Oxygen-isotope systematics in a multiphase weathering system in Haiti

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n Cosmochimica Ado Vol. 56. pp. 2831-2838 Q 1992 Pergamon PreyLtd. inU.S.A.

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Oxygen-isotope systematics in a multiphase weathering system in Haiti MICHAEL I. BIRD, I** FREDERICK J.

LONGSTAFFE,’ WILLIAM S. FIFE, ’ and PIERRE BILDGEN* ‘Department of Geology, University of Western Ontario, London, Ontario N6A 5B7, Canada ‘Laboratoire de GCochimie et MCtallog6nie(CNRS UA196) UPMC, 4 Place Jussieu, 75252, Paris, Cedex 05, France (Received November 20, 1991; accepted in revisedform April 13, 1992)

Abstract-Physical and chemical (partial dissolution ) techniques have been applied to a suite of young karst bauxite and laterite samples from the southern peninsula of Haiti. The 6 ‘*O values have been obtained for nine mineral species and range from +l.O% for anatase/rutile to +33.4% for authigenic quartz. On one hand, results for quartz, calcite, kaolinite, gibbsite, and boehmite compare favourably with 6 “0 values predicted from accepted mineral-water fractionation factors, assuming modem temperatures (25 f 2°C) and water 6’*0 values (-3.1+ 0.5%). On the other hand, the 6 “0 values measured for anatase, illite/smectite, chlorite/smectite, and some iron oxide samples do not compare favourably with predicted values. Departures from expected values for these minerals may be related to crystallization from water with a different 6”O value than modern water, an imprecise knowledge of some mineralwater fractionation factors under surficial conditions, or nonattainment of isotopic equilibrium between mineral and water during formation. Partial dissolution techniques hold considerable promise for obtaining quantitative 6’sO values of individual minerals in single samples of complex, fine-grained regolith material. INTRODUCTION PREVIOUSSTUDIESHAVE SHOWNthat the 6’*0 value of a mineral formed during weathering is related to the 6’sO value of the meteoric water from which it crystallized by a temperature-dependent fractionation factor; this behaviour has generally been taken to indicate a close approach to isotopic equilibrium between mineral and water (LAWRENCE and TAYLOR, 197 1, 1972; SAVIN and EPSTEIN, 1970; YAPP, 1990). The 6 ‘*O value of meteoric water is determined by gross climatic parameters such as temperature and the precipitation/evaporation regime ( YURTSEVER and GAT, 198 1) . Accordingly, the oxygen-isotope compositions of minerals from weathering (regolith) profiles can potentially provide insight into the continental palaeoclimate during their formation (BIRD and CHIVAS, 1988a, 1989; BIRD et al., 1989). Moreover, since many minerals are resistant to post-crystallization oxygen-isotope exchange with later waters, they faithfully preserve palaeoclimatic information (LAWRENCE and TAYLOR, 1971, 1972; BIRD and CHIVAS, 1988b). Despite this potential, there have been comparatively few isotopic studies of regolith deposits, mainly because of the following four problems: 1) The paucity of well-defined mineral-water fractionation factors for minerals at sur!iciaI temperatures; 2) The difficulty of obtaining pure samples from complex mixtures of fine-grained minerals; 3) The ambiguities that are commonly involved in drawing palaeoclimatic inferences from calculated 6 “0 values of ancient meteoric waters; and 4) The problem of demonstrating unequivocally that isotopic equilibrium has been achieved in either natural or experimental systems at surficial temperatures.

In this study, partial dissolution techniques are assessed as a means of determining the isotopic composition of individual minerals in samples where physical purification is ineffective. In addition, measured 6 “0 values for a suite of regolith minerals are compared with their predicted isotopic compositions, based on current knowledge of mineral-water fractionation factors. In this way, the accuracy of published fractionation factors can be assessed. The bauxites of peninsular Haiti provide a suitable test case because 1) the mineralogy of individual samples is comparatively simple, and yet a large range of authigenic minerals (nine in total) can be isolated for analysis; 2) weathering is comparatively young (- 100,000 years; BOULBGUE et al., 1989); 3) the bauxites are in an equatorial region where large temperature variations are unlikely; 4) the climate of the region is not monsoonal (monsoonal conditions can produce large excursions in the 6 “0 value of meteoric water, e.g., BIRD, 1988); and 5 ) the 6 “0 value of modern meteoric waters in the study area is well constrained and shows little variability (-3.1 + 0.5%; B~~LEGIJE et al., 1989). GEOLOGICAL BACKGROUND Lateritic weathering profiles and karst bauxite are widely developed on the southern peninsula of Haiti, and their mineralogical and chemical attributes have been extensively studied ( BILDGEN and HIERONYMUS, 1982; BILDGEN and DEICHA, 1982; BILDGENand BOULI?GUE,1985; BOUL~GUE et al., 1989). The basement geology of the region consists of a sequence of Cretaceous tholeiitic basalts, overlain by Late Cretaceous marine carbonates, Palaeocene turbidites, Eocene marine carbonates, and Pleistocene to Recent reefs. At some locations, the basalts have been thrust over the carbonate

