to the quartz and hematite of U44 is presumed to be 0 ° C. Thus, extrapolation (dotted line) of trajectory A to the 0°C isotherm of Fig. 2 places a lower limit on the ~L80-value of the presumed depositional waters. This calculated 180-value for the initial water is -6.6%o. The dashed line in Fig. 2 labeled "Equilibrium Marine Precipitates" defines the locus of isotopic compositions of quartz and hematite in equilibrium at different temperatures with waters whose ~80-value is 0.0%o (i.e. presumably ocean water). It is assumed in the prograde exchange model presented here that the initial depositional waters did not have ~'80-values more positive than 0.0%o. Consequently, the intersection of closed-system trajectory A with the curve defined by equilibrium marine precipitates permits the calculation of a possible maximum temperature of deposition in the original chemical sediment. This intersection is indicated by the open square in Fig. 2 and corresponds to a depositional temperature of ~ 35 ° C. The other extreme in the continuum of opensystem isotopic exchange is represented by the limiting condition o f p ~ o o (fluid-buffered exchange) with the aforementioned assumption that grim >> gQt z. Open-system fluid-buffered isotope exchange with fluid whose fi180-value is 0.0%o at an exchange temperature of 135 °C is shown in Fig. 2 as trajectory B. The initial point for trajectory B is the open box in Fig. 2 representing the intersection of closed-system trajectory A with the locus of equilibrium ma-
343
OXYGEN ISOTOPES IN IRON (III) OXIDES, 2
rine precipitation. Note that all open-system exchange trajectories for which 0 > Kqtz will lie within the boundaries defined by trajectories A and B and the isotherm chosen as the temperature of exchange [see Criss et al. (1987) for mathematical details]. It is seen that all six of the data points for case A of Table I lie within this exchange domain or on its boundaries. This implies that the isotopic values of the initial point of trajectories A and B of Fig. 2 are representative of plausible original depositional conditions (i.e. ocean water at ~ 35°C). However, the point represented by the open square in Fig. 2 is not a unique initial value for exchange trajectories which define domains within which the data points of Fig. 2 will lie under the assumptions of this model. In fact, any initial point which lies on, or within, the boundaries of the stippled area in Fig. 2 will yield limiting closedand open-system exchange trajectories which encompass all of the data of Fig. 2 for a prograde exchange temperature not lower than 135°C. Thus, in the context of of this model the temperature of the original environment of deposition might have been anywhere between ~ 0 ° and ~ 35 ° C and the d t80 of the ambient water could have been between ~ - 6 . 6 and ~0.0%o. This proposed possible range of initial environmental conditions is only as credible as the assumptions which were employed to apply the exchange model. At this time there are no objective criteria which can be used to validate the assumptions. In addition there are the previously discussed uncertainties associated with the choices of appropriate calculated quartz and hematite d~80-values. For example, if the d~80-values of case B in Table I had been used in the exchange model, a maximum possible depositional temperature of 54 °C would have been indicated for precipitation o f the chemical sediments from ocean water with a d~80value of 0.0%o. The minimum possible calculated d~80-value of the environmental water would be ~ -8.90/00 at a depositional temper-
ature of 0 ° C using the data of case B of Table I. While the calculated ranges of depositional temperatures and water ~180-values for the Urucum quartz-hematite BIF are rather large, it is important to note that cooler temperatures and brackish or fresh waters in the environment of deposition are permitted by the exchange model. Walde et al. ( 1981 ) discuss the stratigraphy of the Jacadigo Group which contains the Urucum quartz-hematite deposits in the Band' Alta Formation. They note that the arkoses and conglomerates which underlie and are vertically transitional to the quartz-hematite deposits probably indicate a continental depositional environment. Furthermore, Walde et al. ( 1981 ) point out the existence of possible dropstones of glacial origin in the Band' Alta Formation. This sedimentologic evidence for the existence of contemporaneous glaciation and perhaps a nonmarine depositional environment is consistent with the cooler temperatures and freshwater components permitted by the stippled domain of initial values shown in Fig. 2. The exchange model applied to the Urucum data does not address the question of isotope exchange mechanisms. It assumes continuous isotopic exchange at a single temperature and does not consider the isotopic effects of mineralogic phase changes during diagenesis. Knowledge of the relative timing of such transitions for silica and ferric oxides during diagenesis is lacking, but may be of fundamental importance to the ultimate interpretation of 5t80 data arrays such as those of the Urucum deposits. 4. Conclusions
