The geochemistry of the stable isotopes of silicon

Ceochimica

0 Pergamon

er Cosmochimica Aclo Vol. 46. pp. Press Ltd. 1982. Printed in U.S.A.

0016.7037/82/08l449-lOSO

1449 to 1458

00/O

The geochemistry of the stable isotopes of silicon C. Division

of Geological

and Planetary (Received

Sciences,

B. DOUTHITT*

California

Institute

of Technology,

Pasadena,

CA 91125

USA

October 22, 1981; accepted in revised form March 29, 1982)

Abstract-One hundred thirty two new measurements of the relative abundances of the stable isotopes of silicon in terrestrial materials are presented. The total variation of 6-“Si found is 6.2%0, centered on the mean of terrestrial mafic and ultramafic igneous rocks, 6”‘Si = -0.4%0. Igneous rocks show limited ( 1.1 %v) variation; coexisting minerals exhibit small, systematic silicon isotopic fractionations that are roughly r/r the magnitude of concomitant oxygen isotopic fractionations at 1150°C. In both igneous minerals and rocks, @‘Si shows a positive correlation with silicon content, as does 6’*0. Opal from both sponge spicules and sinters is light, with z3’Si = -2.3 and -1.4%0, respectively. Large b3’Si values of both positive and negative sign are reported for the first time from clay minerals (-2.3 to +1.8%0), opaline phytoliths (-1.4 to +2.8%0), and authigenic quartz (+1.4%0). All highly fractionated samples were precipitated from solution at low temperatures; however, aqueous silicon is not measurably fractionated relative to quartz at equilibrium. A kinetic isotope fractionation of ~3.5% is postulated to occur during the low temperature precipitation of opal and, possibly, poorly ordered phyllosilicates, with the silicate phase being enriched in ‘*Si. This fractionation, coupled with a Rayleigh precipitation model, is capable of explaining most non-magmatic @‘Si variations. Chert d3’Si values are largely inherited, but the primary opal 630Si values can be modified by isotopic equilibration of silicate silicon and dissolved silicon during the transformation of opal into quartz.

INTRODUCTION IN the 1950’s, variations in the relative abundances of %i and “Si in nature were investigated several times, both experimentally (Reynolds and Verhoogen, 1953; Marsden, cited in Rankama, 1954; Allenby, 1954; Tilles, 196 1a,b) and theoretically (Grant, 1954). The data of Marsden and Allenby are of historical interest only, in that they indicate a large range of variation which has not been substantiated by any later studies. The lunar sample return program and, more recently, the search for isotopic heterogeneities in meteoritic silicon have sparked a resurgence of interest in the variations of the relative abundances of all three of the stable isotopes of silicon (Epstein and Taylor, 1970a,b, 1971, 1972, 1973, 1974, 1975; Taylor and Epstein, 1970a, 1973a,b; Epstein and Yeh, 1977; Yeh and Epstein, 1978a,b; Clayton et al., 1978a,b; Becker and Epstein, 1981). All prior studies of terrestrial samples have indicated that silicon isotopes do fractionate in a variety of environments, but systematic variations have not been demonstrated except for igneous materials. This paper presents 132 new GJoSi analyses of terrestrial materials, a review of published s3’Si data, a discussion of postulated kinetic and equilibrium fractionations, and a scenario that relates most documented terrestrial populations. EXPERIMENTAL Silicon technique

PROCEDURES

was extracted from samples using the fluorine described by Taylor and Epstein (1962a). Under

* EXXON Minerals Drive, Suite A, Tucson,

Company, AZ 84745,

2425 U.S.A.

N.

Huachuca 1449

standard reaction conditions, the only silicon-bearing species produced is SiF,; neither SizFa nor silicon oxyfluorides are produced. Separation of SiF4 from byproduct O2 and unreacted F2 is effected at liquid nitrogen temperatures. Fluorination of impure samples can, however, produce volatile fluorides and oxyfluorides of carbon, sulfur, phosphorus, and hydrogen that are hard to separate from SiF, (Epstein and Taylor, 1971; Clayton et al., 1978). In many cases, therefore, samples were chemically treated prior to fluorination to extract and purify their silicon. The chemical treatment involves fusion of a mixture of sample and Na2COJ in a platinum crucible at 1 lOO”C, followed by dissolution of the mixture in 7.4 N HCl, which causes precipitation of hydrated, amorphous SiO*. The precipitate is collected and washed in 3.0 N HCI twice, in 1.5 N HCl several times, in distilled water several times, and desiccated at 1200°C for thirty minutes. Yields were commonly better than 90%, with the missing fraction attributable to mechanical losses and the finite solubility of silica. The procedure is a modification of techniques described by other workers (Reynolds and Verhoogen, 1953; Tilles, 1962a; Yeh and Epstein, 1978a). This treatment renders the samples useless for 60 determinations, but &Si determinations on samples before and after treatment indicate that no measurable silicon isotope fractionation accompanies this step. 6’sO data are reported in addition to h3“Si data for most of the chemically untreated samples; the determinations were generally made on the same sample. Oxygen was collected using the technique described by Taylor and Epstein (1962a). The mass spectrometer used in this study is a double collecting, 60” sector, 22.86 cm. radius Nier-type mass spectrometer with an all metal sample introduction system, except for the magnetic switching valves (Epstein and Taylor, 1970a). The standard for silicon work is the Caltech Rose Quartz Standard (hereafter as RQS) (Epstein and Taylor, 1970a). This standard has been shown to be homogeneous with respect to 630Si and 6’sO and ha8 been used for all silicon isotope work done by Epstein and coworkers. The results of the isotopic ratio analysis of both silicon and oxygen are reported in h-notation in parts per mil

C B. DOUTHITT

1450 Table I

630Si analyses and sample descriptions

63OSi

t

*

‘74 ( )’ P Sample

Mariana Island Arc: la lb 2 3 4 5

-0.6 -0.4

6 7 8

-0.4 -0.5 -0.4

1;:; -0.7 -0.5

description

____ (6'80) 'P

islands (DIXON and BATIZA, 1979)

basalt dike interior, AQrigan(+S.xj (tS.61 chilled margin, $1 (+5.6\ basalt, Asuncion (t5.n) basalt, Asuncion basalt, Guguan (+F.6i basalt, Pagan /;$:;]

1;;

basalt, Sarigan Pagan basalt, andesite, Sarigan

(+;.?i (16 i:

Mariana Island Arc: seamounts (DIXON, i'i'iii) Qa 9b 10 11 12 13a 13b 14 15 16 17 18

basaltic andesite flc:wtop base, $9 basaltic andesite dacite andesite basalt, glassy rind interior, t'3 basaltic andesite basaltic andesite basaltic andesite basalt basaltic glass

-0.4 -0.3 -0.5 -0.4 -0.4 -0.4 -0.4 -0.5 -0.5 -0.4 -0.3 -0.1

(t5.4) jC5.6) (t6.21 (t6.5) ri-5.R' (+6.2j (+5.7) (tb.0; (+5.81 jt5.8) i+5.9; (15.4:

Volcanic suite, Hackberry Mountain, Yavnpai Lo., Al 19 20 21 22 23 24 25 26 27 28

-0.2 -0.7 +0.1 0:;

basalt basalt andesite andesite dacite dacite dacite dacite dacite dacite

(3)

0.0 0.0 0.0 -0.2 +0.2

(t7.G; (t6.9; tt7.4: (11o.n: (t8.9) (t9.n (*11.2) (+lO.O) (iR.4

Lachlan Fold Belt, Australia 29 30 31 ;'3 34 35 z

* * * * * * * *

@ O @ 9 @ @ @ @

to.2 -0.3 -0.2 -0.8 -0.5 -0.6 -0.2 to.3 -0.3

I-granite I-granodioritc I-tonalite I-granite S-aplite S-granodiorite S-granite S-granite S-granite

:z 41

* * * *

-0.5 -0.6 -0.5 -0.2

; :i.:

(+10.6)

nephelinite alkali gabbro @ syenlte ekerite

.

(2) 9 basalt, chilled margin basalt (LZ) ferrogabbro (UZc) granophyre, Tinden Sill

-0.5 45 * 0.0 (also see %60)

*

Oka Carbonatite complex, Canada 46 * -0.3 sovite okaite * -0.5 * -0.4 meltelglte z 49a * to.2 (2) @ fourchite @ syenite ocellus in #49a 49b * to.2 (also see #57) (Miscellaneous) igneous rocks and minerals ::

* -0.4

52

* -0.4

53

* -0.3

:: ::

* -0.4 0.0 * -0.4 0.0

58

*

0.0

59a

0.0

59b

-0.2

60 2:

%., ( 1’ 13Sample

*

* -0.82 *

0.02 0.0

63

* -0.2

64

* -0.1

65

* +O.l

MOR basalt, S. Atlantic flood basalt, Columbia River plateau yr;eftii!;ridotite,Bellefontein, olivine, dunite, N.C. (U.C.-1 -0.30 + 0.15%*)'+ mangerite, Lofoien Islands, Norway graphic granite, Arendal, Norway obsidian, Long Valley, CA fenftized gnefss, Chilwa Island, Malawi Precambrian composite, southern Norway (JACOBSEN, 1981) quartz, quartz monzonite, Eagle Mts., CA K-feldsoar, t59

descriptions ,_______ ,.,.ra‘

pyroxene, Skaergaard pyroxene, British Columb'd Rose Quartz Standard, Pala, !:A ;+,s.qc quartz from "core", RuQgIes Pegmatite, NH quartz from "pocket', ii:: Pegmatir.13. Evje, Norway (2) @ opal from garnet skarn, rlrngham L :,+20.!'! Canyon, IUT

Clay minerals, sedimentary rocks, non-igneous $ua?rz 66

-0.2

67

68

-0.2 -0.2

67

* to.1

68

* Cl.4

69 70 71a 71b 72

* * * * *

73

* -0.4

74

* -0.2

:6" 77 78 :

* * * * * *

-0.4 -2.1 -0.5 0.0 -0.9

-0.1 -0.8 -0.1 -1.0 -2.3 +1.8 -0.4

Sinters

Skaergaard intrusion, Greenland 42 ::

#

81

Oslo Graben, Norway 38

Table I (cont.) 630Si analyses and sample desc!:xions I_:*_18, 8 630Si

a2

* -1.6

83

* -0.4

84 85 86 67

* * * *

88

* -3.1

89 90a 90b 91

* * * *

-1.5 -1.9 -2.2 -0.9

92 93 94 95a

* * * *

-0.3 -0.3 to.4 to.4

95b 95c

* +0.1 * to.9

-1.8 -1.6 to.5 -0.5

orthoquartzite, Eagle

Mts , ,..A (,+\I i:i Mts,, CA

orthoquartzite, Eagle orthoquartzite, Pinto Mts , CA {+lO,::' 2;;;-2,;iOOum fraction, AnTrim _ (2) iL24.H) quart; (Herkimer "diamond'). NY (3) (+ ?2 : sepiolite, Nickel Mountain, OR garnierite, Nickel Mounta'n, 0i: smectite, lake bed, NV z sepiolite, coexisting w'tb rlla illite, H-36, Ward's f4aturalSriewi Establishment nontronite, H-33a, Ward's 'natural Science Establishment Ca-montmorillonite, Source iidv Mineral Repository, Univ. ilf Missouri (hereafter as SCMRI Na-montmorillonite, SCMR attapulgite, SCMR halloysite, SCMR kaolinlte (well crysta'lized), SCMK kaolinite (poorly crysta'lized)SCMR (2) (2) kaolinite, Long Valley, CA allophane. New Zealand (hereafter as N.Z.) (WELLS et al _ 19771

(2)

(2)

Yellowstone, W‘ 9zerrlf'-1.34 t 0.2%.)i geyserite, Narciysus Geysel,. Yellowstone, WY geyserite, Beowawi, NY geyserite, Wairakei, N.C. opaline sinter, N.Z. opaline sinter, Ohaki pdo'. aroadlands, N.Z. opaline sinter, Champagne pool, Waitaupo, N.Z. sinter, Whakarewarewa, N.L: opaline sinter, Arakei Karoko, N.2. opaline sinter, Arakei Karoko, N.Z. opaline slnter, Steanmoat Springs (hereafter as S.S.), NV !?+20)4 chalcedonic sinter, S.S. chalcedonfc sinter, S.S. silicified plant casts, S..>. chalcedony, drill hole GS-5-'5'.

S.S.

chalcedony, GS-5-28', S.. chalcedony and D-cristobalite, ii.11 j)i GS-5-84'

Dissolved si'1iCOtl.hot springs 96

* to.1

97

l

98

*

99 100

to.3

(2)

+0.1 * -0.1 * to.4

Cistern, Yellowstone, WY (490 pprnSi02: Porcelain Terrace, Yellowstone (500 ppm SiO2) Beryl, Yellowstone (270 ppm Si02) Ear, Yellowstone (385 ppm SiO2Y Minerva Spring, Mammoth, CA (55 pornSiO i

7’

Biogenic silica 1Ola * -2.1 1Olb * -2.0 102

* -0.9

103

* -3.4

(2)

sponge, Microciona sp.. Woods Hole. MA sponge, Microciona sp.. kioods:ioie.

sponge, Hexactinellida, LY;,*“! California

ISOTOPES Table I (cont.) 630Si analyses and sample descriptions 630Si %. *

# 104 105 106 107 108 109

* * * * *

110

* +1.42

l

to.8 +1.7 +z.a -1.4 -0.4 -0.9

( )’ @Sample

descriptions

(6;:)

phytoliths, Equisetum sp., CA phytoliths, Equisetum sp., CA phytoliths, Equisetum sp., CA phytoliths, Bambusa sp., CA diatomite, Monterev Shale, CA flint, Dover Chalki England (U.C.-16, -1.14%.) chert