* Present address: Research School of Earth Sciences, Australian National University, PO Box 4, Canberra, A.C.T., Australia. 2831

2832

M. 1. Bird et al.

sequence during uplift of the peninsula (BILDGEN and BOUL&GLJE, 1985 ) . Weathering of the basalts has led to development of thick lateritic profiles. Continued uplift and erosion of the peninsula caused lateritic detritus to be shed from the basaltic terranes into karst structures in the surrounding limestones. Further weathering of this detritus has produced commercial reserves of karst bauxite. Subsequently, the basalts were eroded to the point where they no longer contribute detritus to the carbonates. BGULeGUEet al. ( 1989) have shown that the modem Al, Ca, XOz, and Si chemistry of local groundwaters is consistent with the reaction kaolinite = gibbsite + dissolved silica, the latter partly precipitating and silicifying limestone beneath the karst. Geochemical balances show that the bauxites formed over a period of 40,000-160,000 years, whereas the peninsula has been above sea level for 300,000-700,000 years (BOUL~~GUE et al., 1989). The modem climate is hot, with mean annual temperatures of 25-27’C, and wet, with annual rainfall of up to 2000 mm and a distinct dry season. A 10.5 ka oxygen-isotope record from ostracod tests in Lake Miragoane (close to the study area) has revealed fluctuations of up to 1.7% in the 6’*0 values of lake waters ( HODELLet al., 199 1) . These variations were attributed to changes in the inflow/evaporation budget of the lake rather than variability in the oxygen-isotope composition of local meteoric waters. As in other tropical regions, temperatures during the last glacial period were probably a few degrees cooler than presently observed (e.g., COLINVAUX,1989). In the study area, lower regional temperatures would have been partially offset

by the lower (warmer) elevations of the plateaux in the past; the region is currently being uplifted at l-2 mm/ yr (DESREUMAUX, 1985). Therefore, while realizing that 25 + 2°C may understate the lower temperature boundary condition slightly, this value has been used in all subsequent calculations. An additional consideration is that ocean surface waters, and therefore the meteoric waters derived from them, are thought to have been - 1% enriched in I80 during glacial periods as a result of the sequestering of low- I80 ice in the polar icecaps (SAWN and YEH, 198 1). SAMPLES

AND EXPERIMENTAL

METHODS

Fourteen samples covering the range of available mineralogieswere collected from an area of approximately 40 by 20 km in the central peninsula region (Fig. 1). Bulk samples were airdried and ground to - I25 pm in an agate mortar. The mineralogy of each sample was determined by X-ray diffraction analysis (XRD) of oriented and random samples as required. The degree of aluminium substitution in iron oxides was determined by XRD, using the d-spacings of the goethite- 111 and hematite- 110ditfractions according to the techniques of %HWERTMANN et al. (1979) and %HUUE (1984). The chemistry of four samples was determined by X-my fluorescence spectrometry (XRF). The international reference bauxite sample BX-N was analysed concurrently, and the results compared favourably with recommended values. In some samples, pure mineral separates could be obtained by handpicking, magnetic separation, chemical treatment, or separation of the <2 pm fraction by standard sedimentation techniques. However, in many samples, small quantities of impurities remained. To correct for impurities, a weighed ahquot of the contaminated sample was treated to destroy the major phase, concentrating impurities into an insoluble residue. This residue was washed with I N (NH&CO3 and distilled water, dried and reweighed to determine the percentage

1

GEOLOGY AND SOILS OF PENINSULAR HAITI

1OOOkm

1 !xaLE

FIG. 1. Location of the study area. Sample localities (see Table 2) are as follows: ( 1) Plateau de Terre Rouge, (2) Plateau de Cavelier, (3) Plateau de Besace, (4) Plateau du Cap Rouge, (5) Plateau du Rochelois, (6) Plateau Savane Zombi, (7) Beloc, and (8) For&t des Pins.

