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The oxygen and hydrogen isotope geochemistry of clay minerals
GTeocNmica et CosmocNmlca Acts,1970,Vol.34,pp.25to 42. Pergamon Press.Printedin Northern Ireland
The oxygen
and hydrogen isotope geochemistry
of clay
minerals*
SAMUEL M. SAYIN? and SAMUEL EPSTEIN Division of Geological Sciences, California Institute of Technology Pasadena, California 91109 (Received 31 October 1968; accepted in revised form
4 July 1969)
&s&t-Oxygen and hydrogen isotope analyses have been made on a variety of clay minerals of sedimentaxy and diagenetic origins. The interlayer water of clay minerals ws~ found to exchange rapidly with atmospheric water. Conditions under which the interlayer water could be removed from the clays without rtffecting the isotopic compositions of their aluminosilicate oxygen and hydrogen were therefore determined, and the interlayer water was routinely removed and discarded prior to isotopic analysis. Approximate fractionation factors for clay mineral-water systems at sedimentary temperatures, inferred from the isotopic compositions of natural samples, are: Oxygen
%&-Hz0 Montmorillonite Kaolinite Glauconite
1.027 1.027 1.026
Hydrogen %in-H,O 0.94 0.97 0.93
Consistency of a relationship between oxygen and hydrogen isotope ratios of kaolinites taken from a variety of areas and having different isotope ratios was demonstrated. This relationship results from exchange of clays with different meteoric waters under conditions in which the fractionation factors are relatively constant. This strongly suggests that the fraction. ation factors are equilibrium ones and that kaolinite forms in isotopic equilibrium with its environment. Deviations of the isotopic composition of montmorillonites from a similar rel&,ionship have been interpreted aa resulting from isotopic exchange at slightly elevated temperatures. There was no clearly demonstrable case of clay minerals undergoing isotopic re-equilibration at sedimentary temperatures, although this is one possible interpretation of the data of some Upper Cretaceous glctuconites. INTRODUCTION
of studies of oxygen isotope ratios in sedimentary carbonates (EPSTEIN DEIJENS and EPSTEIN, 1964; KEITH and WEBER, 1964, and others), have led to a greater understanding of the geochemistry of carbonates. In contrast, there exists only a small body of data pertaining to the stable isotope geochemistry of noncarbonate sediments. SILVERMAN (1951) first showed that oxygen isotope ratios of silicate minerals in sedimentary rocks can be indicative of their modes of formation. For example, detrital components from igneous and metamorphic rocks generally contain less 018 than minerals which have low temperature origins. Prior investigations have also indicated that isotopic exchange can occur in sediments which have not undergone obvious metamorphism. Ancient tie-grained marine A NUMBER
et al., 1961;
* Contribution No. 1540, Division of Geological Sciences, California Institute of Technology. t Present Address: Department of Geology, Case Western Reserve University, Clevelmd, Ohio 44106. 26
26
S. SAVIX
mid
S. ICESTEIN
limestones and cherts often undergo a change of oxygen isotopic composition withont showing any obvious metaKnorphi~features (DEOENSand EPSTEIX,1962). GABLICK and DYHOND(1967) observed that volcanic glass shards in Miocene and oIder ocean bottom sediments had undergone isotopic exchange with sea water. The present series of studies (see also GAVINand EPSTEIN,197Oa, b) consists of investigations of the oxygen and hydrogen isotope geochemistry of silicate minerals which form in sedimentary or dingenetic environments. In particular, it has been of interest to determine whether minerals which are formed under sedimentary conditions are formed in isotopic equilibrium with their environments, and under what conditions these authigenic minerals undergo subsequent isotopic exchange. This information is necessary to evaluate the possibilities of using the isotopic compositions of minerals as aids in understanding some of the problems of sedimentary geology and geochemistry. TERMINOLOCZV AND
NOTATION
All isotope ratios are reported in the 8 notation.
The &values are defined as:
1 C SD = (D/H) sampk _11x100 .
‘01* =
(01s/016) sample (018~016)~td
-
1
x
1000
(D/H) std
Oxygen is reported as per mil deviation from the standard while hydrogen is reported as per cent deviation. In both cases the standard reported throughout this work is Standard mean ocean water (SMUT) , m defined by Crwa ( 196la). The fractionation factors between two compounds or phases containing oxygen or hydrogen are defined by the equations:
uF_g
=
018 ( 0161 A 018 (i0’6 B
D and
0FiA u;?_n = D 0Hts
where A and B denote the compounds or phases. BIKELI1VIINARY GONSIDEIUTIONS A number of factors can affect the isotopic compositions which a clay mineral acquires during sedimentary and diagenetic processes. Among these are: (I) The isotopic composition of the various waters with which the clay mineral may have come in contact during and after its formation; (2) The temperature of the environment at any time during which the clay mineral was subject to isotopic exchange ; and (3) Whether the mineral reached isotopic equilibrium with its env~onment. CRAIG(1961b) has shown that most meteoritic waters, other than those in evaporating basins and hot springs, have isotopic compositions lying near a straight line described by the equation BD = 0.8 601* + 1-O. The isotopic compositions of many ground waters behave according to this relationship (EPSTEINet d., 1966). However,
!L’heoxygen and hydrogen isotope geochemistry of clay minerals
27
CLAYTON etal.(1966) have shown that the isotopic compositions of a number of oil field brines lie to the right (low 8D or high 801s) side of the line. They interpret this to be the result of isotopic exchange of meteoric water with a relatively large amount of marine carbonate rock, a material relatively rich in 0’8, causing the water to be enriched in 01* while maintaining its original D/H ratio. It may be expected that the isotopic compositions of most waters encountered in sedimental processes and some waters encountered in diagenetic processes are meteoric and can be described by the above straight line relationship. Since the 01*/016 and D/H ratios of meteoric water are related and since the isotopic compositions of clay minerals depend on the isotopic composition of water, there should also exist a relationship between the oxygen and hydrogen isotopic impositions of clay minerals. It can be shown that if clay minerals undergo isotopic exchange with a variety of meteoric waters under conditions such that the fractionation factors OL&-HpO and a%ik-EpOare constant (i.e. isotopic equilibrium at constant temperature) then the isotopic com~sitions of the clays will lie along a line with the equation 6D = A 601* + B where
hy A = 6.96 x %%&!
M 9.99
OL&-HpO
and hy
B = 800 x afzy-Hso - 699o(;&+~ - 100.
aclay-H10
Thus, the slope of the line is determined by the ratios of the a’s, and its inbrcept is determined by the absolute values of the a’s. It should be noted that the value of the intercept, B, is extremely insensitive to large variations in the values of the a’s, and thus fractionation factors cannot be calculated with any degree of precision directly from the slope and intercept of a line through a set of experimental data. In the case of eq~ib~um isotopic exchange the effect of temperature on the ~actionation factors may be estimated. Analogy with mineral-water systems for which the oxygen isotope fractionation factors have been experimentally determined indicates that the clay-water oxygen isotope fractionation factors should approach unity with increasing ~rn~rat~e. ~though less information is available with respect to hydrogen isotope fractionation factors, there is evidence (SAVIN, 1967) that these, too, approach unity with increasing temperature in the temperature range of interest. As a result of the temperature dependence of the fractionation factors it may be expected that curves drawn through the isotopic compositions of clay minerals which have exchanged with meteoric waters will lie closer to the line of meteoric water at higher temperatures than at lower ones. These relationships provide a useful working model with which the oxygen and hydrogen isotope vacation in clay minerals may be interpreted. ANALYTICAL TECHNIQUE Oxygen Oxygen for isotopicanalysis was liberated from minerals using fluorine.oxidation techniques similartothose desoribed by TAYLOR and EPSTEIN (1962)andCLAYTON and MAYEDA (1963).
28
S.
SAWN and 8. EPSTEIN
The oxygen was converted to carbon dioxide which was then circulated over hot mercury prior to analysis in order to remove any fluorine or bromine which might inadvertently have escapecl entrapment in the liquid nitrogen traps. The step was instituted to safeguard the mass spcct.rorl, eter from contamination. The isotopic composition of the carbon dioxide was measured on a double collecting mass spcctromcter (NIER, 1947; MCKINNEY rt al., 1950). The reproducibility of the analyses was generally & 0.2 X0 (average deviation) for minerals which do not contain molecular water, 10.3 x0 for clay minerals containing interlayer wat.er, and $:0*7 :“;, for zeolit,os. The measured S-values were related to the SMOW standard through a secondary standard of St. Peter Sandstone which was analyzed with each set of six samples. The sOla value of this sandstone from the same bat!ch used by GARLICK (1964) and GARLICE and EPSTEIT (1967), was taken to bo + lo*9 per mil. Hydrogen. Hydrogen was liberated from hydroxyl bearing minerals by heating in a vacuum to 900°C after preliminary outgassing (see below). Most of the hydrogen was liberated in the form of water, but some was liberated as molecular hydrogen. This was converted to water by reaction with copper oxide which was placed in the furnace next to the sample. The resulting water was quantitatively converted to hydrogen by reaction with hot uranium in the manner described by CRAIG ( 1961 b) and GODFREY ( 1962). Isotopic analyses of the hydrogen were performed on a mass spectrometer of the type described by NIER (1947), and FRIEIXXAN (1953). Reproducibility of the analyses was generally better than 10.3 % (average deviation).
All day minerals contain both hydrogen and oxygen in their silicate frameworks. In addition, many clays contain molecular water which occupies structural positions between the silicate layers, They also contain varying amounts of adsorbed water. Prior to making routine isotopic analyses it was necessary to determine the ease with which the interlayer water exchanged isotopically with the atmosphere and t,o determine the optimum conditions for separating interlayer and adsorbed water from the clay minerals without altering the isotopic compositions of the minerals themselves. Details of experiments made to study the above questions are given by SAVIN (1967). The results are summarized below. 1. The interlayer water of cllays isotopically exchanges almost completely with atmospheric water vapor in a few hours. Thus, the isotopic composition of the interlayer water reflects the isotopic composition of the water or water vapor in which the clay was last stored. Therefore in routine analyses the interlayer water was discarded and the hydrogen and oxygen of the silicate frameworks of the minerals were analyzed. 2. Interlayer water can be removed from clay minerals by heating in a vacuum at temperatures of lOO”-250°C for two hours or more. No isotopic exchange between the water and the hydroxyl hydrogen of clays was noticed when the water was removed in this manner. In routine analysis of hydrogen in clay minerals, samples were outgassed in this manner prior to evolution of the silicate (hydroxyl) hydrogen. 3. Nest of the interlayer water of clay minerals can be removed in 24 hr in a drybox in an atmosphere dried by p&. The small amount of interlayer water remaining in the clay does not normally affect the 6018 value of the aluminosilicate oxygen by more than a few tenths of a permil. Therefore, in the analysis of oxygen, the pretreatment described above w&8 routinely used.
RESULTS AND DISCUSSION
Oxygen and hydrogen isotopic analyses were performed on a number of kaolin&es, montmorillonites, and glauconites from a wide variety of areas. The results of these analyses are listed in Table I. and presented in Figs. Z and 2 where they are compared with the analyses of some igneous and metamorphic rocks and minerals. The measured 01*/016 ratios of the kaolinites and bentonites have been corrected for the
29
The oxygen and hydrogen isotope geochemistry of clay minerals Table 1. Isotopic composition of clay minerals Av. dev.
&s8 (%,)
Av. dev.