(2) (2) (2)

O'NEIL and CHAPPELL, 1977 R. BECKER, pers. coann.,1981 D. SIBLEY, written comm., 1978 J. O'NEIL, pers. coma., 1981 J. GOSS, pers. comn., 1981 Silicon was chemically extracted prior to fluorination Parentheses contain number of fluorinations, if more than one Precision is 0.6%, at 20 Sample analyzed by TILLES (1961a& parentheses enclose sample designation and 6 Si Se#ple analyzed by REYNOLDS and VERHOOGEN (1953) & Si recalculated by TILLES (1961a); parentheses e lose original sample designation and recalculated 6&i

(%),

where: &LE -_

bSi =

RRQS

where

R-2,

“osi 28Si

’

-I80 I60

ASi/\_,, = 653, - 8SiB asi

_

As-

_

1 + w1000 1 + 6,/1000

During a a30 analysis, “Si and %.i are collected in the same Faraday cup. The correction for dual collection is small, ~30SicoasncTsn = 1.025 6’%iMEAsosno [after Craig (1957) assuming that 6*%i = 0.5 a3’Si, and that SiF: is the only species being collected], and has been overlooked by all previous workers. All data taken in this study or borrowed from published sources (except Tilles, 1961a,b) are reported here as b”“SicoaancTEo. Precision of b”Si data, based on repeated analysis of the same gas and repeat fluorinations of both untreated and chemically processed RQS and of silicon metal, is about 0.30% at 20. Precision for 6*?Si is somewhat higher, about 0.27~ at 20. Analyses made during periods of degraded precision (estimated as 0.6% at 20) are indicated in Table 1. The oxygen isotopic data are reported on the SMOW standard by comparison with the results of repeated analyses of RQS, for which a value of 6’sO = +8.45’% is assumed. Precision is estimated to be 0.2% at 20 based on RQS replications, but is somewhat less for analyses of “whole rock’ oxygen. NBS-28 has a value of 6180 = +9.60% on this scale. 6*%i was routinely measured on all samples. As previously reported by Yeh and Epstein (1978a) and Clayton et al. (1978b), no deviation of S*%i of terrestrial samples from the mass fractionation relationship, 6*‘Si N 0.5 b’OSi (Clayton et al., 197813; Matsuhisa et al., 1978), were found that could not be ascribed to the contribution of a contaminant fragment to an analysis peak (Becker and Epstein, 1981). All b?Si analyses are tabulated, along with brief sample descriptions and 6’*0 data where available, in Table 1 and shown in Fig. 1. The histograms of Fig. 1 also include data

1451

OF Si

reported by previous workers, as credited in the caption. Two samples analyzed by Tilles (1961a) were analyzed here (Table 1, #53, 82). Both sets of analyses agree within the combined experimental error, justifying the inclusion in Fig. 1 of both the uncorrected data of Tilles (1961a) and the data of Reynolds and Verhoogen (1953) as corrected by Tilles ( 1961a), with some salient exceptions to be discussed later. Neither these data nor data of the present study reported with precision worse than 0.30’% were used in calculating population means. Sample population means are denoted 8 and, for N > 4, are reported flo. DISCUSSION Igneous

OF ISOTOPE

DATA

rocks and minerals

All terrestrial igneous “whole rock” 630Si analyses are plotted in Fig. la. Systematics do exist in these data, but the comparable magnitude of the 6”Si variations and experimental precision vitiate the significance of single analyses. The emphasis in this study was on the analysis of variations within and between rock suites. The total range of variation found is 1.1%~ Subaerial and submarine volcanic rocks from the Mariana Island Arc (Table 1, #l-18) show a 0.6%0 spread, with a normal distribution of values about ??‘Si = -0.4 + 0.2%. 8”O values show limited spread, 1.2%, about 6”O = +5.9 + 0.3%. No correlation exists between BsoSi and either d’*O or silica content. The Hackberry Mountain, Arizona suite (Table 1, #1928) was collected from a Tertiary silicic volcanic center. Most of these rocks have been somewhat hydrothermally altered. The mean of all samples is $“Si = -0.1 f 0.20/W. The mean of just the andesites and dacites is s3’Si = -0.1 + 0. l%, and is just distinguishable at the 10 level from the mean of the Mariana Island Arc. 8’*0 values from the same suite show a 4.4% spread about 8’*0 = +8.8 f 1.4%. Granitoids from the Lachlan Fold Belt, Australia (Table 1, #29-37) show a wide range in values, from @“Si = -0.8 to +0.3%. These analyses were made during one of the periods of degraded precision, but do indicate that granitoids fall in the “normal” igneous rock range of 6%i values; that hA’Si, unlike 6’*0, cannot be used to distinguish between I- and S-type granitoids (C)‘Neil and Chappell, 1977); and that granitoids may show a larger variation in 8”Si than do the more mafic rocks reported on here. No igneous rocks with values of 630Si > +0.3% were found in this study, in contrast with the (calculated) values of up to +l.S’% reported for granitoids by Tilles (1961a). Fewer a3’Si and no 6”O data were collected on the other suites. Four samples from the Oslo Graben igneous province (Table 1, #38-41) have $“Si = -0.4% and a range of 0.4%. Four samples from the Skaergaard intrusion were analyzed in this study (Table , X42-45); the three basalts are indistinguishable, s”‘Si = -0.5%, and are depleted in “Si relative to the Tinden Sill granophyre, 6%i = 0.0%. Tilles (1961a) reported four analyses of basalts from the Skaerto -0.7%, also with s3’Si gaard, ranging from -0.1 = -0.5%. Five samples from the Oka carbonatite complex, Canada, were analyzed (Table 1, X46-49). Three samples (546-48) have ??‘Si = -0.4%; the two samples measured with lower precision (#49a,b), a lamprophyre sill and cognate syenite ocelli (Quick et al., 1977), are identical, 830Si = +0.2%, and are possibly enriched in %i relative to the other three samples.

The mean of terrestrial mafic and ultramafic igneous rocks is z”Si = -0.4 & 0.2%. The mean of an igneous high silica population comprising a graphic granite, the Tinden Sill granophyre, and eight quartz separates from pegmatites (Table 1, #45, 55, 62-64;

10 8 6 4 2 ,-

0

-2

-3

k

(Cl Ei_cfJenlc Silica: sp~g~s~~p_h~tofrths,cl!n_!!qms.

j u-‘diatamplonktonj

-2 = 8 6

1q

dlotamite

53 ‘a

cttsrt phytoliths

--diatomifes,cherts

:

4

I Dl Clay~enic

und dtaqenel~~g~~~~ b,.,

shales ond sandstones 6 4 2 0

-3

2

.!

+I

t?