The use of oxygen mole fractions is of fundamental importance when calculating the 5 ~80-values of end-member components from the measured compositional and fi 180-values of binary quartz-hematite mixtures. Failure to use oxygen mole fractions can result in errors
344
of up to several per mil in calculated endmember 6 ~SO-values. The application of the kinetic isotopic exchange model of Gregory and Criss ( 1986 ) and Criss et al. ( 1987 ) to the quartz-hematite 6180 data published by Hoefs et al. ( 1987 ) for the Late Proterozoic BIF deposits of Urucum, Brazil, provides broad limits on the original depositional environment of these chemical sediments. Use of the hematite-water fractionation curve of Yapp (1990a) and the quartz-water curve of Knauth and Epstein ( 1976 ) in the model calculations yields a permitted depositional temperature range between 0 ° and ~ 35 °C (or perhaps 54°C). The indicated range of initial water 6180-values is between ~ - 6 . 6 ( o r - 8 . 9 ) and 0.0%o. The cooler temperatures and brackish or fresh waters in the environment of deposition which are permitted by the calculated model conditions are consistent with sedimentologic evidence for contemporaneous glaciation and a possible nonmarine depositional environment for these deposits. Gregory et al. (1989) were not sanguine about the possibility of deducing information on depositional environments from oxygen isotope data of Precambrian quartz-magnetite BIF. While not as pessimistic as Gregory et al. (1989), the conclusion to be drawn from the current treatment of the Urucum data is that the depositional conditions of largely unmetamorphosed Precambrian quartz-hematite BIF can, at present, only be constrained within broad limits using only oxygen isotope data. Acknowledgements The ultimate treatment of, and conclusions concerning, the data from the Urucum deposits were considerably influenced by the thoughtful review of Robert Criss, although he may not necessarily agree with the conclusions
C.J. YAPP
themselves. This research was supported by NSF grant EAR-8719070.
References Becker, R.H. and Clayton, R.N., 1976. Oxygen isotope study of a Precambrian banded iron-formation, Hamersley Range, Western Australia. Geochim. Cosmochim. Acta, 40:1153-1165. Blattner, P., Braithwater, W.R. and Glover, R.B., 1983. New evidence on magnetite oxygen isotope geothermometers at 175 °and 112 °C in Wairakei steam piplelines (New Zealand). Isot. Geosci., 1: 195-204. Clayton, R.N., O'Neil, J.R. and Mayeda, T.K., 1972. Oxygen isotope exchange between quartz and water. J. Geophys. Res., 77: 3057-3067. Criss, R.E., Gregory, R.T. and Taylor, Jr., H.P., 1987. Kinetic theory of oxygen isotopic exchange between minerals and water, Geochim. Cosmochim. Acta, 51: 10991108. Epstein, S. and Mayeda, T.K., 1953. Variation of 180 content of waters from natural sources. Geochim. Cosmochim. Acta, 4: 213-224. Gregory, R.T. and Criss, R.E., 1986. Isotopic exchange in open and closed system. In: Stable Isotopes in High Temperature Geological Processes. Mineral. Soc. Am., Rev. Mineral., 16: 91-127. Gregory, R.T., Criss, R.E. and Taylor, Jr., H.P., 1989. Oxygen isotope exchange kinetics of mineral pairs in closed and open systems: applications to problems of hydrothermal alteration of igneous rocks and Precambrian iron formations. Chem. Geol., 75: 1-42. Hoefs, J., Muller, G., Schuster, K.A. and Walde, D., 1987. The Fe-Mn ore deposits of Urucum, Brazil: an oxygen isotope study. In: N. Clauer and S. Chaudhuri (Editors), Isotopes in the Sedimentary Cycle. Chem. Geol. ( Isot. Geosci. Sect. ), 65:311-319 (special issue ). Knauth, L.P. and Epstein, S., 1976. Hydrogen and oxygen isotope ratios in nodular and bedded cherts. Geochim. Cosmochim. Acta, 40:1095-1108. Walde, D.H.G., Gierth, E., Clausthal-Zellerfled, and Leonardos O.H. 1981 Stratigraphy and mineralogy of the manganese ores of Urucum, Mato Grosso, Brazil. Geol. Rundsch., 70: 1077-1085. Yapp, 1990a. Oxygen isotopes in iron (III) oxides, 1. Mineral-water fractionation factors. Chem. Geol., 85: 329-335 (this issue). Yapp, C.J., 1990b. Oxygen isotope effects associated with the solid state a-FeOOH to a-Fe203 phase transformation. Geochim. Cosmochim. Acta, 54: 229-236.