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The extent of oxygen isotope exchange between clay minerals and sea water
Grochlmsa
et Coamochim~ca
Acta.
1976. Vol
4tl. pp
Press.
743 to 74X. Pergamon
Prmtcd
in Great
Br~tam
The extent of oxygen isotope exchange between clay minerals and sea water* HSUEH-WEN YE& and SAMUELM. SAVIN Department of Earth Sciences, Case Western Reserve University, Cleveland, Ohio 44106, U.S.A. (Received 2 June 1975; accepted in revised form 8 December 1975)
Abstract-The extent of oxygen isotopic exchange between detrital clay minerals and sea water was investigated by analyzing 0’8/016 ratios of separated fine-grained size fractions of deep-sea sediments from three North Pacific ocean cores. Isotopic results were interpreted according to models based on the assumption that the extent of isotopic exchange should increase with decreasing particle size and increasing time of exchange between the sediment and sea water. The data indicate that information concerning the provenance and mode of formation of detrital clay minerals can be obtained from the 018/0’6 ratios of the coarser-than-O.1 Atrn fraction of deep-sea sediments younger than several million years and the finer-than-O.1 pm fraction of deep-sea sediments younger than several tens of thousands of years. Furthermore, if the extent of chemical reaction between detrital clays and sea water is similar to the extent of oxygen isotopic exchange, such reaction may be important in regulating the chemistry of sea water.
INTRODUCTION OXYGENisotopic equilibrium between a clay mineral and the ambient water is generally approached when the mineral crystallizes (SAVIN and EPSTEIN, 1970 a,b; LAWRENCEand TAYLOR, 1971; SHEPPARD et al., 1969). The isotopic composition of the clay is thus a function of the temperature of formation and of the O’S/O’6 ratio of the ambient water. As long as the oxygen isotopic composition acquired by a clay mineral at the time of its formation remains unaltered, it can provide information concerning the nature of the environment in which the mineral formed. There is considerable evidence of a qualitative nature that post-formational isotopic exchange between clay minerals and water is slow at sedimentary temperatures, except when the clays undergo chemical or mineralogical alteration. SAVIN and EPSTEIN (1970b) concluded that there had been no large-scale oxygen isotopic exchange between sea water and the detrital clays of several Recent and Pleistocene ocean core samples they analyzed. LAWRENCE and TAYLOR (1972) concluded that only insignificant post-formational isotopic exchange of the clay minerals of some Quaternary soils had occurred. SHEPPARDet al. (1969) found no evidence for low temperature isotopic exchange of the hydrothermal clay minerals of some Tertiary and Mesozoic porphyry copper deposits. The sluggishness of oxygen isotopic exchange between clay and water is further documented by the laboratory experiments of O’NEIL and * Contribution No. 115, Department of Earth Sciences, Case Western Reserve University. t Present address: Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena. California 91 109. U.S.A. 743
KHARAKA (1976) who found negligible exchange between illite and water and between kaolinite and water at 100°C over times as long as 2 years. While qualitative data indicate that the rate of isotopic exchange is slow, quantitative data on the rate of exchange have been lacking. It is the purpose of the work reported here to place quantitative limits on the rate of low temperature oxygen isotopic exchange between clays and sea water. Such limits should serve to determine the maximum ages of ocean sediments for which oxygen isotope ratios can provide information concerning the environments of formation of the silicates. Additionally, the extent of oxygen isotopic exchange between clay minerals and sea water is probably similar to the extent of chemical interaction between the two phases. Thus inferences can be drawn from the isotopic data concerning the magnitude of such chemical interaction and its importance in regulating the chemical composition of sea water. ANALYTICAL TECHNIQUES Core samples were ultrasonically disaggregated in distilled water, and were subsequently subjected to the following chemical treatments: 1 N sodium acetate buffer (pH = 5) to remove carbonate; 30% H202 to remove organic matter and manganese oxides; and a sodium bicarbonatesodium citrate solution to remove iron oxides (JACKSON, 1956). These treatments have no effect on the oxygen isotopic compositions of clay minerals (YEH, 1974). The silicate residues thus obtained were separated according to particle size (equivalent spherical diameter) by centrifugation. After mineralogical analysis by X-ray diffraction, the separated fractions were isotopically analyzed. After drying in a glove box at a relative humidity near O”/”for at least 24 hr, oxygen was liberated from the samples by reaction with fluorine (TAYLORand EPSTEIN, 1962a). Duplicate extractions of oxygen were done for most samples. The averape deviation from the mean of the results of the two
Hsr I H-WI \ YI II and SA~LI~ I I M. SA\I~
734
was 113s than 0.2(YC,,, for hO”,, of all samples. Isotopic data are reported in J-notation, as deviation< 01‘the isotopic composition in parts per thousand (‘I,,,,)liom that of the SMOW standard (C‘KAI(;. 1961).