No. runs
Langley, So. Carolina Murfreesboro, Arkenses Bath, So. Carolina Mesa Alta. N. Mexico Kadan. Czech. Podlesi, Czech. Sedlec, Czech. b. Halloysite
+21.2 + 22.4 +22.1 + 18.7 + 19.4 + 19.0 + 19.3
0.3 0.3 0.1 0.6 0.0 0.2
2 3 2 2 1 2 3
-6.7 -6.1 -6.6 -7.3 -7.1 -7.2 -7.9
0.1 0.3
Bedford, Indiana o. Dickite
+ 22.0
0.4
2
-6.0
-
1
Kladno Zapotocky, Czech. Horni Slavko, Czech. d. Montmorillonite (bentonite)
+16.6 +16.3
0.6 -
2 1
-3.7 -3.6
-
1 1
Polk&&, Miss. Chambers, Arizona otay, Calif. Clay spur, Wyo. Black Hills, 5. Dakota Little Rook, Arkansas 8. Montmorillonite (no&o&e)
+28.6 +19.1 + 19.0 + 17.6 + 16.9 + 26.0
0.2 0.4 0.4 0.2 0.4 0.0
2 7 4 6 3 2
-6.4 -7.9 -4.1 -9.9 - 12.6 -6.8
0.2 0.3 0.1 0.4 -
Ma&q Wash. f. Montmorillonite (ooean oorea)
+13.2
-
1
-10.9
-
1
RI8 81 EM?‘-RUN g. Illite
t26.1 + 28.6
0.0 1.0
2
-7.8 -6.8
0.8 -
2 1
+ 20.6
0.6
2
-6.1
0.2
2
Sta 1762 (Blake Plateau) Sta 1967 (‘L.I. Sound) Sta 6941 (S. Moniaa Bay) Glauoonite (from rooks)
+ 26.3 t21.8 + 23.8
0.6 0.0 -
6 2 1
-7,4 -9.0 -8.6
0.2 0.1 -
4 3 1
RB-Memhantville Fm R12-Marshalltown Fm R13-Navesink Fm R17-Hornerstown Fm E53/3/14 Crawford Fm Australia E63/3/16 Crawford Fm E63/3/16 Crawford Fm E63/3/18 Crawford Fm E63/3/19 Crawford Fm
+ 20.3 +23.0 + 23.0 + 22.6
0.0 0.6 0.1 0.3
2 3 2 3
-7.8 -8.0 -8.6 -8.6
0.6 0.2 0.2 0.1
3 3 3 2
+ 14.6 +13.7 + 18.6 +17.0 + 16.4
0.3 0.4 0.6 0.2 0.1
2 3 4 2 2
-6.3 -6.2 -4.1 -6.4 -6.4
0.3 0.8 0.2 -
3 4 2 1 1
Sample
6D (%,
No. -
a. Kaolinite
1
Morris. Ill. h. Glaucdnite (from ocean
i.
0.2 0.0 0.1 -
1.0
10 2 1 2 1 2 3
2 3 3 11 1 4
aediients)
contribution of quartz and feldspar, using the mineral percentages given by KERR al. (1950) and KONTA (1965). These corrections generally were between O-1permil and 0.3 permil. As the data of Fig. 1 show, the oxygen in the clay minerals is richer in 018 than that of nearly all the igneous and metamorphic rocks and minerals for et
which data are available. Wherever there are a large number of specimens of a particular type of clay mineral from different localities there is a broad spread in the values of the oxygen isotope ratios. This spread in &values is due to the variety of conditions under which the minerals have formed and exchanged. The hydrogen isotope data likewise show a wide range of &values, but unlike the oxygen isotope data the range of BD values for clay minerals is similar to that for igneous and metamorphic muscovitea and biotites. All the clay minerals are depleted in deuterium with respect to sea water, including a number of samples which
8. SAVIN and S. EPSTEIN
30
+4
Igneousreck?, Meiomorphlc
+6 +8 +I0 +I2 +I4 +I6 +I8 +ZOA2+24+26+2&30 / I I I, I I 1 I I,
I
I
/
t---i
’
I_-_-----,
rock? ’
Metamorphvz muscovlte b-j Kaollnlte Citckl!e
H
Ho!loysite
I
Mcntmorillonite (bentonite) Montmorillonite (ocean cores) Glauconite
I
t ! I:
t1
I
II
,
I It+H--i
!Illk
I o
Taylor 8
b. Gorlick
L
Epstein
1i96.2)
a Epstein
11967)
Fig. 1. The 6CP values (relative to Standard Mean Ocean Water) of clays analyzed in this study compared with pertinent isotopic data of previous workers.
6 -13 -12 -I/( -I?
t
Metamorphic
biotite’
Metamorphic
muscovite
Metamorphic
chlorite
I
-?
D (%)
-B
-7
-7
-T
-4
-2
I
-/
?
I
I
’
1 I Ii
a
/
I
I
Kaolmite
-3
I
4 III H
Dicklte
I
Halloysite Montmorillonite (bentonitel Montmorillanite (ocean cores) Glauconite
t
I
I
I
1
{
H
I
Illde
o Taylor61 Epstein
0966)
Fig. 2. The 8D values (relative to Standard Mean Ocean Water) of clays analyzed in this study compared with pertinent isotopic data of previous workers. appear to have formed in isotopic equilibrium with sea water. Thus, while a$&_np is greater than unity at normal ocean temperatures c&&_~,~ is less than unity.
Kaolinitte group The oxygen and hydrogen isotopic compositions of several kaolin&es, dickites, and a halloysite are listed in Table 1 and are plotted on a graph of 6D vs. 6018 in Fig. 3. All the kaolinite points and the halloysite point lie near a, straight line on this plot while the dickite points fall well off the line. This is consistent with the proposed model in which the isotopic compositions of the kaolinites are determined by
The oxygen and hydrogen isotope geochemistry of clay minerals
31
0
1:
_ 0
q
Koolinite Holloysite
_3
A
Dickite
-4
&
-5 s -6 0 -7 ._ co -6 -9 ; -10
:ii
-II -12 -13 -
”
”
-14 ” ” ” ” ” ” ’ 12 I3 14 15 I6 n I8 I9 20 21 22 23 24 25 26 27 26 29 30 6 o'* ("/ 001
Fig. 3. Graph showing relationship between 6D and &O18values of some keolinitea, dickitaa and a halloyeite. Line L calculated corresponding to a&=,,, = l-0266 and uhy W&&#J = 0.970.
isotopic exchange with meteoric waters under conditions of relatively constant fractionation factors. An estimate of the fractionation factors for the kaolinite-water system has been made by relating the measured isotopic compositions of some of the samples to the best estimates available for the isotopic compositions of the waters in the presence of which they formed. All of the kaolinites analyzed in this work are of residual weathering or sedimentary origin. All have probably formed and been deposited in fresh water environments (see Appendix for geological and petrographic descriptions of the samples). All were colleoted from shallow quarries or outcrops in regions of temperate climate. Hydrogen isotopic compositions of waters from streams and springs in areas near those from which some of the ksolinite samples were collected were published by FRIEDMAN et al. (1964). If it is assumed that the kaolinites attained their present isotopic compositions in the presence of waters of the same isotopic compositions as those measured by FRIEDMAN et al.,under surface temperature conditions, then we are able to estimate the fractionation factors a”&_H,Oand ah&,_uIo. These estimates are shown in Table 2. The weakness of these estimates is apparent. The waters were sampled only once from each location, thereby ignoring any seasonal isotopic variations. They were sampled from streams and springs, and the waters analyzed thus may not represent waters which were in contact with the clay samples. These effects would be expected to be smallest for the samples from the relatively low-lying regions of the southeastern United States, and greatest for the sample from the mountainous region of New Mexico. Finally, the isotopio compositions of the waters at the time during which equilibration occurred may have been significantly different than they are now.
33
S. SAVINand S. EPSTEIT Table
2. Estimate
of the isotopic fructionation factors of tho kaolinit+wator sy&em at sediment,ary temperatures
-5.1
“j- 22.4
-5.1
+
Beth. South Czwolina
-5.6
j21.1
Langley, S. Caroline
--5,7
+21.2
Mess Alta, New Mexico
-7.3
$18.7
Murfreesboro Arkanses
22.4
Red l-t., shrcvcpurt~, La. Ouachita 12.. Monroe, La. Ogoocheo R., Ft. McAllister
-2.7
o.!n;i
1.0271
- 3.0
-5.0
0.978
I.0275
-2.1
-3.9
11.96C
I*0261
-2.1
__ 3.9
0.963
1.0262
I.003
I+0295
c-a
Ogeeoheo . R., Ft. McAllister Ca. Sulfur Spring, Jemez Mtns., Xew Mexico
4.6
-7.6
-10.8
* Hydrogen isotopic compositions of wsters from FF.IEDALU et al. (1984). t Oxygen isotopic oompoeitions of w&em estimated using reletionship SD E 0.8 601* +
1 (CZAIQ.
1981b).
The calculations in Table 2 suggest values of the fractionation factors of a&_H,O 19265 and shy kaol_E,o= 0.970. These are averages of the calculated cc-valuesfor the kaolinites from South Carolina and Arkansas. A line corresponding to these ~a~tio~ation factors has been drawn in Pig. 3. Keeping in mind the assumptions on which the calculations were based the agreement of these values leads us to believe that the fractionation factors are equilibrium ones at surface temperatures. The halloysite point falls quite close to the line through the kaolinite data. It is likely that the aluminosilicate portion of halloysite behaves isotopioally very similarly to kaolinite. KERR et at. (1950) have suggested that this pa~i~ul~ specimen was deposited from solution. If so, it was probably formed from a low temperature solution or else underwent subsequent exchange at low temperatures. The isotopic data of two dickite samples lie well above the line drawn through the kaolinite data. The structural difference between kaolinite and dickite is a very minor one, consisting of a small difference in the way the aluminosilicate layers are stacked, If kaolinite and dickite behave simply with respect to isotope fractionations then factors which might cause the dickite points to lie away from the kaolinite line are: equilibration of the dickite at elevated tempemture, exchange with water of unusual isotopic composition, or incomplete isotopic exchange. One of the dickite samples (Horni Slavkov) is of hydrothermal origin. Its temperature of formation was estimated by KONTA (1957, 1961) to be approxima~ly 3OO’C or 350°C. The second dickite sample (Kl~o) occurs with sulfide minerals in fractures and sedimentary rocks. While KONTA (1957) interprets its origin to be low temperature rather than typically hydrothermal, isotopic evidence indicates that it was formed at temperatures higher than surface temperatures, or was subject after formation, to isotopic exchange at somewhat Ellevated temperatures It appears likely that both dickites underwent final isotopic exchange at similar temperat~s and that these temperatures would be below the ~mperature of formation estimated for the Horni Slnvkov dickite. By analogy with the muscovitewater oxygen isotopic fractionation data of O’NEIL and TAYLOR(1968) the value of 4&ite-n 0 in the temperature range 2OO”C-350% can be estimated to range =
The oxygen and hydrogen isotope geochemistry of clay minerals
33
between l*OOland 1.004.Thus formation of a hydrothermal dickite with a 80is value of j-154 would require the presence of a water whose 6018 value is approximately 13 permil. Another explztnation is that the isotopic compositions of the diokites were established by either partial or complete isotopic re-equilibration with waters of normal b-values at temperatures lower than 200°C. If it is aaeumed that the diokite last equilibrated with water of isotopic composition between - 7.4and 0 permil the temperature of last equilibration can be estimated to range between 40% and 1OO’C. M ontmorillowite
Two montmorillonite samples taken from ocean cores and formed by the aubmarine alteration of volcanic rooks, several bentonites of marine and non-marine origin, and a, nontronite were isotopically analyzed. The hydrogen isotope imposition of these montmor~o~~s range from approximately 8D = -4 to - 10 per cent ctnd the oxygen isotopic compositions range from about 80’* = + 17.5 to + 28G permil. This large range of variation strongly suggests formation or exchange under a wide range of conditions of isotopic environment and temperature. The isotopic data, are hated in Table 1 and have been plotted in Fig. 4 on a graph of 8D vs. 8018.