2833

Oxygen isotope composition of minerals in laterite of impurities in the original sample (e.g., HODGESand ZELAZNY, 1980). The quantity of sample dissolved using any of the techniques in this study was reproducible to l-296, provided that exactly the same protocol was followed each time; similar results were reported by HODGESand ZELAZNY ( 1980). Poorer reproducibility can result if reagent concentrations or reaction times are varied. Details of the chemical treatments are given in Table 1; future references in the text to the treatments employ the descriptors provided in Table 1. Oxygen-isotope compositions of both bulk sample and residue were then determined, and the isotopic composition of the pure mineral calculated by mass balance, based on the theoretical oxygen yields of the minerals present in the bulk sample and residue. Oxygen yields of bulk samples and residues were measured directly as part of the fluorination experiments (see below). The uncertainty in the calculated isotopic compositions for pure minerals depends on ( 1) the difference in 6 “0 values between the mineral and contaminant, (2) the error in measuring the percent oxygen contribution, and (3) the analytical error of the d “0 values for both mineral and contaminant. The size of these errors has been assessed in footnotes to the tables. The use of chemical pretreatments, particularly at elevated temperatures, also leads to concerns that the oxygen-isotope composition of the sample will have been modified by isotopic exchange during the treatment. Sodium hypochlorite (at room temperature and SO’C) to destroy organic matter, and boiling HCI to destroy gibbsite and iron-oxide impurities, have been previously employed with no adverse effects reported (e.g., LONGSTAFFE,1986; AYALONand LONGSTAFFE, 1990). Iron oxides are resistant to isotopic exchange during repeated boiling in 5 N NaOH to remove silicate imourities orovided dissolution-reprecipitation does not occur ( YA~P, 1991j . YEH ( 1980) showed that acid treatment, followed by hydrogen-peroxide treatment and then dissolution of iron oxides using the sodium citratedithionite technique of MEHRA and JACKSON( 1960), did not affect the hydrogen-isotope (and by implication oxygen-isotope) composition of chlorite, kaolinite, ilhte, and smectite mixtures. In general, available evidence suggests that most minerals are resistant to oxygen-isotope exchange provided that dissolution-repmcipitation reactions do not occur. However, while partial dissolution techniques apparently do not cause direct isotopic exchange, few are strictly mineral specific. Some dissolution of other phases can be expected, which has the potential to affect the accuracy of results. Oxygen for isotopic analysis was liberated quantitatively from dried IO-15 mg samples by reaction with bromine pentafluoride at 550°C (CLAYTONand MAYEDA, 1963). Prior to reaction, samples were outgassed for 2 h in vacua at temperatures which varied from 110°C for gibbsite and iron oxides to 150°C for kaolinite and titanium oxides and 300°C for quartz. Oxygen was converted to carbon dioxide by reaction with an incandescent carbon rod, and the 6 ‘*O value of the

gas measured using a Micromass 602D mass spectrometer. Results are reported in per mil (L) relative to Vienna Standard Mean Ocean Water (V-SMOW). A mean value of +9.71 + 0.18% (1~; n = 16) was obtained for the NBS-28 quartz standard over the period during which samples were analysed. Analytical reproducibility for unknowns is somewhat lower, *0.2-0.25%0. RESULTS AND DISCUSSION

Samples analysed for this study exhibit a range of 6 ‘*O values in excess of 32k (Tables 2 and 3). Measured b’*O results for minerals are compared in Fig. 2 with predicted values, which were calculated using available mineral-water fractionation factors, a temperature of 25 +- 2°C and a meteoric water 6 “0 value of -3.1 + 0%~ The predicted ranges shown on Fig. 2 do not incorporate uncertainties inherent in published estimates of the mineral-water fractionation factors for surficial temperatures. Instead, our goal is to assess the potential magnitude of those uncertainties by comparison of theoretical ranges with measured compositions. Agreement is generally good for most minerals, but some discrepancies exist, as discussed below.

QAuthigenic quartz occurs in Haitian regolith profiles as veins, powdery fine crystallites, and large (up to 2 mm) euhedral crystals. It is found in the lower portions of profiles developed in lateritic detritus accumulated on limestone and in silicifying limestone beneath karst structures (BILDGEN and DEICHA, 1982). The 6180 values of all authigenic quartz samples (+32 + 1%o) lie within experimental error of the range of values predicted using the fractionation equations of CLAYTON et al. (1972), MATSUHISA et al. (1979), and

(1985). Oxygen-isotope fractionation in the quartz-water system is the most temperature-dependent of any common regolith mineral at sutlicial temperatures. The agreement between predicted and observed values for quartz suggests that both the magnitude and range of temperature and water 6 “0 values assumed in the calculations are good approximations. KITA et al.