extractions
ISOTOPIC
EXCHANGE MODELS
When detrital &I! minerals of continental origin are incorporated in marine ‘;ediment\ their O’XKI’h ratios are generalI> lower than equilibrium values for clays in the marine en\ironmcnt. Thi\ reflcctz: (I ) the greater 0’“‘O’” ratio of sea water than of most fresh waters: and (2) the colder temperatures of the ocean bottom than of most continental en\ironmcnt\. Isotopic exchange of the detritus toward an equilibrium value can be monitored and expressed using the rclationship:
dctcrmincd cxperimentall! at higher temperatures: comparison of the isotopic compositions of clay minerals which formed at known temperatures with the isotopic compositions of the waters in which they formed: and estimation based on the crystal chemistry of minerals and/or analogy with the isotopic fractionations of similar minerals (TAKOK and EPSTLIN. 1962b; SAVIN. 1967: O’NFIL and TAYI.OK. 1967). Equilibrium fractionation lktors, estimated for a tetnperature of I C. are listed in Table I. The values should be considered approximate. Errors of I T,,,, in the values for smectite and illite. and 3- 4’:,,, for chlorite and kaolinite will not greatly affect the conclusions drawn. a. .S/~~c,c,rirc, ~t‘~l/c’r.The oxygen isotopic fraction bctwecn \mectitc and water was estimated from the anal>Gs of :I Gngle sample of authigenic smcctite taken from core WAHINE I I P (8 16’N. I.53 01’W) supplied h! the Scripp\ Institute of Oceanography. The fractionation factor cstimatcd from the data for this sample is approximately 3.5”,,,, greater than estimated h! S/t\ IU :tnci EP\I I.IV ( I97Oa) from the analyses of t\vo authigcnic smcctitc\. Howc\er. Satin and Epstein did not treat their \mectites to rcmo\‘e oxide mineral\ mid organic maltcr. b. Illitem I~U~CI’. The illilc water fractionation factor at I c‘ wm determined by extrapolating (on a graph of 1000 In 2 vs 1,T’) the illiteewater fractionation estimated for higher temperatures (YEH and SAVIK. in press: ESL.IU,I I< and SAYI\. 1973). c. k’troli~lirv ~v~lrvr. This fractionation factor was estimated by adjusting the value for 17 C given by SA\‘IN and EI?TEI~%(197Oa) to the lower temperature conditions of the ocean bottom. d Chlorirc~ -,c'clter. The chlorite -water fractionation factor was estimated b> combining the fractionation factors for various oxygen containing bonds, as estimated by SARIS (1967).
no,:,:,, ,, is the isotopic composition of the detritus prior to entering the marine environment. CKI,l,;,,,,,,,~,,, iy the value the detritus would ha\,c if it were in cquilibrium with sea water at ocean-bottom temperature. Valid application of the above expression dcpenda on the accurate estimation of 80,t,~,, ,, and &I~‘,;,,,,,,,,,j,,,. 21s discussed below. It should be kept in mind that ;I change in the isotopic composition of a clah mineral. in the absence of evidence for chemical or mineralogical alteration, ma> reflect either of two different mechanisms. A change in isotopic composition is most readily interpreted as reflecting isotopic exchange. Howe\cr, the isotopic data b! themselvzs do not rule out the possibility that the shift in isotopic composition reflects the formation of additional. authigenic material in the marine environment. This material would ha\c to be mineralogicalI> similar to the detritus. and might perhaps form bq the reconstitution of amorphous detrital material as proposed b! MAG+NLII. and GAKKI:LS(1966). Regardless of which mechanism produced the change in isotopic composition of the cla~c. the intcrprctation of the isotopic The O’“‘0”’ ratio of a detrital mineral prior to data and the application of the results arc Gmilar. deposition on the ocean Hoor must be estimated indircctl!. The method most appropriate for this eatimation differs for different samples, and is a function PreciseI) determined oxygen isotope fractionation of sample mineralogy. Two methods applied to the factors between claq minerals and water are not acailresults of this study are described below. Xl‘~tllod .‘I w/1cw r,rinPra/ogJ 01‘ .sun1ple dors irot able. Howebcr, these fractionation factors ma! bc C/KUI~~C~ with prwtidr~ ,sizc. When the mineralogy of approximated in a number of ways, including: extrapolation to low temperatures of fractionation factors each particle-size fraction of a sample is the same, Table
I. Estimated
clay mineral-
water isotopic
Mineral
Fractionation FaCtOr*
smectice llllte Kaolinite Chlorite ____
1.03083 1.02789 1.02645 1.02115
fritctionation
factor5
at
I C
Eq”s-.---,;0181 t30.62 +21.b8 t26.24 +22.95
*
Fractionation factors = (O’X~O’~)~,.,~,(O’X~O1n~~i~~. t Thcsc values are calculated using a sea water ci0”
v;due of
--(Uo”,,,,.
isotope exchange
Oxygen
745
Table 2. Locations, ages and water depths of samples used in exchange rate studies
JYN-II-40
10 to
15
33'04'N
174°15'w
4
31,000
5530
JYN-IV-11G
33 to 39
27s42'N
175*1O'W
2.5
14.000
5750
SCAN 29P
15 to 52
33'16'N
153'4b'E
4
84,000
5857
SCAN 29P
1065 to 1116
33°16'N
153"44'E
4
2,725,OOO
585,
*
Estimated from data in
OPDYKE
and FOSTER (1970).
assumption, the 0L8/0”’ ratio of each size fraction in the uppermost layer of the core may be taken to approximate the initial O18/O16 ratio of that size fraction of the detritus.