-5 -6 -7 -6 -9 -IO -I -I2 -13 -14
12 13 14 15 16 I?
I8
19 20 21 22 23 24 25 80’8
26
27 28 29
30
(%o)
Fig. 4. Gr@h showing rektionship between 6D and 6018 values of some montmorillonites from ocean cores and bentonite deposits. Line is e&ulsted using constants u&,~_~~~ = l-0273 and &5,+EIo = O-938.
The line on the diagram has the equation 6D = O-73 6018 - 26. This line was drawn using the fractionation factors &$+n,o = 1.0273 and eQnt_n,O = 0.938, which were estim&~d by taking the evemge of the ~lc~&~ ~&otion&tion factors between ~a. water and the four montmorillonitea richest in 018. It mrtybe noted that the line dra,wn pressesabove the average value of the four marine montmorillonites used to determine this line. This resulti from the fact that the line describing the isotopic compositions of meteoric waters lies a similar distance above the point 3
:i4
S. SAVLNand S. E:psmm
corresponding to the composition of ocean water (SMOW). All four of these samples were taken from marine beds or ocean cores and have similar isotopic compositions, suggesting that they were formed under marine conditions. The values of the fractionation factors estimated for these samples are of comparable magnitude to fractionation factors estimated between other clay minerals and water. The points corresponding to the isotopic compositions of the bentonites with the lower 018/016 ratios lie somewhat above the line drawn using the isotope data for samples richest in On?. The isotopic compositions of these sa,mples of lower 01s content are best understood in the light of the model of isotopic exchange with meteoric waters at temperatures above those at which the O’* rich samples equilibrated. By analogy with the effect of temperature on fractionation factors in other systems (e.g. carbonate-water) it may be inferred that these temperatures need not have been considerably warmer than surface temperatures, probably below 50°C for most samples. Other possible explanations for the isotopic data of these samples include variations of the isotopic compositions of meteoric waters from the ideal straight line relationship, incomplete exchange, and variations in the fractionation factors due to differences in the chemical compositions of the minerals. Most deviations of the isotopic compositions of meteoric waters from the normal straight line relationship lie on the right hand side (low 6D and/or high 6018) of the line (CRAIG, 1961b; CLAYTONetal., 1966). This is in theoppositedirection to the effect observed for the clay minerals, and thus variations in the isotopic compositions of waters cannot satisfactorily explain the observed isotopic compositions of the clay minerals. Chemical analyses and structural formulas of all the bentonites analyzed in the present study have been reported by KERR et al.(1950). There appears to be no correlation between chemical composition and isotopic composition of the minerals, and thus there is no suggestion that the chemical composition is responsible for the low 018/01a and/or high D/H ratios. Incomplete isotopic exchange cannot be ruled out as a possibility for the montmorillonites studied. While it is quite possible that the hydrogen of clay minerals could exchange with water more rapidly than the oxygen, it is unlikely that the reverse be the case. Thus, the points representing the isotopic compositions of clay minerals which formed under marine conditions and subsequently underwent incomplete isotopic exchange with fresh waters might lie below the equilibrium line on a plot of 6D vs. 60 18, but would not lie above. Similarly, the data representing the isotopic composition of a clay equilibrated with a light water and subsequently exchanged incompletely with a heavier water could lie above the equilibrium line but not below. Thus, if marine bentonites having isotopic values which lie above the equilibrium line acquired their isotopic compositions through incomplete isotopic exchange it must have been through a complicated series of events. The fact that the isotopic compositions of meteoric waters plays a major part in determining the isotopic composition of the clay minerals is emphasized in the relationship between the d-values of the clays and the locations from which they were sampled. Those clays with the lowest 01* and deuterium values are from the northern United States (Wyoming, S. Dakota and Washington). Those richest in 01* are from ocean cores and the southeastern United States.
The oxygen and hydrogen ieotope geochemietry of clay minerals
36
In summary the most reasonable interpretation of the montmorillonite data is that the four samples which are richest in 01s formed in isotopic equilibrium with sea water at sedimentary temperatures and did not undergo subsequent isotopic exchange, and that the Ol*-poor samples have undergone isotopic exchange with meteoric waters at temperatures greater than sea bottom temperature.
In the present study only one sample was analyzed. This sample, of non-marine origin, was taken from the underclay of a Pennsylvanian cyclothem at Morris, Illinois. It has a 80’s value of +204 permil and a 6D value of -51 per cent. The sample is thus isotopically not unlike other non-marine clays. Because only one sample was measured, it is not possible to discuss isotopic variations of illites in sediments. Blauwnite The isotopic data for the glauconite samples are listed in Table 1 and shown in Fig. 6. There are two distinct groups of samples evident in this figure, one set richer in 0’s and poorer in deuterium than the other. The O%.ich set of samples consists entirely of glauconites formed during or since the Late Cretaceous, and none have been subject to deep burial (OWHINS et al., 1961; EMERY,1960). The Ols-poor set of samples are from a suite of Precambrian Australian rocks and are thought to have been buried at some time during their history to depths ranging from one thousand to seven thousand feet (PLUMB,1965). The Mesozoic and Cenozoic samples can be considered to be more representative of glauconites forming or exchanging on or near the earth’s surface under conditions similar to those found at the present time.
0 b
-5
-
-8
-
-9
-
A
Recombrim
0
Ocean sediments
rocks
-10 -II
-
-12 -13 -14
’ ’ ’ ’ ’ ’ ’ ’ ’ ’ ’ 12 13 14 15 16 17 19 I9 20 21 22 23 do’*
’ ’ ’ 24 25 26
’ ’ ’ 27 26 29
(%I
Fig. 6. Graph showing relation&p between 6D and 601e values of Borne glaucon&e taken from rocks and ocean eedimente. Line a ie calculated wing conetants ~&+H,o
= l-0263
ami abh$ac_H,O = 0.936. Line b is drawn through Precambrian data parallel to line a.
A line corresponding to the equation 6D = O-72 6018 - 25 llas been drawn OII Fig. 5. This line corresponds to the fractionation factors ~~~Uc_II,O -_ O-926 and ~$L-H20 = 1.0263 which are calculated from the assumption that sample Sta 1752 has undergone complete isotopic exchange with sea water. This sample, the richest, in Ols of all the glauconites mea,sured, was dredged from the surface of the Blake Plateau (31’28*7’N, 72”29*O’W) in the Atlantic Ocean. A temperature of approxi. mately 6°C prevails on the sea floor in this area (STE~AK~, 1962). The isot,opic composition of the water in this area is probably close to 0.0 permil. It is uncertain whether this glauconite is of detrital origin from the continents or is forming in sit,u (TRUMBULL, 1965). With only one exception (R-8), data for the remaining Mesozoic and Cenozoic glauconites lie on or very close to the line drawn, in a manner compatible with the model of exchange with fresh waters under conditions of constant fractionation factors. If this model is a valid interpretation of the isotopic compositions of these minerals, and if it is also true that they have formed under marine conditions, as interpreted by OWENS et al. (1961), then it is the only case where minerals formed in a marine environment appear to have undergone isotopic exchange at sedimentary temperatures. While the above explanation is possible, there is another interpretation of the data. The isotopic compositions of the five samples which lie near the line are clustered together in a small region. It is possible that some or all of these points represent glauconites which were also in isotopic equilibrium with sea water and did not subsequently undergo isotopic exchange, provided that the sea waters with which the samples equilibrated were appropriately diluted with isotopically light fresh water, and/or that the temperatures of formation were warmer than for the Blake Plateau sample. The two facts which raise the suspicion that this may actually be the case are: (1) Th e isotopic compositions of these samples cluster together, and it is unlikely that all of them would have exchanged completely with waters of similar isotopic compositions at similar temperatures: and (2) One of the samples (Sta. 6941) was dredged from Santa Monica Bay, California and has very likely formed in sea water and remained there since formation (EMERY, 1960). Factors which support the proposed mechanism for obtaining the observed isotopic compositions are that all samples were formed in areas which might be diluted by fresh water, and that four of them were formed during the late Cretaceous when temperatures were warmer than at present and when the 6018 and 6D of the oceans could have been lower because of the absence of isotopically light polar ict* caps. The isotopic composition of one Tertiary sample (R-8) lies well above the line drawn. The deviation of this sample from the line can be caused by exchange with fresh water under conditions of elevated temperature. The deviation may also be due to the difficulty encountered in removing 018 poor detrital minerals from this sample. A line of slope similar to that through the data of the younger samples may be drawn through the points representing the Precambrian samples, but is considerably displaced from the other. One possible explanation for this displacement is that the isotopic composition of sea water may have changed since the time of formation of
The oxygen and hydrogen
isotopegeochemistryof clay minerals
37
the samples 1390 million years ago (MCDOUGALL et uZ., 1966). If the isotopic composition of sea water were approximately 3 per cent richer in deuterium and 6 permil poorer in O1s 1390 million years ago than at the present time, then the two sets of data,would be compatible. A change in the oxygen isotopic composition of se&water in this direction has been postulated by WEBER (1965) and further supported by PERRY (1968). However, the isotopic data of the Precambrian glauconites can also be explained by exchange with ground water at an elevated temperature which need not have been above about 100°C. The latter appears to be a likely explanation for at least part of the observed effect since the sandstones from which the samples were taken are cemented with secondary quartz and the rocks were therefore probably subject to invasion by silica-bearing solutions with which isotopic exchange might have occurred. In addition, there is a spread of five permil in the oxygen isotopic compositions of the glauconites indicating that at least some of them underwent exchange after formation. The observed range of O1s/O1sratios is larger than seems reasonable to have occurred in marine waters in a fairly short length of time and in a fairly small area. The K-Ar dates for these samples are the same within experimental error (MCDOUGALL et al., 1965). SUMMARYAND CONCLUSIONS The variations of the oxygen and hydrogen isotope ratios in kaohnite, illite, montmorillonite and glauconite are relatively large. The ratios are dependent on the conditions of formation and exchange of the clay minerals. When a series of clays have undergone isotopic exchange with different meteoric waters under conditions of constant a~&R,o and a&+rlo (i.e. equilibrium exchange at constant temperature) their isotopic compositions show a relationship of the form 8D = A 601s + B. Clays which have exchanged under equilibrium conditions with water at elevated temperatures will have more positive values of 8D and more negative values of 601e. They will therefore lie above the line drawn for sedimentary temperatures. Of all the clay mineral data, the kaolinite data show the most systematic behavior in terms of exchange with different waters under conditions of constant aE,,l_HIo and c&,~_R~~.The near constancy of these parameters strongly suggests that the exchange resulted in isotopic equilibrium. Although kaolinite forms under a wide range of climatic conditions, pure kaolinite deposits such as the ones considered here tend to be formed in zones of quite intense weathering. This weathering is favored by a mild wet climate and so the kaolinite samples measured here were probably formed within a fairly narrow temperature range, thus more closely fulfilling the conditions of the model than other sample suites. The persistence of a lsrge difference between the isotopic compositions of kaolinite and dickite found in sediments indicates that isotopic exchange at low temperatures is relatively slow even on a geologic time scale. The montmorillonite data cannot be interpreted ss simply as the kaohnite data in that all the samples do not appear to have reached isotopic equilibrium under similar conditions. However, all of the data are compatible with the model of isotopic equilibration over a variety of diagenetic temperatures. The isotope data can be interpreted as evidence that some of the glauconites
38
S.