Table 1. Chemical pretreatmenta employed in this study Descriptor

Reagents

Conditions

Result

Reference

NaOCl

Na-hypochlorite

room temperature overnight

removal of organic matter

amorphous

5% Na2C03 followed by acid ammonium oxalale

room temperalure overnight

removal of amorphous aluminosilicates; Fe and Ti oxides

Jackson (1956)

CBD

Citratedithionile-HCOj-

pH-7; 8O’C;1 hour

removal of crystalline Fe-oxide

Mehra and Jackson (IW)

05N NaOH

OSN NaOH

boil 2.5 minutes

removal of gibbsite and amorphous aluminosilicates

Jackson (1956)

5.ON NaOH

5.ON NaOH

boil 30-60 minutes

removal of aluminosilicates

KBmpf and Schwerlmann (1982)

HCI

6NHCI

boil 30-60 minutes

removal of Fe oxide, gibbshe, amorphous aluminosilicates

Long8taEe (1986) Ayalon and Longstaffe (1990)

fusion/dissolution

isolation of quartz

Syer8 et al. (1968)

dissolution at 48”C/l week

isolation of Ii-oxides

Sayin and Jackson (1975)

H2SiF6 H2TiF6

HCI-H2TiF6

Fitzpatrick et al. (1978)

M. 1. Bird et al.

2834 Table 2. Oxygen-isotope

Haiti-l HCI Haiti-2 HCI Haiti-2 HCI Haiti-3 HCI Haiti-11 H,SiFK Haiti-13 haidpgked Haiti-14 H#iFh

composition of other mmerals

powdery Qz (100) powdery Qz (100) vein Qz (100) vein Qz (100) QZ (100) euhedral Qz (100) Qz (100)

Haiti-4 handpicked Haiti-5 handpicked Haiti-12 handpicked Ha&-E HzTiFh Haiti-11 H 2TiF 6 Haiti-10 mags Haiti-11 mags Haiti-14 mags Haiti-7 ~2 pm CBD Haiti-9 c2 pm CBD Haiti-13 <2 pm HCI

5.1

l/S (95) Bo/Anfl
33.3 30.3 31.0 33.4 12.4 31.7 14.6

B&c, laterite Bewe, laterite Besaw, lakrite Besace, laterite Cap Rouge, karsl bauxlle Belw, lalaite Savane Znmbi, bauxite

28.1 26.2 26.1

spring,Vallei de la Gosselme spring, Cap Rou Diamant spring, K p Rouge

1.1 1.0

Cavalier, altered basalt Rcchelois bauxite

4.7 4.5 4.4

Terre Rouge, kaolinitlc deposit Cap Rouge, karst bauxite Cap Rouge, karat bauxite

;::I

Xi

21.2

22.0

Savane Zombi, altered basalt For& des Pins, altered basalt B&c, laterite

lrelative to V-SMOW in 9& ~rmr associated with analyses of pure minerals is +0.25%+; where a wrrecdon for the presence Of a contaminating mineral has been applied, the error is d *OS+&. ‘Mineral proportions have been estimatedby XRD (given in parentheses). Qz = qunz; Cc = calate: An = anatase; Ru = rutile: Mh = maghemire: Hm = hematite: Ba = bxhmile; l/S = illite/smcctite; Ka = kaolinite; C/S = chlorite/smectitc.

Two quartz separates from bauxite samples ( Haiti- 11 and Haiti-14) have considerably lower 6 “‘0 values ( + 12.4 and + 14.6%0; Table 2), indicating that this quartz is of residual origin. The quartz is probably derived from erosion of Tertiary elastic sediments in the region and was deposited on the karst surface along with lateritic detritus derived from basalt weathering. Calcite Calcite occurs as nodules and sparry coatings at springs which issue from the limestone terrane. Two calcite samples (Haiti-5 +26.2%0; Haiti- 12, +26.1 L ) have b “0 values in good agreement with values predicted from the fractionation curves of O'NEIL et al. (1969) and FRIEDMANand O’NEIL ( 1977). A third sample, from a spring in the Vallei de la Gosseline (Haiti-4; +28.1%0), has a higher b “‘0 value, suggesting that the spring water may have been partly evaporated prior to calcite deposition. BOUL!&JE et al. ( 1989) noted that evaporation had affected the oxygen-isotope composition of some water samples from the region. Clay Minerals A variety of clay minerals is present in the regolith profiles of the study area. The most common clay mineral is kaolinite, but illite/smectite and chlorite/smectite occur in weathered basalt samples and in black earth (Terre Noire) soils developed along the southern coast (BILDGEN and BOUL~GUE, 1985). Two kaolinite samples (Haiti- 10, +2 1.4%0;Haiti 13, +22.0%0) have 6180 values in good agreement with predictions from the fractionation equations of both SAVIN and LEE (1988) and LAND and DUTTON (1978). In contrast, illite/smectite (Haiti-7, +24.4%0) and chlorite/smectite (Haiti-9, f2 1.8%0) have 6 “0 values which are l-2%0 higher than even the maximum values predicted from the fractionation equations of SAWN and LEE ( 1988). These equations have been widely used with success, and it is not probable