it may be assumed that SO~~~~i~~ of each particle-size fraction was also the same. Additionally, it is logical to assume that the extent of isotopic exchange has been greatest for the finest fractions and that, when the sediments are not too old, the coarsest fraction has undergone little or no isotopic exchange. Under those conditions, the isotopic composition of the coarsest fraction can be taken to approximate that of the unexchanged continental detritus. Method B-when mineralogy of a single particle size does not vary with depth in the core. Ocean sediments are complex mixtures of many minerals, and, after the chemical treatments discussed above, generally consist of illite, smectite, mixed-layer clays, chlorite and kaolinite, as well as quartz and feldspar. Commonly the relative proportions of these minerals change as a function of particle size, and in such cases it cannot be assumed that the initial isotopic compositions of any two size fractions were the same. In some such instances, however, the mineralogy of a single particle size fraction may be quite constant throughout the length of a core, reflecting fairly constant geological conditions in the source area(s) of the detritus. Under such circumstances, it is reasonable to assume that the initial isotopic composition of each particle size has remained constant over the time of deposition of the core material. Given that
SAMPLES STUDIED Studies of the rate of isotopic exchange between detrital clays and sea water were made using samples from thrqe North Pacific cores. Data concerning the locations, ages and water depths of the samples are given in Table 2. All samples are predominantly detrital, although SCAN 29P, 1065-l 116 cm contains minor volcanic ash fragments. Depositional ages of samples were estimated from the sedimentation rates determined for the North Pacific by OPDYKE and FOSTER (1970) on the basis of paleomagnetic measurements. None of the samples in the study were from sites studied by Opdyke and Foster. and we have therefore taken averages of sedimentation rates which they calculated for samples from surrounding locations. Ages determined in this way must be considered approximate. Dr. Harvey Sachs of Case Western Reserve University examined radiolarian assemblages from the samples. He determined a biostratigraphic age of 0-400,000 yr B.P. for JYN-II-4G, 10-15 cm (compared to a sedimentation-rate age estimate of 31,000 yr), and a biostratigraphic age of 800,00@5 million yr B.P. for SCAN 29 P, 1065-t 116 cm (compared to a sedimentation-rate age estimate of 2,725,OOO yr). Although precision is not great, ages estimated by these two independent methods are not in conflict. X-ray diffractograms were made of all samples after chemical treatment and size-separation as described ab0ve.t Mineralogic data determined from the diffractograms are summarized in Table 3.
t Diffractograms are available from the authors.
Table 3. Mineralogic
SCAN SCAN SCAN SCAN
29P 29P 29P 29P
SCAN 29P SCAN 29P
SCAN 29P SCAN 29P
data for samples
used in studies of percentage
of exchange
10 to 15cm II II
1 - 0.5 0.5 - 0.2 0.2 - 0.1
7, 80 80
20 20 15
Trl 0 5
3 Tr 0
TT 0 0
+2,.36 +*,.36 +2,.36
33 to 39cm II I,
2 1 1 - 0.5 0.5 - 0.1
74 77 80
20 20 15
0 0 5
5 3 0
1 TT 0
f27.36 +2,.36 l27.36
15 to 52cm I. II II
2 1 1 - 0.5 0.5 - 0.1 so.1
70 72 50 40
13 12 0 0
10 12 50 60
5 3 0 0
2 1 0 0
c27.70 f27.80 +29.2, +30.40
1065 to 1216mo II I, I,
2 1 1 - 0.5 0.5 - 0.1 ‘0.1
70 72 50 40
13 12 0 0
10 12 50 60
5 3 0 0
2 1 0 0
+*,.,o +2,.80 +29.2, +30.40
* Illite. ** Kaolinite
and chlorite.
*** Smectite.
@ Quartz.
$ Feldspar.
# Trace.
HSI I~I-WI u YI H and SAM\ TV_ M. %\I\
746 Table 4. Isotopic
data and percentage of exchange JYN-IV-I IG
for samples
JY N-II-4G
and
L These values have been obtained using method B. * These values have been obtained by adding to the minimum values the percentage exchange determmed for sample JYN II 4G. IO I5 cm. The mineralogy does not vary with size fraction of tither JYN II-4G or JYN IV-I IG. The isotopic data for these samples are therefore amenable to treatment as outlined in method A above, for the determination of hO&,,. Mineralogy varies markedly with size fraction for both SCAN 29 P samples. However, the mineralogic compositions of corresponding size fractions of the two SCAN 29 P samples are similar. The isotopic data for these samples are therefore amenable to treatment as outlined in method 8.
ISOTOPIC
RESULTS
The percentages of exchange calculated for samples from each of the three cores are plotted against particle size in Fig. I. While SCAN 29 P. 1065~.I1 10 cm appkars to have undergone more exchange than either of the younger two samples. this might simply reflect the fact that maximum values for exchange, as described above. arc plotted for SCAN 29 P. If minimum values had been plotted an age effect would not have been nearly so apparent. A particle size etrect is. however. quite evident. This suggests that isotopic exchange at crystal surfaces probably occurs much more rapidly than does diffusion into the interiors of crystals. This is expected in the case of simple
The isotopic data and calculated percentages of isotopic exchange of samples JYN II-4G and JYN IV- 11G are shown in Table 4. Percentage of exchange generally increases with decreasing particle size and ranges to values as high as 16”,, for the 03X).1 itrn / Legend fraction of JYN II-4G. Isotopic results for the two SCAN 29 P samples are given in Table 5. The percentage of exchange was calculated, using method B, for the deeper sample onl>. since the isotopic data from the shallower sample were used as ~io,!,:~,,,, values for the different size fractions. The percentage of exchange calculated in this way is a minimum value. since the iK3,~,~K,,;,, values used were, in fact, obtained from the analysis ......... of samples to which we have assigned an age of 84,000 yr. Maximum values for the percentage exchange of SCAN 29 P. 1065 to 1116 cm can be obtained by ..s.ew adding to the minimum values the percentages of exchange obtained for the different size fractions of a3 10 20 30 .Ol 02 .05 0.1 02 Particle Size (pm) JYN II4G (assigned age 31,000 yr). As with the other two samples. the percentage of exchange of SCAN Fig. I. Per cent exchange vs particle sire for three North 29 P, 1065-I I 16 cm increases markedly with decreasPacific core samples. Values of per cent exchange for SCAN 29 P. 1065 to I I IO are maximum values (see text). ing particle size. I.
-
. .
.
.
. .
. .
.
.
Oxygen isotope exchange diffusion models. If the data are interpreted as reflecting, in part, the growth of authigenic mineral phases rather than the isotopic exchange of existing phases, it could be inferred that the growth of the authigenic phases occurs early in the post-depositional history of the sediment and primarily in the finer size fractions.