_I_ _“,.“.-““ll’l

. :,

FIG. I. Histogram 0f6~‘Si forvarious terrestriaf sample popuiations, Colored corners indicate published data source: 0 Tilles, 196la; U Yeh and Epstein, I978a,b q Clayton et al., l978b, Means are indicate<1 by arrows with a crossbar spanning + I CT: .f,, terrestrial igneous mafic and ultramaIic rocks (,V I ,i 21. lunar igneous mafic and ultramafic rocks (N = 16, calculated from data in Epstem and T‘ayl~i 1970b, 197t. 1972; Taylor and Epstein, 1973a,bJ., f3, meteorites, excluding Ailende incXusions (;V -. ?‘r. calculated from data in Yeh and Epstein. 19788; Clayton ez a!., 1978b); & chalcedony iN ~- 5 j: .k., opafine sinters (N = I1 1; &,, dissofved silicon {i”i = 5); & sponges [iV = 6); X8, clay minerafs f;Y Q~ excluding #70, 79. 80). Squares in fa) have initials within to designate sample suite: M, Mariana Island Arc; M, Hackberry Mountain: L, Lachlan Fold Belt; 0. Oslo Grahen; S, Skaergaard; C, f)ka carbonatitc

Yeh and Epstein, 1978a) is $?3i = 0.0 ri- 0.1 %a,and is not distinguishable from the mean of Hackberry ~o~ota~~

andesites

and da&es.

Ali available G3?Si data on igneous minerals, excluding those of Tilles (196ia,b) are plotted in Fig. 2. As deduced by Epstein and Taylor (1970a), the sequence of ?Si-enrichment for coexisting minerals is (pyroxene, biotite, hornblende) < (plagioclase, Kfeldspar) < (quartz, cristobaiite). Tilles (1961 a,b) reported a different ~Si-enrichment sequence for coexisting phases from Sierra Nevada granitoids and the Rose Quartz pegmatite: biotite e quartz < plagioclase 4 K-feldspar. The discrepancies between the data set of Tilles and the one presented here are not

specifically resolved by reanalysis of his samples, but it is suggested that both the discrepancies and the problem tie in the feidsgar analyses of T&s t i 95 I a.b), which in&de a 3.36 range of variation k3r K-f&spar (6”Si I^ -0.15 to +3.17%‘00)and a Z..i%e range for plagioclase (630Si = +0.5 to +2.8%0), which have not been Gunfirmed by any later workers, All data on igneous mineral separates from granitoids PCported by Tiltes ( 1961a,b) were excluded from Figs. I and 2. Consideration of the data in Figs. 1a ad 2, and of the means of the various suites, indicates that a positive correlation of 3oSi-enrichment with increasing silicon content exists in both rwks and minerals, as predicted by Grant ( 1954) on thcoreticai

ISOTOPES

OF Si

1453

grounds. Similar correlations of ‘80-enrichment with increasing silicon content have been demonstrated by Taylor and Epstein (1962b), Taylor (1968) and Matsuhisa (1979). Analysis of the simultaneously gathered 6”Si and 6”O data on lunar mineral separates (sources listed in the caption of Fig. 2) yields the following means: cristobalite, z3’Si = +0.3%0, 8’*0 = +7.0%0; plagioclase, s3’Si = -0.5%0, 8’*0 = +6.1%0; and pyroxene, z3’Si = -0.24%0, $“O = +5.8%0. Differences between the means of lunar plagioclase (abbreviated F), pyroxene (abbreviated P), and cristobalite (abbreviated Q) are: &3’s& = +0.19%0, ii”O,_, = +0.3%0; & _ 3oSio-p = +0.52%0, L\‘soq_p = + 1.2%0; A30Si QF = +0.33%0, &‘sOq_F = +1.2%0. These and similar 6’*0 data have been interpreted to indicate an approach to equilibrium at T = 1120-1200°C (Taylor and Epstein, 1970b; Onuma et al., 1970; Mayeda et al., 1975). This indication of oxygen isotopic equilibrium, coupled with correlation of 30Si-enrichments with “O-enrichments in coexisting minerals, suggests that the calculated A3’Si values for mineral pairs can be used to approximate equilibrium fractionation factors, cu”Si, at T = 1150°C. Data on comparable terrestrial coexisting minerals that could be used to help calibrate the variation of cu3’Si with 1/T2 are limited (Table 1, #59; Taylor and Epstein, 1970a; Yeh and Epstein, 1978a) but suggest that cy3’Si for mineral pairs shows little variation over the magmatic temperature range. The ratios of lunar values of 6’80/&30Si indicate that the magnitude of silicon isotopic fractionations between coexisting silicates will be two to four times smaller than those of oxygen. Consideration of the data on isolated samples in Table 1, the variations in 630Si exhibited within and between igneous suites, and the small values and limited temperature dependence of a3’Si for igneous minerals suggests that normal igneous processes, such as crystal fractionation and vapor phase separation, do not and cannot produce large variations in magmatic 630Si values. For the purposes of comparison, a3’Si + 1 u of terrestrial mafic and ultramafic rocks, lunar igneous rocks, and meteoritic silicon (excluding Allende inclusions) are indicated in Fig. la. Yeh and Epstein (1978a) have pointed out that lunar and meteoritic means are separable with some degree of confidence, and Epstein and Taylor (1970b) concluded that the isotopic composition of lunar silicon is identical with that of terrestrial silicon. Sinters

and dissolved

silicon

Siliceous sinters are accumulations of opaline silica deposited on the aprons of hot springs by venting mineralized waters. The debauching waters are generally saturated with respect to quartz but undersaturated with respect to amorphous silica; precipitation of opal occurs during cooling and evaporation

0

quartz or cristobalite

0

plogioclase

CD K-feldspar A

pyroxene

a

olivine

v

hornblende

3

muscovite

rLUMI

$@J”

igneous mincrolr

=

I*

-i

0

+I

S 3oSi (%o) FIG. 2. 6?Si of igneous minerals. Terrestrial data from this study, Tilles (1961a), Epstein and Taylor (1970a), Yeh and Epstein (1978a). Lunar data from Epstein and Taylor (1970a, 1971, 1972), Taylor and Epstein (1973b). Means of lunar minerals indicated by heavy vertical bars.