that they are greatly in error. It is more likely that illite/ smectite and chlorite/smectite in the profiles formed under somewhat different conditions than kaolinite. Illite/smectite and chlorite/smectite are not common alteration products in humid tropical weathering environments. Their unusual isotopic compositions may have resulted from one of the following several mechanisms: ( 1) the minerals formed from slightly evaporated waters during an earlier, more arid phase (e.g., HODELL et al., 199 1); (2) the 6 I80 values of meteoric waters during the last glacial period were - 1‘%Ihigher than present, and cooler temperatures increased mineral-water fractionation factors, both effects favouring higher 6 180 values for minerals formed during that period; and (3) these clays are not of pedogenic derivation but instead formed in oxygen isotopic equilibrium with seawater during low-temperature submarine alteration of the basalts (e.g., HOWARDand FISK, 1988; MCMURTRY et al., 1983). Oxide Minerals Gibbsite with minor boehmite is the dominant Al-bearing mineral in the karst bauxites of the Haitian peninsula. The dominant Fe-bearing minerals in all regolith profiles are hematite and goethite, with minor maghemite. The dominant Ti-bearing phase is anatase, with lesser amounts of r-utile. Calculation of 6”O values for pure gibbsite from two of the three bauxite samples is complicated by the presence of crandallite ( CaAls ( P04b ( OH )5 - Hz0 ) , a common lateritic phosphate (SCHWABet al., 1989). This mineral is also readily dissolved by the 0.5 N NaOH treatment intended to dissolve gibbsite, and gibbsite 6”O values need to be adjusted accordingly. The proportion of crandallite was estimated from the P205 content of the bulk sample (Table 4); the calculated gibbsite 6 “0 value was adjusted assuming a crandahite 6 “0 value of +20.0%0 (see Table 3 for explanation). The 13~~0 value for gibbsite in Haiti-6 (+ 12.9%0) did not require correction for crandallite and is therefore the most precise value. Nevertheless, the corrected 6180 values for pure gibbsite in Haiti-l 1 (+12.6%0) and Haiti-14 (+13.2%+) are similar to

Table 3. Haitian bauxite oxygen-isolope Hati- 6 Fraction

raw amorphous free CBD OSNaOH lI;;&Oi$OH CBD/O.SNaOH

%loss

8.9 28.0 50.3 89.1 76.6 73.4

calculati gibbsbe cakulati inn oxides2 calculated kaolinbe

rcsulu aid-14

H&i-l1

Halli-10

6180.

%loss

aw

Iloss

b180

%lOss

12.2/12.2 12.0 14.3 il.0 14.0 15.7 15.9

IS.9 34.2 46.2 90.1 78.3 73.9

14.0114.1 14.3 15.9 13.4 n.d 17.0 16.3

7.1 33.6 32.8 88.7 70.6 62.3

13.4 13.8 15.4 12.9 13.3 15.7 16.1

9.5 25.5 7.5 88.8 33.7

12.9 2.5

12.6 7.5

13.2 6.8

@O

19.2 19.0 20.6 n.d 14.2 n.d

6.2 21.4

‘relative to V-SMOW m s; ad. = not determined this read1 has been calculated from asO, and a180,,5N NaOH assuming that the prOportion of gibbsbe in tbe sample is equivaknt to weight loss during chc O.SN NnOH treatment. Calculated gibbsite 6lsO values have been adjusted for the presence of crandallite in the samples @ble 4), assuming a alsO value for crandallite of +2C%a We are not aware of published cnndallite oxygen-isotope analyses; howcvcr, the assumed value is nasonable for a phmphatic mineral in a tropical environment (e.g., Ayliffe et al., 1992). The CTTOIfor cakulsted pure gibbsite alsO values is + +lBo. and blsOcsD assuming that the proportion of iron 2Tbis result has been c&ulated from6lsO, oxides in the sample is cquivaknt to weight loss during the CBD lreatment. Error for c&ulati rmre iron oxide 6180 values is 5 +2so. and 61s05,0N NaOH assuming that the proporl10n +f’his result has been calculated from @O,o of kaolinife in the sample is quivaknt to weight loss during the J.ON NaOH treatment. Enor for the calculated pure kaolinite 6180 value. is d +0.5!&.