DISCUSSION Implications
for
provenance
studies
The results of this study permit qualification of the conclusion drawn by SAWN and EPSTEIN (1970b) that oxygen isotopic exchange between clay minerals and sea water is slow, and that 601’ values of detrital minerals do indeed reflect the environment of mineral formation and can be used as indicators of provenance. The data presented here suggest that over periods longer than a few million years, the 0.1-0.5 pm fraction and coarser fractions of detrital clay minerals in ocean sediments can be expected to have retained their initial oxygen isotopic compositions to the extent that their 60’* values can provide information concerning provenance. Such information can be obtained from the finer-than-O.1 pm size fraction of clay minerals in ocean sediments younger than a few tens of thousands of years. Implications
regarding
the chemistry
of sea water
Chemical equilibrium considerations led SILLEN (1961) and subsequent investigators (HOLLAND, 1965; GARRELS, 1965; etc.) to the conclusion that the interaction between silicates and sea water and the formation of authigenic silicate phases were important in buffering and regulating the chemistry of sea water. MACKENZIE and GARRELS (1966) came to a similar conclusion on the basis of mass balance calculations. Despite the prediction of the formation of authigenic clay minerals and of alteration of detrital clay minerals in the marine environment, several mineralogic and isotopic studies showed no evidence of the presence of significant amounts of authigenic material or of significant alteration of tletrital cla! minerals in mo
141
results of these studies did not preclude the presence of small amounts (up to several per cent) of authigenic or altered detrital material. The approach used in this study to estimate the percentage of isotopic exchange of separated size fractions of a sediment may be extended to obtain an estimate of the extent of isotopic exchange of the entire sediment. Since isotopic exchange involves the breaking and reforming of the oxygen-containing bonds of a mineral, it is unlikely that a great deal of isotopic exchange would occur without chemical reaction also taking place. It is therefore probable that the extent to which a sediment approaches isotopic exchange equilibrium with sea water corresponds to the extent to which detrital clay minerals have been chemically altered or authigenic clay minerals have been formed. The total percentage of isotopic exchange of a sample can be calculated from the extent of exchange of the different size fractions. Total per cent exchange
= 1 Aei,
L
where fi is the fraction of a sediment which falls in size range i and ei is the percentage of isotopic exchange of size range i. Values offi were determined only for the two SCAN 29 P samples, and were virtually identical for those two samples (> 2 pm, 30%; 2-I pm. 147;; l-O.5 pm, 14%; 0.5-0.1 pm, 37%; ~0.1 pm, 5%). In the absence of such data for the other two samples, this size distribution was assumed to approximate that of all samples. Calculated values of total percentage of isotopic exchange are given in Table 6. They indicate small but significant amounts of isotopic exchange of all sediments, with 4-9% of the detrital ocean sediments having undergone isotopic exchange in period of 104-lo6 yr. If a corresponding amount of chemical reaction occurred, the effect of this reaction on the chemistry of sea water could be considerable. For example, on the basis of material balance calculations, MACKENZIE and GARRELS (1966) proposed that the chemical composition of sea water would be maintained at steady state by the reconstitution of degraded aluminosilicates and the formation of a quantit? of authigenic cla! minerals equivalent to about 7”,, of the mass of ocean sediments.
Table 6. Total percentage of exchange of samples
Ackrruwlr~yemerlrs-Wi: thank Dr. JOHN HOWEK for providing advice and guidance throughout the study. Dr. HARVEY SACHS br his biastratigraphic determinations, Ms. IVY %3 and Ms. ?RMF, ~~EHLI for assistance in sample analysis and the Scripps ~~st~~nr~onof Oceanography for providing samples. Financiai support was provided by National Science Foundation Grants GA-4016 and GA- 16827. REFERENCES
~i~era~o~ and se~irn~n~a~~on of recent deep sea dlay in the Atlantic Ocean and adjacent seas and oceans. Bull. Geol. Sm. Amer. 76, 803-832. Ca~ra H. (1961) Standard for reporting concentrations of deuterium and oxygen 18 in natural waters. Science 133. 1833, DASCH E. J. (1969) Strontium istlfopes in rvrathcring prlibs, deep-sea sediments and sedimentary rocks. &rt+ B~SCAYE P, E. (1965)
claim. Cosmf3c&m. Aetn 33. 152I-1552. ESIJNGER E. V. and SAVKN S. M. (1973)
Mineraology and oxygen isotope geochemistry of the hydrorhermally altered rocks of the Qhaki-Broadlands, New Zealand geothermal area. Amer. 1. Sci 273. 240-267. CARRELS R. M. (1965) Silica--rote in the ht&ring of natural waters. S~imc~ 1J8. 69. ~c%.DBERGE. D. arid @GFPii'4 J. J. [1363) Ck%y-rninerai distributions in the Pacific Ocean. fn The Sea. (editor M. N. Hill), pp. 72%741, Interscience. %tL,ANI) H. D. (1965) ‘The history of occi!n water and its effect on the chemistry of the atmosphere. In SUII/?. E~alution of rhe Earth’s Armosphere. Pro?. rl’rtrl Acat/_ St% S3_ t $73-1 I83
L. (1956) Soil Chemical A~ul~!sis-.-Advanced Course, 894 pp. M. L. Jackson. LAWRENCE J. R. and TAYLOR I-i.P.. JR. (1972) Hydrogen and oxyg~s ismope sysfeematics in ~~~t~~~~~~ pro&s. Gr0&im. ~U~~~~~~. Actn %* 2377~1393. LAWRENCI-: J. R. and TavKm H. i)., JR. (1971) Deurerium irnd oxygen correlation: clay minerals and hydroxides
JACKWN
M.