SAVIN and S. EPSTEIN
which originally formed under marine conditions may have subsequently undergone complete isotopic exchange with fresh waters at sedimentary temperatures. There is, however, no case in which such post-formational exchange at sedimentary temperatures can be unequivocally demonstrated to have occurred. Also, no evidence has been found for incomplete isotopic exchange under sedimentary conditions since no points lie beneath the lines on plots of 6D vs. ~30~~. This suggests that exchange under sedimentary conditions may occur only slowly and with great difficulty. Since clay minerals appear to form in isotopic equilibrium with their environment and since their isotopic compositions respond to changes in the environment in a predictable manner at elevated temperatures, isotopic variations should prove useful for determining conditions of formation of clay minerals and alterations during their subsequent history. AchmowZedgemerbts-Samples and associated information were generously provided by Drs. E. GOLDBERG and J. J. GRIFFIN of the University of California at La Jolla, Dr. J. KONTA of Charles University, Prague, Drs. H. A. LO~ENSTAM and R. F. SCOTT.of California Institute of Technology, Dr. J. RICH-S of the Australian National University, and Drs. J. SCHLEE and J. TRTJD~~L of the U.S. Geological Survey. We are grateful for valuable discussions with Drs. A. ALBEE, P. 0. BANKS, R. N. CLAYTON, G. D. GARLICK, J. HOFVER, H. A. LO~~NSTA&I, J. R. G’NEIL, L. T. SILVER and H. P. TAYLOR. Financial support was provided by the National Science Foundation Grant number GA-992 and the Atomic Energy Commission Contract number AT(O4-3)-427. During a part of the time in which this research was performed, one of the authors (S. M. S.) was supported by a National Science Foundation Graduate Fellowship.
REFERENCES CALLAGHAN E. (1948) Endellite deposits in Gardner Mine ridge, Lawrence County, Indiana. Dept. of Conservation, Div. of Geol., Bull. 1. CLAYTON R. N. and EPSTEIN S. (1958) The relationship between O1s/O1a ratios in coexisting quartz carbonate and iron oxides from various geological deposits. J. Beol. 66, 352-373. CLAYTON R. N., FRIEDMAN I., GRAF D. L., MAYEDA T. K., MEENTS W. F. and SHIMP N. F. (1906) The origin of saline formation waters. I. Isotope composition. J. Ceophys. Res. 71,3&X+3882. 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. CLEVELAND G. B. (1960) Geology of the Otay bentonite deposit, San Diego County, California. Calif. Div. Mines Spec. Rept. 84. CRAIG H. (1961a) Standard for reporting concentrations of deuterium and oxygen 18 in natural water. Scieme 133,1833. CRAIG H. (1961b) Isotopic variations in meteoric waters. Science 153,1702-1703. DARTON N. H. (1928) Red beds and associated formations in New Mexico. U.S. BeoE. Swrv. BuuEl.
704. DEGENS E. T. and EPSTEIN S. (1964) Oxygen and carbon isotope ratios in coexisting calcites and dolomites from recent and ancient sediments. Beochim. Cosmochim. Acta 28, 23-44. DEGENS E. T. and EPSTEIN S. (1962) Relationship between O1*/Ors ratios in coexisting carbonates, cherts, and diatom&es. Bull. Amer. Assoc. Petrol. Beol. 46, 634-542. EIV~RY K. 0. (1960) The Sea ofJSouthwn California, 363 pp. John Wiley. EPSTEIN S., BIJCHSBAUMR., LO-NSTAM H. A. and UREY H. C. (1951) Carbonate-water isotopic temperature scale. Beol. Sot. Amer. Bdl. 68, 417-426. EPSTEIN S., SHARP R. P. and Gow A. J. (1965) Six year record of oxygen and hydrogen isotope variations in South Pole flrn. J. Beophys. Res. 70, 1809-1814. FRIEDMAN I. (1953) Deuterium content of natural waters and other substances. GLeochim. Cosmodim. Acta 4, 89-103.
The oxygen and hydrogen isotope geochemistry of clay minerals
39
F~IED~~~~LN I., REDFIELDA. C., SCHOENB. and HARRXSJ. (1964) The variation in the deuterium content of natural waters in the hydrologic oyole. Rev. &ophg8. S, 177-224. G-ox G. D. (1964) Oxygen isotope ratios in coexisting minerals of regionally metamorphosed rooks. Ph.D. Thesis, California Institute of Technology. GARLICKG. D. and D~adoxn J. R. (1967) Oxygen-isotope exchange between volcanic glass and ocean water. (Abstract) Trans. Amer. Ueoplqp. Union 48, 236. GA~LICK G. D. and EPSTEIN S. (1967) Oxygen isotope ratios in coexisting minerals of regionally metamorphosed rooks. Qeochim. Coamochim. Actu 31, 181-214. GODFREYJ. D. (1962) The deuterium content of hydrous minerals from the East-Central Sierra Nevada and Yosemite National Park. Beochim. Coamochim. Acta 26, 1215-1245. GOLDBERUE. D. and GRIFFIN J. J. (1963) Clay-mineral distributions in the Pacific Ocean. In The Sea, (editor M. N. Hill), pp. 728-741. Interscience. HEROLD P. G. and HJZYLG. R. (1942) Kaolin deposits of southern Pike County. Arkaneae Geol. Swv. Bull. ‘7. KEITH M. L. and WEBER J. N. (1964) Carbon and oxygen isotopic composition of selected limestones and fossils. Geochim. Cosmochim. Acta 28, 1787-1816. KERR P. F., -TON P. K., PILL R. J., WHEELER G. V., LEWIS D. R., BURKHA~DT W. RENO D., TAYLOR G. L., Mmrxuz R. C., KINK M. E. and SC~IELTZN. C. (1960) Analytical data on reference clay minerals. API Project 49, Prelim. Rept. No. 7, Amer. Petrol. Inst. KERR P. F., KULP J. L. and HAB~ILTONP. K. (1949) Differential thermal analyses of reference clay mineral specimens. API Project 49, Prelim. Rept. No. 3, Amer. Petrol. Inst. KERR P. F., MAIN M. S. and HAMILTONP. K. (1950) Occurrence and microscopic examination of reference clay mineral specimens. API Project 49, Prelim. Rept. No. 6, Amer. Petrol. Inst. KONTA J. (1966) Der Rohhaolin von Sedlec bei Karlsbad. SIyrec&ruZ Q&3-12. KONTA J. (1961) Crystallization temperatures of clay minerals in the molybdenite and cassiterite Genese et Synthese des Argiles, Colloques wolfram&e ore veins of northern Bohemia. Internationaux de Centre National de la Recherche Scientiflque No. 106, 130-137. KONTA J. (1957) Jilovd ilfineraly &.skosbveneka, 319 pp. Prague. McDona~ I., DUNN P. R., COMPSTON W., WEBB A. W., RICEA~DS J. R. and BOB?NUEXV. M. (1965) Isotopic age determinations on Precambrian rocks of the Carpentaria region, Northern Territory, Australia. J. Qeol. Sot. Au&&a 12, 67-90. MCKINNJEYC. R., MCCREAJ. M., EPSTEINS., ALLEN H. A. and T_JnxyH. C. (1950) Improvements in mass spectrometers for the measurement of small differences in isotopic abundance ratios. Rev. Sci. In&r. 21, 724-730. NEDMANNF. R. (1927) Origin of the Cretaoeous clays of South Carolina. Ewn. Geol. 22,374-387. NLER A. 0. (1947) A mass spectrometer for isotope and gas analysis. Rev. Sci. Irastrum. 18, 398-411. NOTING P. G. (1943) Absorbent clays, their distribution, properties, production, and uses. U.S. cfeol. Suw. Bull. 928-C, 127-219. O’NEIL J. R. and TAYLOR H. P., JR. (1968) Oxygen isotope equilibrium between muscovite and water. In press. OWICNSJ. P., MINARD J. P. and BLAC~ON P. D. (1961) Distribution of clay-sized sediments in the coastal plain formations near Trenton, New Jersey. U.S. aeol. Sm. Prof. Paper 42d-C. 317-319. PERRY E. C., JR. (1968) The oxygen isotope chemistry of ancient oherts. Earth Planet. Sci. Lett. a, 62-66. PLUMB K. A. (1965) Personal communication to S. EPSTEIN. SAVIN S. M. (1967) Oxygen and hydrogen isotope ratios in sedimentary rooks and minerals. Unpublished Ph.D. Thesis, California Institute of Technology. SAVIN S. M. and EPSTEIN S. (1970a) The oxygen and hydrogen isotope geochemistry of ocean sediments and shales. Geochim. Coemochim. Acta 84, 43-63. SAVIN S. M. and EPSTEIN S. (1970b) The oxygen isotopic compositions of coarse grained rooks and minerals. Geochim. Cosmochim. Acta to be published. SIL~~R~~AN5. R. (1951) The isotopic geology of oxygen. &o&m. Coemochim. Acta 2,26-42.
H.
L&EWART
B.,
Jn.
(1962)
Oceanographic
Cruise
Report,
U.S.
Ex$orer-1960, 162 pp. U.S. Dept. (lommerce, Washington. TAYLOR H. P., JR. and EPSTEIN S. (1962) Relationship between minerals of igneous and metamorphic Bull. Beol. Sot. Amer. 73, 461- 450.
rocks.
Part
1. Principles
(‘onst
c1n.d GeotE. Suw.
#Ship
01s/016 ratios ill coexistllq and experimental results.
T~XJMBX~LL J. (1965) Personal communication. VEATCH 0. (1908) Kaolins of the Dry Branch region, Georgia. ,%~?a.Cr’eol.3, 109-117. WEBER J. N. (1965) The 01e/016ratio in ancient oceans. Beokhimi~o 674-680.
APPENDIX-SAMPLE
DESCRIPTIONS
AND GEOLOGICAL
BACKGROUND
A. Kaolinite 1. Murfreesboro, Arkansas (API Standard H 1). Upper Cretaceoue Tokio Fm. Waterlaid volcanic material altered in part after deposition (HEROLD and HEYL, 1942). 601s = +22*4 6D = -5.1. 2. Bath, S. Carolina (API Standard H 5). Cretaceous Upper Hamburg Fm. or Middendorf Fm. (VEATCE, 1908) Depositional environment non-marine or marine (NEUMANN, 1927). Some alteration since deposition (KERR et aE., 1960). 601* = +22-l, SD = -5.6. 3. Langley, 8. Carolina (Wards No. 492). A few miles from Bath, S. Carolina. 60’s = +21.2, 6D = -5.7. 4. Mesa Alta, N. Mexico (API Standard H 9). Cretaceous Dakota Sandstone. Post depositional alteration (DARTON, 1928). 601s = +18.‘7, SD = -7.3. 5. Sedlec, Czechoslovakia (provided by J. KONTA). Residual weathering kaolinite formed on granite. Oligocene and early Miocene age. 80’s = +19*3, 6D = -7.9. 6. Kadan, Czechoslovakia (provided by J. KONTA). Weathering origin, 33 km NE of the Sedlec kaolinite (KONTA, 1957). 60”’ = $19.4, 6D = -7.1. 7. Podlesi, Czechoslovakia (provided by J. KONTA). Similar to Sedlec and Kadan samples and of weathering origin (KONTA, 1957). 6018 = $19.0, SD = -7.2. B . Halloysite
1. Bedford, Indiana (API Standard H 12). Chester formation, possibly deposited from solution (CHAPMAN, SO’s = $22.0, 6D = -6-O.