2835

Oxygen isotope composition of minerals in laterite

calcite kaolinite illitelsmectite chlorite/smectitc

-+

boehmite

-+

I

gibbsite

&I

iron oxides anataselrutik

20.0

15.0

25.0

30.0

35.0

6180 (per mil) FIG. 2. Predicted (bars) and observed (crosses) oxygen-isotope compositions of regolith minerals from the southern peninsula of Haiti. The ranges of predicted ii’s0 values were calculated assuming a temperature of 25 + 2°C and a meteoric water 6 “0 value of -3.1 f 0.5%. The predicted range in d ‘*O values for individual minerals does not include errors inherent in the fractionation equations (seetext). Mineral-water fractionation factors were derived from the following sources. Quartz: (1)CLAYTON et al. (1972); (2)KITA et al. ( 1985); (3) MATSUHISA etal. (1979). Calcite: (1) FRIEDMAN and O’NEIL (1977); (2) O'NEIL et al. (1969). Kaolinite: ( 1) LAND and DUTTON(1978): (21 SAVIN and LEE (1988).Illite/smectite: SAVIN and LEE ( 1988).Chlorite/smectite: SAVIN and LEE(1988). Bo&n&: BIRD etal.( 1989). Gibbsite: CHEN et al. ( 1988): BIRDet al. ( 1990). Iron oxides: YAPP (1990).Titanium oxide: c 1j anatase (Zheng, pers. commun.); (2) ru&: ZH&G ( 1991). &au& there is no fractionation equation availabld fbr either gibbsite or boehmite (only a single fractionation factor applicable to “surficial temperatures”), the error associated with the temperature dependence of fractionation has been assumed to be +0.5% in the range 25 + 2’C. Errors associated with measured d “0 values depend on the proportion and isotopic composition of contaminants in each sample (see.text). For pure phases (quartz, calcite, anatase/rutile, and “magnetic” iron oxides), analytical error is kO.25960.Where a correction has been applied, the error increases as follows: clay minerals and boehmite 5 +0.5%; gibbsite 5 +I .Ok; “nonmagnetic” iron oxides I k2L.

the value for Haiti-6, and all lie within the predicted range of values. Only one analysis of boehmite was possible (Haiti-6, +15.8%0). Its 6’*0 value is consistent within error of the fractionation factor proposed by BIRD et al. ( 1989). An oxygen-isotope fractionation curve for iron oxides (hematite/goethite ) has been determined by YAPP ( 1990). Calculated 6 ‘*O values for iron oxides ( fine-grained mixtures of hematite, gcethite, and minor amorphous material) from this study show a comparatively wide range, extending to values considerably higher than predicted, even allowing for the comparatively large error associated with the calculation of d “0 values for iron oxides by mass balance (up to +2460;see Table 3). Aluminium commonly substitutes to some degree for iron in both hematite and goe.thite; such substitution may explain some of the higher values. The maximum observed Al-substitution in this study is 2 1.2%; but because the dominant mineral in all samples is hematite, the maximum degree of Al-substitution for bulk iron oxides in any sample probably does not exceed 10% (Table 5 ) . The effect of Al-substitution can be calculated by assuming that the fractionation factor for goethite (or hematite) increases linearly towards that of

boehmite with increasing Al-substitution. For 10% Al-substitution, the maximum predicted 6 180 value would increase from +3.6 to +4.7%0. However, this increase is insufficient to explain still higher 6 “0 values of several samples (up to +7.5?50). We suggest that evaporation may have been important in the formation of such iron-oxide samples; such an effect has been noted for other regions (BIRD et al., 1992). Another possibility may be that amorphous iron oxides present in the samples may have substantially different 6’*0 values than the crystalline iron oxides. There is a large difference between the oxygen isotopic composition of magnetic iron oxides, which exhibit a small range (+4.4 to +4.7’S), and the bulk iron-oxide 6 I80 values for the same samples (+6.2 to +7.5%0). This variation may result from a difference in the fractionation factor between maghemite and hematite/goethite or may indicate different modes of formation for these phases. For example, maghemite may have formed as a pseudomorphic replacement of primary magnetite (e.g., ANAND and GILKES, 1984; MORRIS, 1983) rather than precipitating from solution, the latter being the likely mechanism for formation of most hematite and goethite in the samples.