(Fournier and Rowe, 1966). The opal is deposited slowly and represents the dumping of only a portion of the silicon in solution. As recognized by Reynolds and Verhoogen (1953) opaline sinters are, in general, quite light. There is a considerable range of values (Table 1, #82-91) -3.1 to +0.5%0, with z3’Si = -1.4 + l.O%o. The dissolved silicon samples (Table 1, #96- 100, abbreviated Si(aq)) were obtained by evaporating to dryness waters collected from hot springs at Yellowstone, Wyoming and Mammoth, California. Temperatures calculated from dissolved silicon content range from 75°C to 270°C (R. Fournier, written comm., 1979). The samples analyzed show a 0.5%0 spread, with a3’Si = +0.2 +- 0.2%0. There is no correlation between b3’Si and the amount of dissolved silicon and, by extension, no correlation between a3’Si and the temperature at which the silicon content was equilibrated with quartz (Fournier and Rowe, 1966). Comparison of zsoSi of dissolved silicon with either the Hackberry Mountain silicic volcanics, $“Si = -0.1 -t 0.1 b, or the igneous high silica suite, as’Si = 0.0 t O.lL, indicates that the solution of silicon from plausible source rocks at T > 75°C probably does not measurably fractionate silicon isotopes. If there is a fractionation, dissolved silicon is probably enriched in “Si. However, comparison of the opaline sinter mean, z3’Si = -1.4 + l.OL, with that of dissolved silicon shows that a large fractionation must occur during the precipitation of opal from cooling and evaporating hot springs waters. A maximum value of ASioPAL-sicaq) = - 3.5%0 is estimated from the

1454

C. B. DOUTHITT

difference between the extreme analyses of opaline sinter, d?Si = -3.1% (#88), and dissolved silicon, s30Si = $0.4960 (#LOO). Opal spontaneously transforms into thermodynamically more stable forms of silica; the rate of this transformation is increased greatly by increasing temperature (Fournier and Rowe, 1966). Transformation of opal into quartz involves solution and redeposition during which oxygen isotopic equilibrium between silicates and water is approached (Murata et al., 1977; Kastner et al., 1977). At Steamboat Springs, Nevada, the surface sinters are negative, @OSi= -0.9%0 (Table 1, #91); chalcedonic sinter exposed at the surface is negative but heavier, d”Si - -0.3% (Table 1, #92,93); ~halcedonyencountered in drill hole GS-5 is positive (Table 1, #95a,b,c), with 630Si ranging from $0.1 to +0.9%0. Steamboat Springs chalcedony has a3’Si = +0.2 * 0.5%0. The similarity between the isotopic composition of chalcedony and hot springs dissolved silicon, x30Si = -to.2 1 0.2%+ together with oxygen isotopic evidence for equilibration (J. O’Neil, unpub. data) suggests that silicon isotopic equilibrium has been established, with a fractionation factor smaller than experimental precision. Biogenic silicon Silicon, while present in trace amounts in almost all organisms, is, in the form of opal, an essential structural component in several diverse groups of organisms and a non-essential but high percentage dry weight component of many taxa of vascular land plants. Biogenic silicon, plotted in Fig. ic, shows the widest range of S30Sivalues of any group of samples analyzed, and indeed, contributes the extrema to the currently established terrestrial spreads: -3.40/w, from a marine sponge, and +2.8%0, from a terrestrial plant. While quantitative data are few (Lovering and Engel, 1967), it is clear that plants, especially graminoids, remove vast quantities of silicon from soils. This silicon is precipitated intracellularly in the form of opaline phytoliths (literally, “plant stones”), which are cycled into soils, sediments, and aerosols. Phytoliths separated from two quite different taxa of vascular land plants, B~mbusa sp. and E~~i~~~u~ spp. (Tabfe I, #104- 107), collected from four iocalities in southern California, show a wide spread of values, ranging from s3’Si = -1.4 to +2.80/w. Horsetail (Equisetum spp.) phytoliths show variable but consistently positive 630Sivalues: +0.8. f I .‘7,and i-2.8?&. In contrast, bamboo ph~oliths are *‘Si-enriched, with S3?Si = -1.4%. Sponge spicules, like opaline sinters, are consistently light. Six samples (Table 1, #lOl-103; Yeh and Epstein, 1978a) have a range of @‘Si values, from -3.4 to -0.9%0, with $‘OSi = -2.3 + 0.9%~. These data confirm the earlier findings of Tilles ( 196 1a), who reported similar values for two sponges, a30Si = -1.3 and -2.2L. Sponges are immobile,

grow slowly, and are likely to convert only a small fraction of the dissolved silicon load in the ambient water column into skeletal opal. As such, sponges may be the best place to look for distinguishable expression of any fractionation between dissolved silicon and biogenic opal. The isotopic composition of marine dissolved silicon has not been measured but, for the sake of argument, is assumed to be between -0.4460 (the presumed value of silicon input at mid-ocean ridges) and +0.2%0 (the mean of hot springs dissolved silicon). The maximum value derived for ASiorAL_sicaq)is either ---Xl or -. 3 7% (de. pending on the assumed value for seawater silicon1 which is comparable to the earlier derived --3.4%0 maximum difference between sinter opal and dissolved silicon. Since both the abstraction of silicon from ocean water and its subsequent prectpitation as skeletal opal are mediated by biochemical reactions, the existence of “vital effects” and potentially species specific values of ASioPAL_.S,(aq) cannot be discounted. This ~ssibility was not concIusively tested in this study. While two sponges of the same genus (Table 1, #IOla,b) do have indistinguishable values of a3’Si, they were collected from the same general location. All other taxa of sponges have distinct @OSi values, but no two are from the same locatton. Opal-precipitating organisms are &aimed to have been the major route for removal of silicon from ocean water since the late Precambrian or early Paleozoic (Lowenstam, 1974). The planktonic organisms that secrete opaline tests (diatoms, sihcoflagellates, ebridia, radiolaria) certainly dominate the silicon budget of near surface ocean waters. where the ~n~entration of dissolved silicon is generally a limiting growth factor. This raises the strong posstbility that any fractionation between dissolved silicon and biogenic opal may be masked by the conversion of variable but significant proportions of the available pool of dissolved silicon into opaline silicon. A value of S3?Si = i-0.8%+ has been reported for diatom plankton (Tilles, 196la). Four analyses of diatomites (Table I, #lOS; Yeh and Epstein, 1978a; Tilles, 1961a) reveal a 2.6%a range of values, from 6’“Si = -0.9 to +1.7%0. Assuming that there is httle occasion for fractionation of silicon isotopes during the accumulation of diatom tests and subsequent compaction into diatomites, this 2.6%0 range of variation reflects variability inherited from the diatom precursors. Thirty four analyses of (presumably biogenie) cherts (Table 1, #109, 110; Yeh and Epstein, t978a,b; Tiles, 1961 a) yield a 3. I %Orange, from @‘Si = - 1.4 to + 1.7%0. Abiogenic chert from Lake Magadi (O’Neil and Hayes, 19731, with GwSi .I- O.O%o (Clayton et al., 1978b), sits right in the middle of the biogenic chert spread. Clay minerals. authigenic and diagenetic quartz. sedimentary rocks The @?Zi data on clay minerals, non-igneous quartz, and sedimentary rocks are plotted in Fig. id.