1948).
C. Dickite 1. Horni Slavkov, Czechoslovakia (provided by J. KONTA). Hydrothermal origin. Temperature of formation estimated at approximately 350% (KONTA, 1957). 601* = +153, SD = -3.6. 2. Kladno Zapotocky, Czeohoslovakia (provided by J. KONTA). From a sedimentary deposit (KONTA, 1957). 601s = +15*6, 6D = -3.7.
300 to
The oxygen and hydrogen isotope geochemistry of clay minerals
41
D. Mc&morillmdte (jrom ocean cora) I. RIS 81 (provided by E. GOLDBERU and J. J. GR~FZ?XN). Pacific Ocean floor (14”08’S, 138OO6’W) probably altered volcanic m&terisl (GOLDBERG and G~FFIN, 1963). Purified by J. J. GRWIN. 60’s = $26.1, 6D = -7.8. 2. EM-7 Run 1 (provided by J. J. GRI~BIN). From experimental Mohole test site 28”58’N, 117”28’W, approximately 163 m below sediment surface. Volcanic fragments are found in nearby sediment and basalt Iayer is 15 m below sample. J. J. GRIFFINsepamted the finest gmined, very pure montmorilloni~ fraction. 601s = +28*5, 6D = -6.8. E. Mot&norzXonite(from bentonites) 1. Polkeville, Mississippi (API Standard H 19). Ohgoeene Vicksburg formation. Highly altered volcanic ash. Masine origin (Kxnx et al., 1950). 6018 = +285, 6D = -6.4. 2. Chambers, Arizona (API Standard H 23). Altered volcanic tuff, non-marine (NUTTING,1943). so’8 = -+X9.1, 6D = -7.9. 3. Otay, California (API Standard H 24). Interbedded with Pliocene masine ssnds. Formed from wemtheredvolcanic ash (CLEvnLAxD, 1989). 601s = +19.0, 6D = -4.1. 4. Clay Spur, Wyoming (API Standard H 20). Highly crystallinemontmorillonitefrom the marine Upper CretaceousMowry formation. Altered volcanic ash. Seoondary alteration present (UN, in KERR et al., 1950). 6018 = -l-17-6, 6D = -99. 5. Little Rock, Arkansas (API Standard H 28). Marine (P) Eocene Wilcox formation (KERRet al., 1950). 6018 = +2&o, 6D = -6.0. 0. Blaok Hills, S. Dakota. The exact source is not known, nor was its mineralogy studied in detail. Bentonites in this area, of the United St&es are generally formed by altemtion of volcmic material. 6018 = + 16.9, 6D = --I26 7. Manito, W~hin~on (API Standard H 33b). Nontronite rather than a bentonite, occurs as fissure fillings in volcanic material (KERR et al., 1950). 6018 = +13.2, 6D = -10.9.
1. Morris, Illinois (API Standsrd H 36). From the non-marine underclay of a Pennsylvanian cyolothem. SOi = $20.5 96D = -51. G.
Glaworaite (from ocean sedimenta) I. Strt 1752 (provided by J-S THUMBS). From a surface sample taken by the Woods Hole Oa~o~&phic ~titution at the inner edge of the Blake Plateau in the Atlantic Ocean f31°28*7’N, 79*29.O’W). This glaueonite may be detrital, origixmtingin Tertiary rocks in the Carohnas (3. Scmmr, personal communication) or it may be forming st the present time (R. PRArT, personal communication). 6018 = +26.3, 6D = -7.4.
42
2. Sta 1957 (provided by JAMES TRUIVIBULL). From a surface sample taken in shallow water between Long Island, New York anti Block Island, Rhode Island (40”21’N, 75”57’W); probably detrital in origin. X-ra) analysis has not indicated the presence of any impurities. 6018 = +21*8 6D = -9.0. 3. Sta 6941 (pro;ided by H. A. LOWENSTAM). From glauconite rich sediments in Santa Monica Bay, California. It may have remained in marine waters since formation or may be detrital from some of the glauconite bearing rocks of the surrounding land area. 601s = 238, 6D = -8-R. H. Gkxuconite (porn New Jersey Coastal Plain povided by JOHN SCHLEE)
Upper Cretaceoua and Early l’ertky
Rock+--
Next four samples were deposited in shallow water marine conditions (OVENS et al., 1961). R-8, Merchantville Fm.- (Upper Cretaceous), Kinkora, N. J. Poorly consolidated, fine grained sandstone. Glauconite is fine sand or silt sized. Analyzed sample may be impure. 601* = +20*3, 6D = -7% R-12, Marshalltown Fm. (Upper Cretaceous), Wallenford, N.J. Similar in appearance to R-8. 601* = $23.0, 6D = -8.0. R-13, Navesink Fm. (Upper Cretaceous), New Egypt, N.J. Similar in appearance to R-8 and R-12 but glauconite is coarse sand sized, bright green. 601* = +23-O, SD = -8.6. R-17 Hornerstown Fm. (Paleocene), north of New Egypt, N.J. Almost entirely coarse glauconite pellets in fine glauconite matrix. The coarse glauconite was analyzed. 60** = +22.5, 6D = -8.5. I.
Ghwmite
(from the Precambrian Crawford Fm. of Au&ral~ovided
@ J. RIU~~ARDS)
All samples were taken from the same area on the west side of the Gulf of Carpentaria. All contain glauconite dated by K/A at 1390 million years (MCDOUGIALLet al., 1966). 1. E53/3/14 601* = +14.6; 6D = -6.3. 2. E53/3/16 601* = $13.7; 6D = -6.2. 3. E63/3/16 Differs from E53/3/14and E63/3/16in that much of the glauconite appears to be altered to a brownish material. The brown grains were hand picked and discarded rtnd the nnaltered glauconite was analyzed. Sol8 = + 18.5; 6D = -4-l. 4. E53/3/18 601* = 1-17.0; 6D = -6.4. 5. E53/3/19 6018 = +16*4; 6D = -6.4.
The oxygen
and hydrogen isotope geochemistry
of clay
minerals*
SAMUEL M. SAYIN? and SAMUEL EPSTEIN Division of Geological Sciences, California Institute of Technology Pasadena, California 91109 (Received 31 October 1968; accepted in revised form
4 July 1969)
&s&t-Oxygen and hydrogen isotope analyses have been made on a variety of clay minerals of sedimentaxy and diagenetic origins. The interlayer water of clay minerals ws~ found to exchange rapidly with atmospheric water. Conditions under which the interlayer water could be removed from the clays without rtffecting the isotopic compositions of their aluminosilicate oxygen and hydrogen were therefore determined, and the interlayer water was routinely removed and discarded prior to isotopic analysis. Approximate fractionation factors for clay mineral-water systems at sedimentary temperatures, inferred from the isotopic compositions of natural samples, are: Oxygen
%&-Hz0 Montmorillonite Kaolinite Glauconite
1.027 1.027 1.026
Hydrogen %in-H,O 0.94 0.97 0.93
Consistency of a relationship between oxygen and hydrogen isotope ratios of kaolinites taken from a variety of areas and having different isotope ratios was demonstrated. This relationship results from exchange of clays with different meteoric waters under conditions in which the fractionation factors are relatively constant. This strongly suggests that the fraction. ation factors are equilibrium ones and that kaolinite forms in isotopic equilibrium with its environment. Deviations of the isotopic composition of montmorillonites from a similar rel&,ionship have been interpreted aa resulting from isotopic exchange at slightly elevated temperatures. There was no clearly demonstrable case of clay minerals undergoing isotopic re-equilibration at sedimentary temperatures, although this is one possible interpretation of the data of some Upper Cretaceous glctuconites. INTRODUCTION
of studies of oxygen isotope ratios in sedimentary carbonates (EPSTEIN DEIJENS and EPSTEIN, 1964; KEITH and WEBER, 1964, and others), have led to a greater understanding of the geochemistry of carbonates. In contrast, there exists only a small body of data pertaining to the stable isotope geochemistry of noncarbonate sediments. SILVERMAN (1951) first showed that oxygen isotope ratios of silicate minerals in sedimentary rocks can be indicative of their modes of formation. For example, detrital components from igneous and metamorphic rocks generally contain less 018 than minerals which have low temperature origins. Prior investigations have also indicated that isotopic exchange can occur in sediments which have not undergone obvious metamorphism. Ancient tie-grained marine A NUMBER
et al., 1961;
* Contribution No. 1540, Division of Geological Sciences, California Institute of Technology. t Present Address: Department of Geology, Case Western Reserve University, Clevelmd, Ohio 44106. 26
26
S. SAVIX
mid
S. ICESTEIN
limestones and cherts often undergo a change of oxygen isotopic composition withont showing any obvious metaKnorphi~features (DEOENSand EPSTEIX,1962). GABLICK and DYHOND(1967) observed that volcanic glass shards in Miocene and oIder ocean bottom sediments had undergone isotopic exchange with sea water. The present series of studies (see also GAVINand EPSTEIN,197Oa, b) consists of investigations of the oxygen and hydrogen isotope geochemistry of silicate minerals which form in sedimentary or dingenetic environments. In particular, it has been of interest to determine whether minerals which are formed under sedimentary conditions are formed in isotopic equilibrium with their environments, and under what conditions these authigenic minerals undergo subsequent isotopic exchange. This information is necessary to evaluate the possibilities of using the isotopic compositions of minerals as aids in understanding some of the problems of sedimentary geology and geochemistry. TERMINOLOCZV AND
NOTATION
All isotope ratios are reported in the 8 notation.