in Quaternary soils compared to meteoric waters. Groilctir 36, I377- 1393. M~(‘li~~7.n: F. T. and GARKtitS R. M. (10661 Chemical mass balance between rivers and oct”iins. Amiii-. J. Sri. 264. 5@7-52% O’NEK 3. R. and KHGKAKAY. K. (IQ%) Hydrogen and oxygen isotope exchange reactions between clay minerals and WDt&l*.Cic’ocllim. Co,srtfr~c~liifli. 4cfu 40, 141 246. O*i%n. I. R. and TAYLOR H. P., JR. (1967) The oxygen isotope and cation exchange chemistry of FeIdspars. Arurrr.i%#i~erizt.9, t414-.L437. C&L-ME T-4. D. and %SEK i. H. (1970) ~aieomagn~t~sm of cores from the North f%CifiC. GKJk Sot. Amer. Mem. 126, 83-l 19. S~VIN S, M. (1967) Oxygen and hydrogen isotope ratios in sedimentary rocks and minerals. Ph.D. Thesis, Califbrnia Institute of Technology. SAGAN S. M. and E~sr~rni S. jfY70a) The oxygen and hydragen isotope geochemistry of day minerals. &o&&z. Cnsmochiin?.ACflz-34,2.142. SAWN S. M. and EPS’!I?IN S. (1970b) T&oxygen and hydrogen isotope geochemistry of ocean sediments and shales. CieoLhim.Cosmochim. Actn 34, 43.-63. SHEPPARD S. M. F., NIRSEN R. L. and TAYLORH. P., JR. (1969) Oxygen and hydrogen i.sotope ratios of clay miner&s frmn porphyry copper deposits. &XX G&. 64, 75-777. d~inl. Cnsmochim
?‘AYL.OK H. P.. JK. and
EIJST~IN S. (196&j Relationship between O’s:O” b ratios in coexisting minerals of igneous and rn~~at~~~rph~crock-s. Part I. Friucip$es and
igneous and metamorphic rocks, Part II, Application to petrologic problems. &I!/. &ol. Sot. Amar. 73, 675 -693. YI:I,fH. W. (1974) Oxygen isotope studies of ocean sediments during sedimentation and huriai &genesis. Ph.D. Thesis, Case Western Reserve tmiversity. YEI-IH. W. and S~vrr: S. M. [in press) The rn~han~s~~ of burial n~etarnor~~hjsm of argillaceous sediments: 3, Oxygen isotopic evidence. Eltli. G&. Sot. 4nzrr. in press.
et Coamochim~ca
Acta.
1976. Vol
4tl. pp
Press.
743 to 74X. Pergamon
Prmtcd
in Great
Br~tam
The extent of oxygen isotope exchange between clay minerals and sea water* HSUEH-WEN YE& and SAMUELM. SAVIN Department of Earth Sciences, Case Western Reserve University, Cleveland, Ohio 44106, U.S.A. (Received 2 June 1975; accepted in revised form 8 December 1975)
Abstract-The extent of oxygen isotopic exchange between detrital clay minerals and sea water was investigated by analyzing 0’8/016 ratios of separated fine-grained size fractions of deep-sea sediments from three North Pacific ocean cores. Isotopic results were interpreted according to models based on the assumption that the extent of isotopic exchange should increase with decreasing particle size and increasing time of exchange between the sediment and sea water. The data indicate that information concerning the provenance and mode of formation of detrital clay minerals can be obtained from the 018/0’6 ratios of the coarser-than-O.1 Atrn fraction of deep-sea sediments younger than several million years and the finer-than-O.1 pm fraction of deep-sea sediments younger than several tens of thousands of years. Furthermore, if the extent of chemical reaction between detrital clays and sea water is similar to the extent of oxygen isotopic exchange, such reaction may be important in regulating the chemistry of sea water.
INTRODUCTION OXYGENisotopic equilibrium between a clay mineral and the ambient water is generally approached when the mineral crystallizes (SAVIN and EPSTEIN, 1970 a,b; LAWRENCEand TAYLOR, 1971; SHEPPARD et al., 1969). The isotopic composition of the clay is thus a function of the temperature of formation and of the O’S/O’6 ratio of the ambient water. As long as the oxygen isotopic composition acquired by a clay mineral at the time of its formation remains unaltered, it can provide information concerning the nature of the environment in which the mineral formed. There is considerable evidence of a qualitative nature that post-formational isotopic exchange between clay minerals and water is slow at sedimentary temperatures, except when the clays undergo chemical or mineralogical alteration. SAVIN and EPSTEIN (1970b) concluded that there had been no large-scale oxygen isotopic exchange between sea water and the detrital clays of several Recent and Pleistocene ocean core samples they analyzed. LAWRENCE and TAYLOR (1972) concluded that only insignificant post-formational isotopic exchange of the clay minerals of some Quaternary soils had occurred. SHEPPARDet al. (1969) found no evidence for low temperature isotopic exchange of the hydrothermal clay minerals of some Tertiary and Mesozoic porphyry copper deposits. The sluggishness of oxygen isotopic exchange between clay and water is further documented by the laboratory experiments of O’NEIL and * Contribution No. 115, Department of Earth Sciences, Case Western Reserve University. t Present address: Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena. California 91 109. U.S.A. 743
KHARAKA (1976) who found negligible exchange between illite and water and between kaolinite and water at 100°C over times as long as 2 years. While qualitative data indicate that the rate of isotopic exchange is slow, quantitative data on the rate of exchange have been lacking. It is the purpose of the work reported here to place quantitative limits on the rate of low temperature oxygen isotopic exchange between clays and sea water. Such limits should serve to determine the maximum ages of ocean sediments for which oxygen isotope ratios can provide information concerning the environments of formation of the silicates. Additionally, the extent of oxygen isotopic exchange between clay minerals and sea water is probably similar to the extent of chemical interaction between the two phases. Thus inferences can be drawn from the isotopic data concerning the magnitude of such chemical interaction and its importance in regulating the chemical composition of sea water. ANALYTICAL TECHNIQUES Core samples were ultrasonically disaggregated in distilled water, and were subsequently subjected to the following chemical treatments: 1 N sodium acetate buffer (pH = 5) to remove carbonate; 30% H202 to remove organic matter and manganese oxides; and a sodium bicarbonatesodium citrate solution to remove iron oxides (JACKSON, 1956). These treatments have no effect on the oxygen isotopic compositions of clay minerals (YEH, 1974). The silicate residues thus obtained were separated according to particle size (equivalent spherical diameter) by centrifugation. After mineralogical analysis by X-ray diffraction, the separated fractions were isotopically analyzed. After drying in a glove box at a relative humidity near O”/”for at least 24 hr, oxygen was liberated from the samples by reaction with fluorine (TAYLORand EPSTEIN, 1962a). Duplicate extractions of oxygen were done for most samples. The averape deviation from the mean of the results of the two
Hsr I H-WI \ YI II and SA~LI~ I I M. SA\I~
734
was 113s than 0.2(YC,,, for hO”,, of all samples. Isotopic data are reported in J-notation, as deviation< 01‘the isotopic composition in parts per thousand (‘I,,,,)liom that of the SMOW standard (C‘KAI(;. 1961).