2836

M. I. Bird et al.

Theoretical fractionation curves for rutile and anatase have been determined recently ( ZHENG, 1991; Zheng, pers. comm.). Two anatase samples (with minor rutile) analysed for this study have very similar 6 “0 values (Haiti-8, + 1.1%o; Haiti- 11, + 1.O%O ), but these results are lower than predicted values by - 1L. Whether this behaviour reflects crystallization from lower ‘*O waters (or at higher temperatures), an imprecise estimate of the fractionation factor, failure to achieve isotopic equilibrium during formation, or modification of the anatase 6 ‘*O value during the H2TiF6 pretreatment is not known. Bauxite Three bauxite samples and one sample of lateritic detritus (on karst) were investigated in detail. The effects of various dissolution techniques on the isotopic composition of each sample are summarized in Table 3. Surprisingly, the removal of significant quantities of amorphous or poorly crystalline Al-W and/or Fe-compounds had little effect on the 6”O value of any sample. Amorphous aluminosilicates are likely to have considerably higher 6 I80 values than amorphous iron oxides. The similarity of 6 ‘*O values between raw and amorphousfree samples suggests that contributions from amorphous Fe and Al-Si minerals balance each other.

Table 4. Chemical and mineralogical compmitmns of selected samples Sample

(BX-N’)

Chemical composition (wt.%) SiO, La.19 7.61 17.40)

Haiti-6

Haiti-l 1

Haiti-14

Haitl-10

5.00 1.82 48.38

Il.10 I.90 36.30

‘2.19 2.42 40.84

33.49 2.17 31.34

21.90

17.50

21.32

0.06 0.04 0.00 0.00

0.90 0.38 1.86 0.10

0.30 0.46 0.54 0.06

18.05 0.18 0.58 0.20 0.24

0.20 22.52 0.00

3.32 26.39 0.00

1.36 20.38 0.00

0.21 0.00 13.47

99.92

99.75

99.87

99.93

Mineralogical composition (wt.%)3 0.1 quanz 1.6 anatas&tile 0.7 cranjailite 10.7 kaolinite 0.0 illit&mectite 28.0 iron oxides 50.5 gibbsite 8.2 kehmite

1.9 2.0 11.9 18.6 2.8 34.2 25.4 3.2

0.5 2.5 4.7 24.4 1.6 33.6 26.0 6.8

1.8 2.2 0.7 62.6 6.3 25.5 0.0 0.9

calculated% H20+(4) obserwd% HzO+

21.4 21.5

15.7 21.2

15.8 17.3

10.7

calculakd% gibbsitc observed% gibbsit&

51.2 49.6

27.2 34.3

26.7 28.1

0.0 6.8

12.1

14.6

14.5

18.8

12.2

14.0

13.4

19.2

2.22 i2.37) TiG a01 A@& +0.08 54.57 (54.21) Fc,Oq do.06 22.72 (23.17)

0.11 (o.i7j

CG3 an K20

+a03

0.06

(0.05)

PZOS *cl.02 0.14 Na&o to.10 12.34 0.00

(0.13) (0.04) (12.36)

Total

Cahhled

loo.06 (lW.o6)

bW,,6

Observedb@O,

‘XRF Sardardbaune Bx-N (Go”m*a,u, 15-a), rccLlmmewed values1”brack8. W+DStandard~rflQ~.

_.

11.7

Table 5. Al-substitution in hematite and gwtbite sample Haiti-6 Haiti-10 Haiti-11 Haiti-14