The &values are defined as:
1 C SD = (D/H) sampk _11x100 .
‘01* =
(01s/016) sample (018~016)~td
-
1
x
1000
(D/H) std
Oxygen is reported as per mil deviation from the standard while hydrogen is reported as per cent deviation. In both cases the standard reported throughout this work is Standard mean ocean water (SMUT) , m defined by Crwa ( 196la). The fractionation factors between two compounds or phases containing oxygen or hydrogen are defined by the equations:
uF_g
=
018 ( 0161 A 018 (i0’6 B
D and
0FiA u;?_n = D 0Hts
where A and B denote the compounds or phases. BIKELI1VIINARY GONSIDEIUTIONS A number of factors can affect the isotopic compositions which a clay mineral acquires during sedimentary and diagenetic processes. Among these are: (I) The isotopic composition of the various waters with which the clay mineral may have come in contact during and after its formation; (2) The temperature of the environment at any time during which the clay mineral was subject to isotopic exchange ; and (3) Whether the mineral reached isotopic equilibrium with its env~onment. CRAIG(1961b) has shown that most meteoritic waters, other than those in evaporating basins and hot springs, have isotopic compositions lying near a straight line described by the equation BD = 0.8 601* + 1-O. The isotopic compositions of many ground waters behave according to this relationship (EPSTEINet d., 1966). However,
!L’heoxygen and hydrogen isotope geochemistry of clay minerals
27
CLAYTON etal.(1966) have shown that the isotopic compositions of a number of oil field brines lie to the right (low 8D or high 801s) side of the line. They interpret this to be the result of isotopic exchange of meteoric water with a relatively large amount of marine carbonate rock, a material relatively rich in 0’8, causing the water to be enriched in 01* while maintaining its original D/H ratio. It may be expected that the isotopic compositions of most waters encountered in sedimental processes and some waters encountered in diagenetic processes are meteoric and can be described by the above straight line relationship. Since the 01*/016 and D/H ratios of meteoric water are related and since the isotopic compositions of clay minerals depend on the isotopic composition of water, there should also exist a relationship between the oxygen and hydrogen isotopic impositions of clay minerals. It can be shown that if clay minerals undergo isotopic exchange with a variety of meteoric waters under conditions such that the fractionation factors OL&-HpO and a%ik-EpOare constant (i.e. isotopic equilibrium at constant temperature) then the isotopic com~sitions of the clays will lie along a line with the equation 6D = A 601* + B where
hy A = 6.96 x %%&!
M 9.99
OL&-HpO
and hy
B = 800 x afzy-Hso - 699o(;&+~ - 100.
aclay-H10
Thus, the slope of the line is determined by the ratios of the a’s, and its inbrcept is determined by the absolute values of the a’s. It should be noted that the value of the intercept, B, is extremely insensitive to large variations in the values of the a’s, and thus fractionation factors cannot be calculated with any degree of precision directly from the slope and intercept of a line through a set of experimental data. In the case of eq~ib~um isotopic exchange the effect of temperature on the ~actionation factors may be estimated. Analogy with mineral-water systems for which the oxygen isotope fractionation factors have been experimentally determined indicates that the clay-water oxygen isotope fractionation factors should approach unity with increasing ~rn~rat~e. ~though less information is available with respect to hydrogen isotope fractionation factors, there is evidence (SAVIN, 1967) that these, too, approach unity with increasing temperature in the temperature range of interest. As a result of the temperature dependence of the fractionation factors it may be expected that curves drawn through the isotopic compositions of clay minerals which have exchanged with meteoric waters will lie closer to the line of meteoric water at higher temperatures than at lower ones. These relationships provide a useful working model with which the oxygen and hydrogen isotope vacation in clay minerals may be interpreted. ANALYTICAL TECHNIQUE Oxygen Oxygen for isotopicanalysis was liberated from minerals using fluorine.oxidation techniques similartothose desoribed by TAYLOR and EPSTEIN (1962)andCLAYTON and MAYEDA (1963).
28
S.
SAWN and 8. EPSTEIN
The oxygen was converted to carbon dioxide which was then circulated over hot mercury prior to analysis in order to remove any fluorine or bromine which might inadvertently have escapecl entrapment in the liquid nitrogen traps. The step was instituted to safeguard the mass spcct.rorl, eter from contamination. The isotopic composition of the carbon dioxide was measured on a double collecting mass spcctromcter (NIER, 1947; MCKINNEY rt al., 1950). The reproducibility of the analyses was generally & 0.2 X0 (average deviation) for minerals which do not contain molecular water, 10.3 x0 for clay minerals containing interlayer wat.er, and $:0*7 :“;, for zeolit,os. The measured S-values were related to the SMOW standard through a secondary standard of St. Peter Sandstone which was analyzed with each set of six samples. The sOla value of this sandstone from the same bat!ch used by GARLICK (1964) and GARLICE and EPSTEIT (1967), was taken to bo + lo*9 per mil. Hydrogen. Hydrogen was liberated from hydroxyl bearing minerals by heating in a vacuum to 900°C after preliminary outgassing (see below). Most of the hydrogen was liberated in the form of water, but some was liberated as molecular hydrogen. This was converted to water by reaction with copper oxide which was placed in the furnace next to the sample. The resulting water was quantitatively converted to hydrogen by reaction with hot uranium in the manner described by CRAIG ( 1961 b) and GODFREY ( 1962). Isotopic analyses of the hydrogen were performed on a mass spectrometer of the type described by NIER (1947), and FRIEIXXAN (1953). Reproducibility of the analyses was generally better than 10.3 % (average deviation).
All day minerals contain both hydrogen and oxygen in their silicate frameworks. In addition, many clays contain molecular water which occupies structural positions between the silicate layers, They also contain varying amounts of adsorbed water. Prior to making routine isotopic analyses it was necessary to determine the ease with which the interlayer water exchanged isotopically with the atmosphere and t,o determine the optimum conditions for separating interlayer and adsorbed water from the clay minerals without altering the isotopic compositions of the minerals themselves. Details of experiments made to study the above questions are given by SAVIN (1967). The results are summarized below. 1. The interlayer water of cllays isotopically exchanges almost completely with atmospheric water vapor in a few hours. Thus, the isotopic composition of the interlayer water reflects the isotopic composition of the water or water vapor in which the clay was last stored. Therefore in routine analyses the interlayer water was discarded and the hydrogen and oxygen of the silicate frameworks of the minerals were analyzed. 2. Interlayer water can be removed from clay minerals by heating in a vacuum at temperatures of lOO”-250°C for two hours or more. No isotopic exchange between the water and the hydroxyl hydrogen of clays was noticed when the water was removed in this manner. In routine analysis of hydrogen in clay minerals, samples were outgassed in this manner prior to evolution of the silicate (hydroxyl) hydrogen. 3. Nest of the interlayer water of clay minerals can be removed in 24 hr in a drybox in an atmosphere dried by p&. The small amount of interlayer water remaining in the clay does not normally affect the 6018 value of the aluminosilicate oxygen by more than a few tenths of a permil. Therefore, in the analysis of oxygen, the pretreatment described above w&8 routinely used.
RESULTS AND DISCUSSION
Oxygen and hydrogen isotopic analyses were performed on a number of kaolin&es, montmorillonites, and glauconites from a wide variety of areas. The results of these analyses are listed in Table I. and presented in Figs. Z and 2 where they are compared with the analyses of some igneous and metamorphic rocks and minerals. The measured 01*/016 ratios of the kaolinites and bentonites have been corrected for the
29
The oxygen and hydrogen isotope geochemistry of clay minerals Table 1. Isotopic composition of clay minerals Av. dev.
&s8 (%,)
Av. dev.
No. runs
Langley, So. Carolina Murfreesboro, Arkenses Bath, So. Carolina Mesa Alta. N. Mexico Kadan. Czech. Podlesi, Czech. Sedlec, Czech. b. Halloysite
+21.2 + 22.4 +22.1 + 18.7 + 19.4 + 19.0 + 19.3
0.3 0.3 0.1 0.6 0.0 0.2
2 3 2 2 1 2 3
-6.7 -6.1 -6.6 -7.3 -7.1 -7.2 -7.9
0.1 0.3
Bedford, Indiana o. Dickite
+ 22.0
0.4
2
-6.0
-
1
Kladno Zapotocky, Czech. Horni Slavko, Czech. d. Montmorillonite (bentonite)
+16.6 +16.3
0.6 -
2 1
-3.7 -3.6
-
1 1
Polk&&, Miss. Chambers, Arizona otay, Calif. Clay spur, Wyo. Black Hills, 5. Dakota Little Rook, Arkansas 8. Montmorillonite (no&o&e)
+28.6 +19.1 + 19.0 + 17.6 + 16.9 + 26.0
0.2 0.4 0.4 0.2 0.4 0.0
2 7 4 6 3 2
-6.4 -7.9 -4.1 -9.9 - 12.6 -6.8
0.2 0.3 0.1 0.4 -
Ma&q Wash. f. Montmorillonite (ooean oorea)
+13.2
-
1
-10.9
-
1
RI8 81 EM?‘-RUN g. Illite
t26.1 + 28.6
0.0 1.0
2
-7.8 -6.8
0.8 -
2 1
+ 20.6
0.6
2
-6.1
0.2
2
Sta 1762 (Blake Plateau) Sta 1967 (‘L.I. Sound) Sta 6941 (S. Moniaa Bay) Glauoonite (from rooks)
+ 26.3 t21.8 + 23.8
0.6 0.0 -
6 2 1
-7,4 -9.0 -8.6
0.2 0.1 -
4 3 1
RB-Memhantville Fm R12-Marshalltown Fm R13-Navesink Fm R17-Hornerstown Fm E53/3/14 Crawford Fm Australia E63/3/16 Crawford Fm E63/3/16 Crawford Fm E63/3/18 Crawford Fm E63/3/19 Crawford Fm
+ 20.3 +23.0 + 23.0 + 22.6
0.0 0.6 0.1 0.3
2 3 2 3
-7.8 -8.0 -8.6 -8.6
0.6 0.2 0.2 0.1
3 3 3 2
+ 14.6 +13.7 + 18.6 +17.0 + 16.4
0.3 0.4 0.6 0.2 0.1
2 3 4 2 2
-6.3 -6.2 -4.1 -6.4 -6.4
0.3 0.8 0.2 -
3 4 2 1 1
Sample
6D (%,
No. -
a. Kaolinite
1
Morris. Ill. h. Glaucdnite (from ocean
i.
0.2 0.0 0.1 -
1.0
10 2 1 2 1 2 3
2 3 3 11 1 4
aediients)
contribution of quartz and feldspar, using the mineral percentages given by KERR al. (1950) and KONTA (1965). These corrections generally were between O-1permil and 0.3 permil. As the data of Fig. 1 show, the oxygen in the clay minerals is richer in 018 than that of nearly all the igneous and metamorphic rocks and minerals for et
which data are available. Wherever there are a large number of specimens of a particular type of clay mineral from different localities there is a broad spread in the values of the oxygen isotope ratios. This spread in &values is due to the variety of conditions under which the minerals have formed and exchanged. The hydrogen isotope data likewise show a wide range of &values, but unlike the oxygen isotope data the range of BD values for clay minerals is similar to that for igneous and metamorphic muscovitea and biotites. All the clay minerals are depleted in deuterium with respect to sea water, including a number of samples which
8. SAVIN and S. EPSTEIN
30
+4
Igneousreck?, Meiomorphlc
+6 +8 +I0 +I2 +I4 +I6 +I8 +ZOA2+24+26+2&30 / I I I, I I 1 I I,
I
I
/
t---i
’
I_-_-----,
rock? ’
Metamorphvz muscovlte b-j Kaollnlte Citckl!e
H
Ho!loysite
I
Mcntmorillonite (bentonite) Montmorillonite (ocean cores) Glauconite
I
t ! I:
t1
I
II
,
I It+H--i
!Illk
I o
Taylor 8
b. Gorlick
L
Epstein
1i96.2)
a Epstein
11967)
Fig. 1. The 6CP values (relative to Standard Mean Ocean Water) of clays analyzed in this study compared with pertinent isotopic data of previous workers.
6 -13 -12 -I/( -I?
t
Metamorphic
biotite’
Metamorphic
muscovite
Metamorphic
chlorite
I
-?
D (%)
-B
-7
-7
-T
-4
-2
I
-/
?