ISOTOPES OF Si

All the clay minerals and most of the quartz samples have large positive 6”O values, indicating low temperature isotopic exchange with some portion of the hydrosphere. Three orthoquartzite samples (Table 1, #66a,b,c) from a single formation (R. Powell, pers. comm., 1980) have very uniform @‘Si and &I80 values (-0.2 and + 10.9%0, respectively), suggesting that sedimentary processes have resulted in the homogenization of both silicon and oxygen in large volumes of rock. These rocks have been subjected to high grade metamorphism (Powell, 198 1 ), which may have enhanced primary sedimentation-related homogenization. The St. Peter sandstone has d”Si and 6’*0 values of 0.0 and + 11.O%O,respectively (Yeh and Epstein, 1978a; Clayton and Mayeda, 1963), and the Potsdam sandstone has @OSiand 6180 values of f0.3 and + 15.4%0, respectively (Tilles, 1961a; Clayton and Mayeda, 1963). Enrichments in I8O, whether primary or secondary in origin, seem to be paralleled by enrichments in “Si. The small spread of sandstone &“Si reflects both the limited variability of 630Si and the predominance of 630Si values of =O.O%o for quartz in source materials (cf- Fig. 2). Two analyses of shales, &?Si = +0.3 and -0.2%, were reported by Tilles (1961a). Eight analyses of fine grained diagenetic quartz from shales show a range of values from 630Si = -0.2 to +0.3%0 (Table 1, #67; Yeh and Epstein, 1978b). The samples analyzed by Yeh and Epstein (1978b) have been interpreted as having formed during diagenesis via redistribution of silicon liberated during the formation of illite (Hower et al., 1976) at temperatures in excess of 50°C (Yeh and Savin, 1977). Textural evidence and 6180 data (Yeh and Savin, 1977) strongly suggest that this quartz was in oxygen isotopic equilibrium with the pore fluid from which it precipitated. These S3@Sidata strongly suggest that discernible silicon isotopic fractionation does not accompany the equilibrium precipitation of quartz from solution at temperatures at least down to 50°C. Sample 852 is interpreted as having formed from silicon liberated during the dissolution of sponge spicules (D. Sibley, pers. comm., 1978). If this interpretation is correct, then 630Si cannot be used to discriminate between silicon derived from sponges (biologically processed silicon) and silicon liberated during the formation of illite (inorganically derived silicon). The diagenetic quartz d-“Si values fall in the middle of the spread of chert 630Si values (Fig. lc) and, despite a suggestion to the contrary (Yeh and Epstein, 1978b), 630Si data clearly cannot be used to distinguish between the two populations. Clay minerals (Table 1, #69-S 1) have a distribution of A30Sivalues similar to that of igneous rocks, with BsoSi = -0.5 f 0.4%0 (excluding the three most highly fractionated samples). The most fractionated samples include three kaolinites (#78-80) and a garnierite (#70). Two of the kaolinite samples are light, with 630Si = -1.0 and -2.3y~, while the third is heavy, ?j3’Si = + 1.9%0. Of the two light kaolinite

1455

samples, the more poorly ordered (#79) is the more fractionated. No X-ray diffraction information on ordering was taken on the heavy kaolinite. The garnierite, S30Si = -2.1%0, is from the same nickel laterite as a sepiolite (#69), a3’Si = -0.4%~ It is probable that the garnierite is poorly cystalline; it is certainly not as well ordered as the sepiolite. All highly fractionated clay samples formed during the extreme leaching and alteration of igneous protoliths that presumably had a3’Si values falling within the range shown in Fig. la. The clay data indicate that, while little or no fractionation of silicon isotopes accompanies most instances of clay formation, fractionation can occur under as yet poorly understood conditions. Herkimer ‘“diamond” (Table 1, #68) is authigenic, doubly terminated quartz precipitated from groundwater at low tem~ratures. Like the Long Valley kaolinite (#80), it shows significant 3oSi-enrichment, with 630Si = -l-1.4%0. ORIGIN OF THE NON-MAGMATIC VARIATIONS OF &Si Based largely on the similarities between the means and spreads of 630Si of dissolved silicon, chalcedony, and diagenetic quartz, it is concluded that ~30SioTz-s,~a4~ LI 1.OOOO. Equilibrium between silicates and solution, with respect to both the partitioning of silicon isotopes and the concentration of dissolved silicon, is established at temperatures at least as low as 50°C. To explain the existence of the 28Si-enriched populations of sponge spicules and opaline sinters, a kinetic fractionation that accompanies the low temperature precipitation of opal from solution is postulate to exist. This fractionation, based on estimates of maximum values of A30SiopAL_si(aq), is estimated to have a value(s) of a30SiopAL_si(a4) between 0.9963 and 0.9969. The clay data suggest that fractionation of similar magnitude may also occur during the formation of poorly ordered phyllosilicates. The fractionation is inferred to be kinetic because it is only developed in X-ray amorphous or poorly ordered materials at earth surface temperatures and because it is quite different than the equilibrium fractionation factor, a”Si oTz_sitaq), inferred for only slightly higher temperatures. Enrichment of opaline silica in **Si is expected from a kinetic isotope effect, since the lighter isotope has a greater readiness to react (Goldstein, 1966). The existence of such a kinetic fractionation is presumably masked in all systems except ones that are open with respect to the supply of dissolved silicon (i.e., the number of atoms of Si dissolved in solution ti number of atoms of Si removed by precipitation} by further precipitation from the dissolved silicon reservoir. This is illustrated in Fig. 3 for the Rayleigh process, in which solids once precipitated are removed from further interaction with the dissolved silicon. The open system

1456

C H. DOUTHITT

and, as such, cannot be confidently Interpreted, but may indicate the existence of “Si-enriched ground. waters. CONCLUSIONS

-40L

I

/

02

04

06

08

i \7

tractmn condensed

FIG. 3. Schematic illustration of Rayleigh process in precipitation of opal from dissolved silicon. The form of the Rayleigh equation used was: R = f’“~“, where R = (‘“Si/

**Si)rsr,duals,(aq)/(30Si/28Si),nlt,al~,~aq~, f = fraction of Si(aq) precipitated, cy = solid/dissolved tor. (30Si/28Si)~~l/(30Si/28Si)s,~aq~ the example illustrated; b?Si(aq)

silicon fractionation facassumed to be 0.9965 for = 0.0%.

condition of above refers to f = 0 in Fig. 3. Ciiven the parameters used in calculating the curves of Fig. 3, ?j30Sicaq) = O.O%oand a30SiopAL.si(aq) = 0.9965, the Rayleigh process is capable of yielding a range of values for 6”Si greater than that so far observed. The variations of 630Si found in sponges and opaline sinters are ascribed to Rayleigh precipitation of opal under conditions of limited depletion of the dissolved silicon reservoir. The larger range of 630Si values exhibited by diatomites is also attributed to Rayleigh precipitation of opal, but with much more variable to extreme depletion of the dissolved silicon reservoir. The spreads of a3’Si values of phytoliths (- 1.4 to +2.8%0) and of clay minerals (-2.3 to +1.8%0) are similar to that of diatomites and may be similarly explained, but not enough data have been collected to discuss systematics. The 3.1 %Orange of 630Si values exhibited by cherts (-1.4 to +1.7%0) was probably established during the biological uptake of silicon from seawater. During the transformation of opal into quartz, equilibration of silicate silicon with a pool of dissolved silicon may alter the primary opaline h3’Si values. G30SicHERTdepends on whether the isotopic composition of dissolved silicon is internally or externally buffered, and on the isotopic composition of the buffering pools. Unevaluated factors that would complicate this scenario include species specific fractionations and fractionations accompanying the low temperature solution of silicon from silicates. The value of + 1.4%~ for an authigenic quartz crystal is an isolated analysis