%Al in hematite 10.2 9.4 9.5 6.2

%Al in gcethite 1.4 21.2 0.6 0.6

CBD-treated samples lost 25.5-34.2% oftheir initial mass, corresponding to the proportion of iron oxides in the sample. The 6 “0 values of CBD-treated samples rose by -2%0 because of loss of low- I80 iron oxides. In contrast, the 6 I80 values of 0.5 N NaOH-treated bauxite samples decreased by l-2%0 due to loss of gibbsite (+ crandallite). Although no gibbsite is present in Haiti-lo, the sample lost 7.5% of its mass during the 0.5 N NaOH treatment, corresponding to loss of amorphous aluminosilicates or poorly crystalline kaolinite. The HCl/S.ON NaOH treatment removed -90% of all samples, leaving a residue of anatase/rutile + quartz + boehmite + chlorite/smectite + illite/smectite. The 6 ‘*O values of the residues range from + 13.3 to + 14.2%~. Both the CBD/HCI and CBD/O.S N NaOH treatments should dissolve the same minerals from the samples (amorphous material, gibbsite, iron oxides, and crandallite), yet the former treatment resulted in 3-8% more weight loss. This behaviour suggests a difference between the two methods in the selectivity or severity of chemical attack. However, only in one sample (Haiti- 11) did the 6 “‘0 values of the two fractions differ by more than experimental error. Weight losses from the CBD/O.S N NaOH treatment are 4-6% less than the sum of the weight losses obtained from the CBD and 0.5 N NaOH treatments when used alone. This discrepancy probably results from accumulating errors (minor sample losses during multiple centrifuge washings, etc.) in performing the two treatments separately on different aliquots, rather than a real difference in the selectivity of the two methods. Given the results from the partial dissolution experiments plus chemical compositions for the samples, it is possible to estimate their mineralogical composition. From the mineralogical composition, it is then possible to calculate parameters such as H20+ and bulk 6’*0 values, which should compare favourably with observed values (Table 4). Because the gibbsite concentration in the mineralogical calculation is assumed to be that proportion of the total mass not assigned to other minerals (see Table 4, footnote 2), the calculated% gibbsite can be compared to the measured% gibbsite (defined as weight loss during 0.5 N NaOH treatment minus %crandallite) as another check on internal consistency. The predicted and observed values agree very well for the sample with the simplest mineralogy (Haiti-6). They also agree well for Haiti-lo, except that the mineralogical calculations indicate no gibbsite (confirmed by XRD) while the “observed” gibbsite concentration is 6.8% (arising from dissolution of poorly crystalline or amorphous aluminosilicate material). The comparatively good agreement between predicted and measured values for Haiti-6 and Haiti-10 is not surprising as a large proportion of the total oxygen in these samples is derived from a single mineral (>60% from gibbsite in Haiti-6, >70% from kaolinite in Haiti-lo).

2837

Oxygen isotope composition of minerals in laterite

The agreement is not as good for samples with a more complex mineralogy (Haiti- 11 and Haiti- 14)) where the calculated% gibbsite and, as a result, HzO+ are underestimated, and bulk 6 ‘*O values are overestimated. A rigorous assessment of the errors associated with the calculations in Table 4 was not attempted because we lack information on ( 1) degree of departure from the ideal stoichiometries assumed in the calculations, (2) variations in the hematite/goethite ratio, and (3) error in the assumed crandallite 6 “0 value. Nevertheless, the readily quantifiable errors associated with the measured chemical and mineralogical compositions of the samples, and the calculated 6 ‘*O values of pure minerals, are sufficient to accommodate the observed discrepancies. In sample Haiti- 11, the high proportion of crandallite may have exacerbated the discrepancy between calculated and observed results. Besides hydroxyl groups, crandallite also contains water of hydration, which may contribute to the particularly high observed HzO+ content ofthis sample. This possibility is difficult to evaluate further, however, because of uncertainty regarding the temperature at which crandallite dewaters. Crandallite in Haiti- 11 also contributes - 13% (or more if hydration water is present) of the oxygen in the bulk sample. Any error in the &I80 value assumed for crandallite will have a considerably greater effect on the 6 I80 value calculated for this bulk sample than for the others. CONCLUSIONS This study has investigated the utility of partial dissolution techniques for determining the oxygen-isotope composition of minerals in samples where physical separation techniques cannot produce a pure separate for analysis. Broad agreement has been demonstrated between most observed mineral 6 ‘*O values and those predicted from available fractionation factors. Most discrepancies can be explained adequately by processes that may have operated during formation of particular minerals (e.g., evaporation) or by minor changes in the 6 ‘*O value of meteoric waters in past glacial climates. However, the possibility that some fractionation factors are in error cannot be ruled out. In addition, where predicted and measured values do not coincide, it may be that equilibrium was not attained during mineral formation (perhaps the case for rutile/anatase?). The degree to which kinetic, rather than equilibrium, fractionation influenced the isotopic composition of minerals analysed in this study cannot be assessed from our data. However, if kinetic fractionation was important, the mechanism was both reproducible and predictable in many cases; the predicted values reported in this study were calculated using fractionation equations derived from a number of diverse natural and experimental systems. Further work is required to better define the potential consequences of some chemical dissolution techniques, both in terms of their selectivity and their effect on the oxygen isotopic composition of minerals remaining after the treatment. The accuracy of chemical dissolution techniques can be improved by careful field selection of samples and by physical pretreatments designed to produce as pure a sample as possible, prior to the application of chemical methods. Acknowledgments-We

thank C. Wu for the XRF analyses, P. Middlestead for laboratory assistance, and the Natural Sciences and En-

gineering Research Council of Canada for financial support. C. J. Yapp, J. R. Lawrence, and an anonymous reviewer provided thoughtful reviews of the manuscript. Editorial handling: G. Fame

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