extractions
ISOTOPIC
EXCHANGE MODELS
When detrital &I! minerals of continental origin are incorporated in marine ‘;ediment\ their O’XKI’h ratios are generalI> lower than equilibrium values for clays in the marine en\ironmcnt. Thi\ reflcctz: (I ) the greater 0’“‘O’” ratio of sea water than of most fresh waters: and (2) the colder temperatures of the ocean bottom than of most continental en\ironmcnt\. Isotopic exchange of the detritus toward an equilibrium value can be monitored and expressed using the rclationship:
dctcrmincd cxperimentall! at higher temperatures: comparison of the isotopic compositions of clay minerals which formed at known temperatures with the isotopic compositions of the waters in which they formed: and estimation based on the crystal chemistry of minerals and/or analogy with the isotopic fractionations of similar minerals (TAKOK and EPSTLIN. 1962b; SAVIN. 1967: O’NFIL and TAYI.OK. 1967). Equilibrium fractionation lktors, estimated for a tetnperature of I C. are listed in Table I. The values should be considered approximate. Errors of I T,,,, in the values for smectite and illite. and 3- 4’:,,, for chlorite and kaolinite will not greatly affect the conclusions drawn. a. .S/~~c,c,rirc, ~t‘~l/c’r.The oxygen isotopic fraction bctwecn \mectitc and water was estimated from the anal>Gs of :I Gngle sample of authigenic smcctite taken from core WAHINE I I P (8 16’N. I.53 01’W) supplied h! the Scripp\ Institute of Oceanography. The fractionation factor cstimatcd from the data for this sample is approximately 3.5”,,,, greater than estimated h! S/t\ IU :tnci EP\I I.IV ( I97Oa) from the analyses of t\vo authigcnic smcctitc\. Howc\er. Satin and Epstein did not treat their \mectites to rcmo\‘e oxide mineral\ mid organic maltcr. b. Illitem I~U~CI’. The illilc water fractionation factor at I c‘ wm determined by extrapolating (on a graph of 1000 In 2 vs 1,T’) the illiteewater fractionation estimated for higher temperatures (YEH and SAVIK. in press: ESL.IU,I I< and SAYI\. 1973). c. k’troli~lirv ~v~lrvr. This fractionation factor was estimated by adjusting the value for 17 C given by SA\‘IN and EI?TEI~%(197Oa) to the lower temperature conditions of the ocean bottom. d Chlorirc~ -,c'clter. The chlorite -water fractionation factor was estimated b> combining the fractionation factors for various oxygen containing bonds, as estimated by SARIS (1967).
no,:,:,, ,, is the isotopic composition of the detritus prior to entering the marine environment. CKI,l,;,,,,,,,~,,, iy the value the detritus would ha\,c if it were in cquilibrium with sea water at ocean-bottom temperature. Valid application of the above expression dcpenda on the accurate estimation of 80,t,~,, ,, and &I~‘,;,,,,,,,,,j,,,. 21s discussed below. It should be kept in mind that ;I change in the isotopic composition of a clah mineral. in the absence of evidence for chemical or mineralogical alteration, ma> reflect either of two different mechanisms. A change in isotopic composition is most readily interpreted as reflecting isotopic exchange. Howe\cr, the isotopic data b! themselvzs do not rule out the possibility that the shift in isotopic composition reflects the formation of additional. authigenic material in the marine environment. This material would ha\c to be mineralogicalI> similar to the detritus. and might perhaps form bq the reconstitution of amorphous detrital material as proposed b! MAG+NLII. and GAKKI:LS(1966). Regardless of which mechanism produced the change in isotopic composition of the cla~c. the intcrprctation of the isotopic The O’“‘0”’ ratio of a detrital mineral prior to data and the application of the results arc Gmilar. deposition on the ocean Hoor must be estimated indircctl!. The method most appropriate for this eatimation differs for different samples, and is a function PreciseI) determined oxygen isotope fractionation of sample mineralogy. Two methods applied to the factors between claq minerals and water are not acailresults of this study are described below. Xl‘~tllod .‘I w/1cw r,rinPra/ogJ 01‘ .sun1ple dors irot able. Howebcr, these fractionation factors ma! bc C/KUI~~C~ with prwtidr~ ,sizc. When the mineralogy of approximated in a number of ways, including: extrapolation to low temperatures of fractionation factors each particle-size fraction of a sample is the same, Table
I. Estimated
clay mineral-
water isotopic
Mineral
Fractionation FaCtOr*
smectice llllte Kaolinite Chlorite ____
1.03083 1.02789 1.02645 1.02115
fritctionation
factor5
at
I C
Eq”s-.---,;0181 t30.62 +21.b8 t26.24 +22.95
*
Fractionation factors = (O’X~O’~)~,.,~,(O’X~O1n~~i~~. t Thcsc values are calculated using a sea water ci0”
v;due of
--(Uo”,,,,.