Silicon isotope ratios have potential, albeit very limited, to contribute useful information to a variety of geological and geochemical studies. The mean oi terrestrial mafic and ultramafic igneous rocks is 8"Si = -0.4%~ Silicon isotope ratios show little variation in igneous rocks and minerals; hence, they have little import for studies of igneous rocks. In a general way s3’Si increases with the silicon contents of both igneous minerals and rocks. bsoSi values of dissolved silicon are not weit pinnen down. Marine dissolved silicon is assumed to he generally between -0.4 and tO.2%0. Substanttal ‘“Sienrichment is postulated to occur locally as a resulr of the biogenic precipitation of opal. Neither silicon dissolved in water above 75°C nor quartz deposited from solution above 50°C are distinguishably frae tionated relative to plausible silicate or dissolved siticon sources. This is inferred to be the equilibrium situation, and a3’Siorz-si(aq) 2 I .OOOO. Relatively large fractionations occur in opaline sinters; biogenic opal. including sponge spicules, diatoms, diatomites, cherts, and phytoliths; clay minerals; and authigenic quartL. Opaline sinter and sponge spicules both show mod, erately variable but consistently negative 6’“Si valuea that are interpreted to be indicative of a kinetic fraL tionation attendant upon precipitation of opal from dissolved silicon. This fractionation appears to be dependent on the existence of a substantial degree of crystallographic disorder, and is estimated to have a value(s) between 0.9963 and 0.9969; combined with Rayleigh precipitation of silicates from solution, it is capable of explaining the currently known non magmatic range of variation of fi3’Si. The transformation of opaline silica to quartz is accompanied by equilibration with a fluid phase and its dissolved constituents, potentially causing a loss or “blurring” ot original isotopic ratios. Since the total conversion of silicate stlicon into SiF4 occurs during the reaction of silicates with both F2 and BrF, and, since the complete recovery of SiF, is possible during routine recovery of 0, with fittle modification of conventional equipment and little extra work, SiF., should be routinely collected. Sib, yields and SiF4/02 yield ratios provide valuable quality controls on analyses of both mineral separates and “whole rocks”. Acknowledgments-This research was supported by Na, tional Science Foundation grant #EAR-7816873 awarded to S. Epstein. I would like to thank S. Epstein for support and R. Becker for critical reading of the manuscript and help with mass spectrometry; V. Nenow and M. Carr for mass spectrometer and extraction line maintenance; B. Chappell, T. Dixon, R. Fournier, M. Furst, .J. Goss. Jim Hoffman, S. Jacobsen, R. Lewis, J. O’Neil, Helen of Tech. R. Powell. J. Quick, R. P. Sharp, D. Sibley, N. Wells, and

ISOTOPES D. White for free samples; R. DuPont for drafting; Solomon for snickering at the right time.

G. C.

REFERENCES Allenby R. J. (1954) Determination of the isotopic ratios of silicon in rocks. Geochim. Cosmochim. Acta 5,40-48. Becker R. and Epstein S. (1981) Silicon and oxygen isotopes in some selected Allende inclusions (abstract). In Lunar and Pianetarv Science XII. Vol. I. DD. 56-58. Lunar and Planetary Institute, Houston. -Clayton R. N. and Mayeda T. K. (1963) The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochim. Cosmochim. Acta 27, 43-52. Clayton R. N., Mayeda T. K. and Epstein S. (1978a) Isotopic fractionation of silicon in Allende inclusions (abstract), In Lunar and Planetary Science IX, Vol. I, pp. 186- 188. Lunar and Planetary Science Institute, Houston. Clayton R. N., Mayeda T. K. and Epstein S. (1978b) Isotopic fractionation of silicon in Allende inclusions. Proc. 9th Lunar Plane?. Sci. Con$, Houston. Vol. 1, pp. 12671278. Pergamon Press. Craig H. (1957) Isotopic standards for carbon and oxygen and correction factors for mass spectrometric analysis of carbon dioxide. Geochim. Cosmochim. Acta 12,133- 140. Dixon T. (1979) Volcanic seamounts in the southern Mariana Island Arc (abstract). EOS 60, 969. Dixon T. and Batiza R. ( 1979) Petrology and chemistry of recent lavas in the northern Marianas: implications for the origin of island arc basalts. Contrib. Mineral. Petrol. 70, 167-181. Epstein S. and Taylor H. P., Jr. (1970a) ‘*O/‘6O, %i/**Si, D/H and “C/‘*C studies of lunar rocks and minerals. Science 167, 533-535. Epstein S. and Taylor H. P., Jr. (1970b) The concentration and isotopic composition of hydrogen, carbon and silicon in Apollo 11 lunar rocks and minerals. Proc. Apollo I I Lunar Sci. ConJ, Houston. Vol. 2, pp. 1085-1096. Pergamon Press. Epstein S. and Taylor H. P., Jr. (1971) 0’8/0’6, Si”‘/Si2’, D/H, and C”/Cr’ ratios in lunar samples. Proc. 2nd Lunar Sci. Conf, Houston. Vol. 2, pp. 142 I- 144 1. M.I.T. Press. Epstein S. and Taylor H. P., Jr. (1972) 0’8/0’6, Si”/S?*, C”/C12 and D/H studies of Apollo 14 and 15 samples. Proc. 3rd Lunar Sci. Conj, Houston. Vol. 2, pp. 14291454. Pergamon Press. Epstein S. and Taylor H. P., Jr. (1973) ‘*O/‘6O, 3oSi/z8Si, ‘%/‘% D/H and hydrogen and carbon concentration data on’Apollo 11 soils. EOS 54, 585-586. Epstein S. and Taylor H. P., Jr. (1974) Oxygen, silicon, carbon and hydrogen isotope fractionation processes in lunar surface materials (abstract). Lunar Science V, Vol. 1, pp. 212-214. The Lunar Science Institute, Houston. Epstein S. and Taylor H. P., Jr. (1975) Investigation of the carbon, hydrogen, oxygen, and silicon isotope and concentration relationships on the grain surfaces of a variety of lunar soils and in some Apollo 15 and 16 core samples, Proc. 6th Lunar Sci. Conj, Houston. Vol. 2, pp. 14291454. Pergamon Press. Epstein S. and Yeh H.-W. (1977) The 6’*0,6”0, a3”Si and 6% of oxygen and silicon in stony meteorites and Allende inclusions (abstract). Lunar and Planetary Science VIII, Vol. 1, pp. 287-289. The Lunar and Planetary Science Institute, Houston. Fournier R. 0. and Rowe J. J. (1966) Estimation of underground temperature from the silicon content of water from hot springs and wet steam wells. Amer. J. Sci. 264, 685-697.