I
I
’
1 I Ii
a
/
I
I
Kaolmite
-3
I
4 III H
Dicklte
I
Halloysite Montmorillonite (bentonitel Montmorillanite (ocean cores) Glauconite
t
I
I
I
1
{
H
I
Illde
o Taylor61 Epstein
0966)
Fig. 2. The 8D values (relative to Standard Mean Ocean Water) of clays analyzed in this study compared with pertinent isotopic data of previous workers. appear to have formed in isotopic equilibrium with sea water. Thus, while a$&_np is greater than unity at normal ocean temperatures c&&_~,~ is less than unity.
Kaolinitte group The oxygen and hydrogen isotopic compositions of several kaolin&es, dickites, and a halloysite are listed in Table 1 and are plotted on a graph of 6D vs. 6018 in Fig. 3. All the kaolinite points and the halloysite point lie near a, straight line on this plot while the dickite points fall well off the line. This is consistent with the proposed model in which the isotopic compositions of the kaolinites are determined by
The oxygen and hydrogen isotope geochemistry of clay minerals
31
0
1:
_ 0
q
Koolinite Holloysite
_3
A
Dickite
-4
&
-5 s -6 0 -7 ._ co -6 -9 ; -10
:ii
-II -12 -13 -
”
”
-14 ” ” ” ” ” ” ’ 12 I3 14 15 I6 n I8 I9 20 21 22 23 24 25 26 27 26 29 30 6 o'* ("/ 001
Fig. 3. Graph showing relationship between 6D and &O18values of some keolinitea, dickitaa and a halloyeite. Line L calculated corresponding to a&=,,, = l-0266 and uhy W&&#J = 0.970.
isotopic exchange with meteoric waters under conditions of relatively constant fractionation factors. An estimate of the fractionation factors for the kaolinite-water system has been made by relating the measured isotopic compositions of some of the samples to the best estimates available for the isotopic compositions of the waters in the presence of which they formed. All of the kaolinites analyzed in this work are of residual weathering or sedimentary origin. All have probably formed and been deposited in fresh water environments (see Appendix for geological and petrographic descriptions of the samples). All were colleoted from shallow quarries or outcrops in regions of temperate climate. Hydrogen isotopic compositions of waters from streams and springs in areas near those from which some of the ksolinite samples were collected were published by FRIEDMAN et al. (1964). If it is assumed that the kaolinites attained their present isotopic compositions in the presence of waters of the same isotopic compositions as those measured by FRIEDMAN et al.,under surface temperature conditions, then we are able to estimate the fractionation factors a”&_H,Oand ah&,_uIo. These estimates are shown in Table 2. The weakness of these estimates is apparent. The waters were sampled only once from each location, thereby ignoring any seasonal isotopic variations. They were sampled from streams and springs, and the waters analyzed thus may not represent waters which were in contact with the clay samples. These effects would be expected to be smallest for the samples from the relatively low-lying regions of the southeastern United States, and greatest for the sample from the mountainous region of New Mexico. Finally, the isotopio compositions of the waters at the time during which equilibration occurred may have been significantly different than they are now.
33
S. SAVINand S. EPSTEIT Table
2. Estimate
of the isotopic fructionation factors of tho kaolinit+wator sy&em at sediment,ary temperatures
-5.1
“j- 22.4
-5.1
+
Beth. South Czwolina
-5.6
j21.1
Langley, S. Caroline
--5,7
+21.2
Mess Alta, New Mexico
-7.3
$18.7
Murfreesboro Arkanses
22.4
Red l-t., shrcvcpurt~, La. Ouachita 12.. Monroe, La. Ogoocheo R., Ft. McAllister
-2.7
o.!n;i
1.0271
- 3.0
-5.0
0.978
I.0275
-2.1
-3.9
11.96C
I*0261
-2.1
__ 3.9
0.963
1.0262
I.003
I+0295
c-a
Ogeeoheo . R., Ft. McAllister Ca. Sulfur Spring, Jemez Mtns., Xew Mexico
4.6
-7.6
-10.8
* Hydrogen isotopic compositions of wsters from FF.IEDALU et al. (1984). t Oxygen isotopic oompoeitions of w&em estimated using reletionship SD E 0.8 601* +
1 (CZAIQ.
1981b).
The calculations in Table 2 suggest values of the fractionation factors of a&_H,O 19265 and shy kaol_E,o= 0.970. These are averages of the calculated cc-valuesfor the kaolinites from South Carolina and Arkansas. A line corresponding to these ~a~tio~ation factors has been drawn in Pig. 3. Keeping in mind the assumptions on which the calculations were based the agreement of these values leads us to believe that the fractionation factors are equilibrium ones at surface temperatures. The halloysite point falls quite close to the line through the kaolinite data. It is likely that the aluminosilicate portion of halloysite behaves isotopioally very similarly to kaolinite. KERR et at. (1950) have suggested that this pa~i~ul~ specimen was deposited from solution. If so, it was probably formed from a low temperature solution or else underwent subsequent exchange at low temperatures. The isotopic data of two dickite samples lie well above the line drawn through the kaolinite data. The structural difference between kaolinite and dickite is a very minor one, consisting of a small difference in the way the aluminosilicate layers are stacked, If kaolinite and dickite behave simply with respect to isotope fractionations then factors which might cause the dickite points to lie away from the kaolinite line are: equilibration of the dickite at elevated tempemture, exchange with water of unusual isotopic composition, or incomplete isotopic exchange. One of the dickite samples (Horni Slavkov) is of hydrothermal origin. Its temperature of formation was estimated by KONTA (1957, 1961) to be approxima~ly 3OO’C or 350°C. The second dickite sample (Kl~o) occurs with sulfide minerals in fractures and sedimentary rocks. While KONTA (1957) interprets its origin to be low temperature rather than typically hydrothermal, isotopic evidence indicates that it was formed at temperatures higher than surface temperatures, or was subject after formation, to isotopic exchange at somewhat Ellevated temperatures It appears likely that both dickites underwent final isotopic exchange at similar temperat~s and that these temperatures would be below the ~mperature of formation estimated for the Horni Slnvkov dickite. By analogy with the muscovitewater oxygen isotopic fractionation data of O’NEIL and TAYLOR(1968) the value of 4&ite-n 0 in the temperature range 2OO”C-350% can be estimated to range =
The oxygen and hydrogen isotope geochemistry of clay minerals
33
between l*OOland 1.004.Thus formation of a hydrothermal dickite with a 80is value of j-154 would require the presence of a water whose 6018 value is approximately 13 permil. Another explztnation is that the isotopic compositions of the diokites were established by either partial or complete isotopic re-equilibration with waters of normal b-values at temperatures lower than 200°C. If it is aaeumed that the diokite last equilibrated with water of isotopic composition between - 7.4and 0 permil the temperature of last equilibration can be estimated to range between 40% and 1OO’C. M ontmorillowite
Two montmorillonite samples taken from ocean cores and formed by the aubmarine alteration of volcanic rooks, several bentonites of marine and non-marine origin, and a, nontronite were isotopically analyzed. The hydrogen isotope imposition of these montmor~o~~s range from approximately 8D = -4 to - 10 per cent ctnd the oxygen isotopic compositions range from about 80’* = + 17.5 to + 28G permil. This large range of variation strongly suggests formation or exchange under a wide range of conditions of isotopic environment and temperature. The isotopic data, are hated in Table 1 and have been plotted in Fig. 4 on a graph of 8D vs. 8018.
-5 -6 -7 -6 -9 -IO -I -I2 -13 -14
12 13 14 15 16 I?
I8
19 20 21 22 23 24 25 80’8
26
27 28 29
30
(%o)
Fig. 4. Gr@h showing rektionship between 6D and 6018 values of some montmorillonites from ocean cores and bentonite deposits. Line is e&ulsted using constants u&,~_~~~ = l-0273 and &5,+EIo = O-938.
The line on the diagram has the equation 6D = O-73 6018 - 26. This line was drawn using the fractionation factors &$+n,o = 1.0273 and eQnt_n,O = 0.938, which were estim&~d by taking the evemge of the ~lc~&~ ~&otion&tion factors between ~a. water and the four montmorillonitea richest in 018. It mrtybe noted that the line dra,wn pressesabove the average value of the four marine montmorillonites used to determine this line. This resulti from the fact that the line describing the isotopic compositions of meteoric waters lies a similar distance above the point 3
:i4
S. SAVLNand S. E:psmm
corresponding to the composition of ocean water (SMOW). All four of these samples were taken from marine beds or ocean cores and have similar isotopic compositions, suggesting that they were formed under marine conditions. The values of the fractionation factors estimated for these samples are of comparable magnitude to fractionation factors estimated between other clay minerals and water. The points corresponding to the isotopic compositions of the bentonites with the lower 018/016 ratios lie somewhat above the line drawn using the isotope data for samples richest in On?. The isotopic compositions of these sa,mples of lower 01s content are best understood in the light of the model of isotopic exchange with meteoric waters at temperatures above those at which the O’* rich samples equilibrated. By analogy with the effect of temperature on fractionation factors in other systems (e.g. carbonate-water) it may be inferred that these temperatures need not have been considerably warmer than surface temperatures, probably below 50°C for most samples. Other possible explanations for the isotopic data of these samples include variations of the isotopic compositions of meteoric waters from the ideal straight line relationship, incomplete exchange, and variations in the fractionation factors due to differences in the chemical compositions of the minerals. Most deviations of the isotopic compositions of meteoric waters from the normal straight line relationship lie on the right hand side (low 6D and/or high 6018) of the line (CRAIG, 1961b; CLAYTONetal., 1966). This is in theoppositedirection to the effect observed for the clay minerals, and thus variations in the isotopic compositions of waters cannot satisfactorily explain the observed isotopic compositions of the clay minerals. Chemical analyses and structural formulas of all the bentonites analyzed in the present study have been reported by KERR et al.(1950). There appears to be no correlation between chemical composition and isotopic composition of the minerals, and thus there is no suggestion that the chemical composition is responsible for the low 018/01a and/or high D/H ratios. Incomplete isotopic exchange cannot be ruled out as a possibility for the montmorillonites studied. While it is quite possible that the hydrogen of clay minerals could exchange with water more rapidly than the oxygen, it is unlikely that the reverse be the case. Thus, the points representing the isotopic compositions of clay minerals which formed under marine conditions and subsequently underwent incomplete isotopic exchange with fresh waters might lie below the equilibrium line on a plot of 6D vs. 60 18, but would not lie above. Similarly, the data representing the isotopic composition of a clay equilibrated with a light water and subsequently exchanged incompletely with a heavier water could lie above the equilibrium line but not below. Thus, if marine bentonites having isotopic values which lie above the equilibrium line acquired their isotopic compositions through incomplete isotopic exchange it must have been through a complicated series of events. The fact that the isotopic compositions of meteoric waters plays a major part in determining the isotopic composition of the clay minerals is emphasized in the relationship between the d-values of the clays and the locations from which they were sampled. Those clays with the lowest 01* and deuterium values are from the northern United States (Wyoming, S. Dakota and Washington). Those richest in 01* are from ocean cores and the southeastern United States.
The oxygen and hydrogen ieotope geochemietry of clay minerals
36
In summary the most reasonable interpretation of the montmorillonite data is that the four samples which are richest in 01s formed in isotopic equilibrium with sea water at sedimentary temperatures and did not undergo subsequent isotopic exchange, and that the Ol*-poor samples have undergone isotopic exchange with meteoric waters at temperatures greater than sea bottom temperature.