isotope exchange
Oxygen
745
Table 2. Locations, ages and water depths of samples used in exchange rate studies
JYN-II-40
10 to
15
33'04'N
174°15'w
4
31,000
5530
JYN-IV-11G
33 to 39
27s42'N
175*1O'W
2.5
14.000
5750
SCAN 29P
15 to 52
33'16'N
153'4b'E
4
84,000
5857
SCAN 29P
1065 to 1116
33°16'N
153"44'E
4
2,725,OOO
585,
*
Estimated from data in
OPDYKE
and FOSTER (1970).
assumption, the 0L8/0”’ ratio of each size fraction in the uppermost layer of the core may be taken to approximate the initial O18/O16 ratio of that size fraction of the detritus.
it may be assumed that SO~~~~i~~ of each particle-size fraction was also the same. Additionally, it is logical to assume that the extent of isotopic exchange has been greatest for the finest fractions and that, when the sediments are not too old, the coarsest fraction has undergone little or no isotopic exchange. Under those conditions, the isotopic composition of the coarsest fraction can be taken to approximate that of the unexchanged continental detritus. Method B-when mineralogy of a single particle size does not vary with depth in the core. Ocean sediments are complex mixtures of many minerals, and, after the chemical treatments discussed above, generally consist of illite, smectite, mixed-layer clays, chlorite and kaolinite, as well as quartz and feldspar. Commonly the relative proportions of these minerals change as a function of particle size, and in such cases it cannot be assumed that the initial isotopic compositions of any two size fractions were the same. In some such instances, however, the mineralogy of a single particle size fraction may be quite constant throughout the length of a core, reflecting fairly constant geological conditions in the source area(s) of the detritus. Under such circumstances, it is reasonable to assume that the initial isotopic composition of each particle size has remained constant over the time of deposition of the core material. Given that
SAMPLES STUDIED Studies of the rate of isotopic exchange between detrital clays and sea water were made using samples from thrqe North Pacific cores. Data concerning the locations, ages and water depths of the samples are given in Table 2. All samples are predominantly detrital, although SCAN 29P, 1065-l 116 cm contains minor volcanic ash fragments. Depositional ages of samples were estimated from the sedimentation rates determined for the North Pacific by OPDYKE and FOSTER (1970) on the basis of paleomagnetic measurements. None of the samples in the study were from sites studied by Opdyke and Foster. and we have therefore taken averages of sedimentation rates which they calculated for samples from surrounding locations. Ages determined in this way must be considered approximate. Dr. Harvey Sachs of Case Western Reserve University examined radiolarian assemblages from the samples. He determined a biostratigraphic age of 0-400,000 yr B.P. for JYN-II-4G, 10-15 cm (compared to a sedimentation-rate age estimate of 31,000 yr), and a biostratigraphic age of 800,00@5 million yr B.P. for SCAN 29 P, 1065-t 116 cm (compared to a sedimentation-rate age estimate of 2,725,OOO yr). Although precision is not great, ages estimated by these two independent methods are not in conflict. X-ray diffractograms were made of all samples after chemical treatment and size-separation as described ab0ve.t Mineralogic data determined from the diffractograms are summarized in Table 3.
t Diffractograms are available from the authors.
Table 3. Mineralogic
SCAN SCAN SCAN SCAN
29P 29P 29P 29P
SCAN 29P SCAN 29P
SCAN 29P SCAN 29P
data for samples
used in studies of percentage
of exchange
10 to 15cm II II
1 - 0.5 0.5 - 0.2 0.2 - 0.1
7, 80 80
20 20 15
Trl 0 5
3 Tr 0
TT 0 0
+2,.36 +*,.36 +2,.36
33 to 39cm II I,
2 1 1 - 0.5 0.5 - 0.1
74 77 80
20 20 15
0 0 5
5 3 0
1 TT 0
f27.36 +2,.36 l27.36
15 to 52cm I. II II
2 1 1 - 0.5 0.5 - 0.1 so.1
70 72 50 40
13 12 0 0
10 12 50 60
5 3 0 0
2 1 0 0
c27.70 f27.80 +29.2, +30.40
1065 to 1216mo II I, I,
2 1 1 - 0.5 0.5 - 0.1 ‘0.1
70 72 50 40
13 12 0 0
10 12 50 60
5 3 0 0
2 1 0 0
+*,.,o +2,.80 +29.2, +30.40
* Illite. ** Kaolinite
and chlorite.
*** Smectite.
@ Quartz.
$ Feldspar.
# Trace.
HSI I~I-WI u YI H and SAM\ TV_ M. %\I\
746 Table 4. Isotopic
data and percentage of exchange JYN-IV-I IG
for samples
JY N-II-4G
and
L These values have been obtained using method B. * These values have been obtained by adding to the minimum values the percentage exchange determmed for sample JYN II 4G. IO I5 cm. The mineralogy does not vary with size fraction of tither JYN II-4G or JYN IV-I IG. The isotopic data for these samples are therefore amenable to treatment as outlined in method A above, for the determination of hO&,,. Mineralogy varies markedly with size fraction for both SCAN 29 P samples. However, the mineralogic compositions of corresponding size fractions of the two SCAN 29 P samples are similar. The isotopic data for these samples are therefore amenable to treatment as outlined in method 8.