OF Si

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Goldstein M. J. (1966) Kinetic isotope effects and organic chemical mechanisms. Science 154, 16 16- 162 1. Grant F. S. (1954) The geological significance of variations in the abundances of the isotopes of silicon in rocks. Geochim. Cosmochim. Acta 5, 225-242. Hower J., Eslinger E., Hower M. E. and Perry E. A. (1976) Mechanism of burial metamorphism of argillaceous sediment: 1. Mineralogical evidence. Bull. Geol. Sot. Amer. 87, 725-737. Jacobsen S. B. (1980) Study of crust and mantle differentiation processes from variations in Nd, Sr, and Pb isotopes. Ph.D. dissertation, California Inst. Technology. Kastner M., Keene J. B. and Gieskes J. M. (1977) Diagenesis of siliceous oozes - 1. Chemical controls on the rate of opal-A to opal-CT transformation-an experimental study. Geochim. Cosmochim. Acia 41, 10411059. Lovering T. S. and Engel C. (1967) Translocation of silicon and other elements from rock into Equisetum and three grasses. U.S. Geol. Surv. Prof Paper 594-B, 16 p. Lowenstam H. A. (1974) Impact of life on chemical and physical processes. In The Sea (ed. E. Goldberg), Vol. 5. .on. 715-796. John Wilev and Sons. Matsuhisa Y. (1979) Oxygen isotopic compositions of volcanic rocks from the East Japan island arcs and their bearing on petrogenesis. J. Volcanol. Geothermal. Res. 5, 271-296. Matsuhisa Y., Goldsmith J. and Clayton R. N. (1978) Mechanisms of hydrothermal crystallization of quartz at 250°C and 15 kbar. Geochim. Cosmochim. Acta 42,173182. Mayeda T. K., Shearer J. and Clayton R. N. (1975) Oxygen isotope fractionation in Apollo 17 rocks. Proc. 6th Lunar Sci. Conf, Houston. Vol. 2, pp. 1799-1802. Pergamon Press. Murata K. J., Friedman I. and Gleason J. D. (1977) Oxygen isotope relations between diagenetic silica minerals in Monterey shale, Temblor Range, California. Amer. J. Sci. 277, 259-272. O’Neil J. R. and Chappell B. W. (1977) Oxygen and hydrogen isotope relations in the Berridale-batholith. 3. Geol. Sot. London 133. 559-571. O’Neil J. R. and Hayes’R. L. (1973) ‘8O/‘6O ratios in cherts associated with the saline lake deposits of east Africa. Earfh Planet. Sci. Lett. 19, 257-266. Onuma N., Clayton R. N. and Mayeda T. K. (1970) Apollo 11 rocks: oxygen isotope fractionation between minerals, and an estimate of the temperature of formation. Proc. Apollo I1 Lunar Sci. Co& Houston, Vol. 2, pp. 14291434. Pergamon Press. Powell R. E. (1981) Geology of the crystalline basement complex, eastern Transverse Ranges, southern California: constraints on regional tectonic interpretation. Ph.D. dissertation, California Inst. Technology. Quick J., Albee A., Ma M.-S., Murali A. and Schmitt R. (1977) Chemical compositions and liquid immiscibility of two silicate melts in 120 13. Proc. 8th Lunar Sci. Conj, Houston. Vol. 2, pp. 2153.-2189. Pergamon Press. Rankama K. ( 1954) Isorope Geology. Pergamon Press. Reynolds J. H. and Verhooaen J. (1953) Natural variations in the isotopic constitution of silicon. Geochim. Cosmochim. Acta 3, 224-234. Taylor H. P., Jr. (1968) The oxygen isotope geochemistry of igneous rocks. Contrib. Mineral. Petrol. 19, l-71. Taylor H. P., Jr. and Epstein S. (1962a) Relationships between 0’8/0’6 ratios in coexisting minerals of igneous and metamorphic rocks, Part I: Principles and Experimental Results. Bull. Geol. Sot. Amer. 73, 46 I-480. Taylor H. P., Jr. and Epstein S. (1962b) Relationships between 0’8/0’6 ratios in coexisting minerals of igneous and metamorphic rocks, Part 2. Application to petrologic problems. Bull. Geol. Sot. Amer. 73, 6755694.

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Taylor H. P., Jr. and Epstein S. ( 197Oa) Oxygen and silicon isotope ratios of lunar rock 12013. Earth Planet. Ski. Lert. 9, 208-210. Taylor H. P., Jr. and Epstein S. (1970b) O1*/Oib ratios ol Apollo 11 lunar rocks and minerals. Proc. Apollu I I Lunar Sci. Con$, Houston. Vol. 2, pp. I2 I3- 1226. Pergamon Press. Taylor H. P., Jr. and Epstein S. (1973a)Oxygen and silicon isotope ratios of the Luna 20 soil. Geochim. Cosmochim Acta 37, 1107-l 109.

Taylor H. P., Jr. and Epstein S. ( 1973b) O’“/O” and Si”“/ Si2* studies of some ApoUo 15, 16, and 17 samples. Pr[w. 4th Lunar Sci. Conj, Houston. Vol. 2, pp. 16% 1679 Pergamon Press. Tilles D. (1961a) Natural variations in isotopic abundances of silicon. J. Geophys. Rex 66, 3003.-3014

Tilles D. (1961 b) Variations of silicon isotope rlilios in _: pegmatite. J. Geophys. Rex 66, 3015 3020. Wells N.. Childs C. W. and Dawnes C. J. f 1977) Silica Springs, Tongariro National Park, New %ealarrtl .~n;~l. yses of the spring surface water and characterisation oj the alumina-silicate deposit. Geochim. Co~m~him. .~ivf~ 41, 1497-1506.

Yeh H.-W. and Epstein S. (1978a) “Si/*“Si and ‘*Si/:%t of meteorites and AIlende inclusions Iabstract ). Lunar and Planet. Sci. IX. Vol. 2, pp. 1289,- 1291. II mar and Planetary Science Institute, H&on. Yeh H.-W. and Eostein S. (1978b) ““Sii”“Si ratios in fine, grained quartz ‘separated. from shale; which were sub jetted to deep burial diagenesis. Geoi. Soi. dmcs lihsia Prog. IO, 520.

Yeh H.-W. and Savin S. (1977) Mechanism of bur& metamorphism of argillaceous sediments: 3. O-Isotope cvi dence. Bull. Gent. Sot. Amer. 88, I32 I 1?W