ISOTOPIC
RESULTS
The percentages of exchange calculated for samples from each of the three cores are plotted against particle size in Fig. I. While SCAN 29 P. 1065~.I1 10 cm appkars to have undergone more exchange than either of the younger two samples. this might simply reflect the fact that maximum values for exchange, as described above. arc plotted for SCAN 29 P. If minimum values had been plotted an age effect would not have been nearly so apparent. A particle size etrect is. however. quite evident. This suggests that isotopic exchange at crystal surfaces probably occurs much more rapidly than does diffusion into the interiors of crystals. This is expected in the case of simple
The isotopic data and calculated percentages of isotopic exchange of samples JYN II-4G and JYN IV- 11G are shown in Table 4. Percentage of exchange generally increases with decreasing particle size and ranges to values as high as 16”,, for the 03X).1 itrn / Legend fraction of JYN II-4G. Isotopic results for the two SCAN 29 P samples are given in Table 5. The percentage of exchange was calculated, using method B, for the deeper sample onl>. since the isotopic data from the shallower sample were used as ~io,!,:~,,,, values for the different size fractions. The percentage of exchange calculated in this way is a minimum value. since the iK3,~,~K,,;,, values used were, in fact, obtained from the analysis ......... of samples to which we have assigned an age of 84,000 yr. Maximum values for the percentage exchange of SCAN 29 P. 1065 to 1116 cm can be obtained by ..s.ew adding to the minimum values the percentages of exchange obtained for the different size fractions of a3 10 20 30 .Ol 02 .05 0.1 02 Particle Size (pm) JYN II4G (assigned age 31,000 yr). As with the other two samples. the percentage of exchange of SCAN Fig. I. Per cent exchange vs particle sire for three North 29 P, 1065-I I 16 cm increases markedly with decreasPacific core samples. Values of per cent exchange for SCAN 29 P. 1065 to I I IO are maximum values (see text). ing particle size. I.
-
. .
.
.
. .
. .
.
.
Oxygen isotope exchange diffusion models. If the data are interpreted as reflecting, in part, the growth of authigenic mineral phases rather than the isotopic exchange of existing phases, it could be inferred that the growth of the authigenic phases occurs early in the post-depositional history of the sediment and primarily in the finer size fractions.
DISCUSSION Implications
for
provenance
studies
The results of this study permit qualification of the conclusion drawn by SAWN and EPSTEIN (1970b) that oxygen isotopic exchange between clay minerals and sea water is slow, and that 601’ values of detrital minerals do indeed reflect the environment of mineral formation and can be used as indicators of provenance. The data presented here suggest that over periods longer than a few million years, the 0.1-0.5 pm fraction and coarser fractions of detrital clay minerals in ocean sediments can be expected to have retained their initial oxygen isotopic compositions to the extent that their 60’* values can provide information concerning provenance. Such information can be obtained from the finer-than-O.1 pm size fraction of clay minerals in ocean sediments younger than a few tens of thousands of years. Implications
regarding
the chemistry
of sea water
Chemical equilibrium considerations led SILLEN (1961) and subsequent investigators (HOLLAND, 1965; GARRELS, 1965; etc.) to the conclusion that the interaction between silicates and sea water and the formation of authigenic silicate phases were important in buffering and regulating the chemistry of sea water. MACKENZIE and GARRELS (1966) came to a similar conclusion on the basis of mass balance calculations. Despite the prediction of the formation of authigenic clay minerals and of alteration of detrital clay minerals in the marine environment, several mineralogic and isotopic studies showed no evidence of the presence of significant amounts of authigenic material or of significant alteration of tletrital cla! minerals in mo
141
results of these studies did not preclude the presence of small amounts (up to several per cent) of authigenic or altered detrital material. The approach used in this study to estimate the percentage of isotopic exchange of separated size fractions of a sediment may be extended to obtain an estimate of the extent of isotopic exchange of the entire sediment. Since isotopic exchange involves the breaking and reforming of the oxygen-containing bonds of a mineral, it is unlikely that a great deal of isotopic exchange would occur without chemical reaction also taking place. It is therefore probable that the extent to which a sediment approaches isotopic exchange equilibrium with sea water corresponds to the extent to which detrital clay minerals have been chemically altered or authigenic clay minerals have been formed. The total percentage of isotopic exchange of a sample can be calculated from the extent of exchange of the different size fractions. Total per cent exchange
= 1 Aei,
L
where fi is the fraction of a sediment which falls in size range i and ei is the percentage of isotopic exchange of size range i. Values offi were determined only for the two SCAN 29 P samples, and were virtually identical for those two samples (> 2 pm, 30%; 2-I pm. 147;; l-O.5 pm, 14%; 0.5-0.1 pm, 37%; ~0.1 pm, 5%). In the absence of such data for the other two samples, this size distribution was assumed to approximate that of all samples. Calculated values of total percentage of isotopic exchange are given in Table 6. They indicate small but significant amounts of isotopic exchange of all sediments, with 4-9% of the detrital ocean sediments having undergone isotopic exchange in period of 104-lo6 yr. If a corresponding amount of chemical reaction occurred, the effect of this reaction on the chemistry of sea water could be considerable. For example, on the basis of material balance calculations, MACKENZIE and GARRELS (1966) proposed that the chemical composition of sea water would be maintained at steady state by the reconstitution of degraded aluminosilicates and the formation of a quantit? of authigenic cla! minerals equivalent to about 7”,, of the mass of ocean sediments.
Table 6. Total percentage of exchange of samples
Ackrruwlr~yemerlrs-Wi: thank Dr. JOHN HOWEK for providing advice and guidance throughout the study. Dr. HARVEY SACHS br his biastratigraphic determinations, Ms. IVY %3 and Ms. ?RMF, ~~EHLI for assistance in sample analysis and the Scripps ~~st~~nr~onof Oceanography for providing samples. Financiai support was provided by National Science Foundation Grants GA-4016 and GA- 16827. REFERENCES
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EIJST~IN S. (196&j Relationship between O’s:O” b ratios in coexisting minerals of igneous and rn~~at~~~rph~crock-s. Part I. Friucip$es and
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