In the present study only one sample was analyzed. This sample, of non-marine origin, was taken from the underclay of a Pennsylvanian cyclothem at Morris, Illinois. It has a 80’s value of +204 permil and a 6D value of -51 per cent. The sample is thus isotopically not unlike other non-marine clays. Because only one sample was measured, it is not possible to discuss isotopic variations of illites in sediments. Blauwnite The isotopic data for the glauconite samples are listed in Table 1 and shown in Fig. 6. There are two distinct groups of samples evident in this figure, one set richer in 0’s and poorer in deuterium than the other. The O%.ich set of samples consists entirely of glauconites formed during or since the Late Cretaceous, and none have been subject to deep burial (OWHINS et al., 1961; EMERY,1960). The Ols-poor set of samples are from a suite of Precambrian Australian rocks and are thought to have been buried at some time during their history to depths ranging from one thousand to seven thousand feet (PLUMB,1965). The Mesozoic and Cenozoic samples can be considered to be more representative of glauconites forming or exchanging on or near the earth’s surface under conditions similar to those found at the present time.
0 b
-5
-
-8
-
-9
-
A
Recombrim
0
Ocean sediments
rocks
-10 -II
-
-12 -13 -14
’ ’ ’ ’ ’ ’ ’ ’ ’ ’ ’ 12 13 14 15 16 17 19 I9 20 21 22 23 do’*
’ ’ ’ 24 25 26
’ ’ ’ 27 26 29
(%I
Fig. 6. Graph showing relation&p between 6D and 601e values of Borne glaucon&e taken from rocks and ocean eedimente. Line a ie calculated wing conetants ~&+H,o
= l-0263
ami abh$ac_H,O = 0.936. Line b is drawn through Precambrian data parallel to line a.
A line corresponding to the equation 6D = O-72 6018 - 25 llas been drawn OII Fig. 5. This line corresponds to the fractionation factors ~~~Uc_II,O -_ O-926 and ~$L-H20 = 1.0263 which are calculated from the assumption that sample Sta 1752 has undergone complete isotopic exchange with sea water. This sample, the richest, in Ols of all the glauconites mea,sured, was dredged from the surface of the Blake Plateau (31’28*7’N, 72”29*O’W) in the Atlantic Ocean. A temperature of approxi. mately 6°C prevails on the sea floor in this area (STE~AK~, 1962). The isot,opic composition of the water in this area is probably close to 0.0 permil. It is uncertain whether this glauconite is of detrital origin from the continents or is forming in sit,u (TRUMBULL, 1965). With only one exception (R-8), data for the remaining Mesozoic and Cenozoic glauconites lie on or very close to the line drawn, in a manner compatible with the model of exchange with fresh waters under conditions of constant fractionation factors. If this model is a valid interpretation of the isotopic compositions of these minerals, and if it is also true that they have formed under marine conditions, as interpreted by OWENS et al. (1961), then it is the only case where minerals formed in a marine environment appear to have undergone isotopic exchange at sedimentary temperatures. While the above explanation is possible, there is another interpretation of the data. The isotopic compositions of the five samples which lie near the line are clustered together in a small region. It is possible that some or all of these points represent glauconites which were also in isotopic equilibrium with sea water and did not subsequently undergo isotopic exchange, provided that the sea waters with which the samples equilibrated were appropriately diluted with isotopically light fresh water, and/or that the temperatures of formation were warmer than for the Blake Plateau sample. The two facts which raise the suspicion that this may actually be the case are: (1) Th e isotopic compositions of these samples cluster together, and it is unlikely that all of them would have exchanged completely with waters of similar isotopic compositions at similar temperatures: and (2) One of the samples (Sta. 6941) was dredged from Santa Monica Bay, California and has very likely formed in sea water and remained there since formation (EMERY, 1960). Factors which support the proposed mechanism for obtaining the observed isotopic compositions are that all samples were formed in areas which might be diluted by fresh water, and that four of them were formed during the late Cretaceous when temperatures were warmer than at present and when the 6018 and 6D of the oceans could have been lower because of the absence of isotopically light polar ict* caps. The isotopic composition of one Tertiary sample (R-8) lies well above the line drawn. The deviation of this sample from the line can be caused by exchange with fresh water under conditions of elevated temperature. The deviation may also be due to the difficulty encountered in removing 018 poor detrital minerals from this sample. A line of slope similar to that through the data of the younger samples may be drawn through the points representing the Precambrian samples, but is considerably displaced from the other. One possible explanation for this displacement is that the isotopic composition of sea water may have changed since the time of formation of
The oxygen and hydrogen
isotopegeochemistryof clay minerals
37
the samples 1390 million years ago (MCDOUGALL et uZ., 1966). If the isotopic composition of sea water were approximately 3 per cent richer in deuterium and 6 permil poorer in O1s 1390 million years ago than at the present time, then the two sets of data,would be compatible. A change in the oxygen isotopic composition of se&water in this direction has been postulated by WEBER (1965) and further supported by PERRY (1968). However, the isotopic data of the Precambrian glauconites can also be explained by exchange with ground water at an elevated temperature which need not have been above about 100°C. The latter appears to be a likely explanation for at least part of the observed effect since the sandstones from which the samples were taken are cemented with secondary quartz and the rocks were therefore probably subject to invasion by silica-bearing solutions with which isotopic exchange might have occurred. In addition, there is a spread of five permil in the oxygen isotopic compositions of the glauconites indicating that at least some of them underwent exchange after formation. The observed range of O1s/O1sratios is larger than seems reasonable to have occurred in marine waters in a fairly short length of time and in a fairly small area. The K-Ar dates for these samples are the same within experimental error (MCDOUGALL et al., 1965). SUMMARYAND CONCLUSIONS The variations of the oxygen and hydrogen isotope ratios in kaohnite, illite, montmorillonite and glauconite are relatively large. The ratios are dependent on the conditions of formation and exchange of the clay minerals. When a series of clays have undergone isotopic exchange with different meteoric waters under conditions of constant a~&R,o and a&+rlo (i.e. equilibrium exchange at constant temperature) their isotopic compositions show a relationship of the form 8D = A 601s + B. Clays which have exchanged under equilibrium conditions with water at elevated temperatures will have more positive values of 8D and more negative values of 601e. They will therefore lie above the line drawn for sedimentary temperatures. Of all the clay mineral data, the kaolinite data show the most systematic behavior in terms of exchange with different waters under conditions of constant aE,,l_HIo and c&,~_R~~.The near constancy of these parameters strongly suggests that the exchange resulted in isotopic equilibrium. Although kaolinite forms under a wide range of climatic conditions, pure kaolinite deposits such as the ones considered here tend to be formed in zones of quite intense weathering. This weathering is favored by a mild wet climate and so the kaolinite samples measured here were probably formed within a fairly narrow temperature range, thus more closely fulfilling the conditions of the model than other sample suites. The persistence of a lsrge difference between the isotopic compositions of kaolinite and dickite found in sediments indicates that isotopic exchange at low temperatures is relatively slow even on a geologic time scale. The montmorillonite data cannot be interpreted ss simply as the kaohnite data in that all the samples do not appear to have reached isotopic equilibrium under similar conditions. However, all of the data are compatible with the model of isotopic equilibration over a variety of diagenetic temperatures. The isotope data can be interpreted as evidence that some of the glauconites
38
S.
SAVIN and S. EPSTEIN
which originally formed under marine conditions may have subsequently undergone complete isotopic exchange with fresh waters at sedimentary temperatures. There is, however, no case in which such post-formational exchange at sedimentary temperatures can be unequivocally demonstrated to have occurred. Also, no evidence has been found for incomplete isotopic exchange under sedimentary conditions since no points lie beneath the lines on plots of 6D vs. ~30~~. This suggests that exchange under sedimentary conditions may occur only slowly and with great difficulty. Since clay minerals appear to form in isotopic equilibrium with their environment and since their isotopic compositions respond to changes in the environment in a predictable manner at elevated temperatures, isotopic variations should prove useful for determining conditions of formation of clay minerals and alterations during their subsequent history. AchmowZedgemerbts-Samples and associated information were generously provided by Drs. E. GOLDBERG and J. J. GRIFFIN of the University of California at La Jolla, Dr. J. KONTA of Charles University, Prague, Drs. H. A. LO~ENSTAM and R. F. SCOTT.of California Institute of Technology, Dr. J. RICH-S of the Australian National University, and Drs. J. SCHLEE and J. TRTJD~~L of the U.S. Geological Survey. We are grateful for valuable discussions with Drs. A. ALBEE, P. 0. BANKS, R. N. CLAYTON, G. D. GARLICK, J. HOFVER, H. A. LO~~NSTA&I, J. R. G’NEIL, L. T. SILVER and H. P. TAYLOR. Financial support was provided by the National Science Foundation Grant number GA-992 and the Atomic Energy Commission Contract number AT(O4-3)-427. During a part of the time in which this research was performed, one of the authors (S. M. S.) was supported by a National Science Foundation Graduate Fellowship.
REFERENCES CALLAGHAN E. (1948) Endellite deposits in Gardner Mine ridge, Lawrence County, Indiana. Dept. of Conservation, Div. of Geol., Bull. 1. CLAYTON R. N. and EPSTEIN S. (1958) The relationship between O1s/O1a ratios in coexisting quartz carbonate and iron oxides from various geological deposits. J. Beol. 66, 352-373. CLAYTON R. N., FRIEDMAN I., GRAF D. L., MAYEDA T. K., MEENTS W. F. and SHIMP N. F. (1906) The origin of saline formation waters. I. Isotope composition. J. Ceophys. Res. 71,3&X+3882. 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. CLEVELAND G. B. (1960) Geology of the Otay bentonite deposit, San Diego County, California. Calif. Div. Mines Spec. Rept. 84. CRAIG H. (1961a) Standard for reporting concentrations of deuterium and oxygen 18 in natural water. Scieme 133,1833. CRAIG H. (1961b) Isotopic variations in meteoric waters. Science 153,1702-1703. DARTON N. H. (1928) Red beds and associated formations in New Mexico. U.S. BeoE. Swrv. BuuEl.
704. DEGENS E. T. and EPSTEIN S. (1964) Oxygen and carbon isotope ratios in coexisting calcites and dolomites from recent and ancient sediments. Beochim. Cosmochim. Acta 28, 23-44. DEGENS E. T. and EPSTEIN S. (1962) Relationship between O1*/Ors ratios in coexisting carbonates, cherts, and diatom&es. Bull. Amer. Assoc. Petrol. Beol. 46, 634-542. EIV~RY K. 0. (1960) The Sea ofJSouthwn California, 363 pp. John Wiley. EPSTEIN S., BIJCHSBAUMR., LO-NSTAM H. A. and UREY H. C. (1951) Carbonate-water isotopic temperature scale. Beol. Sot. Amer. Bdl. 68, 417-426. EPSTEIN S., SHARP R. P. and Gow A. J. (1965) Six year record of oxygen and hydrogen isotope variations in South Pole flrn. J. Beophys. Res. 70, 1809-1814. FRIEDMAN I. (1953) Deuterium content of natural waters and other substances. GLeochim. Cosmodim. Acta 4, 89-103.
The oxygen and hydrogen isotope geochemistry of clay minerals
39
F~IED~~~~LN I., REDFIELDA. C., SCHOENB. and HARRXSJ. (1964) The variation in the deuterium content of natural waters in the hydrologic oyole. Rev. &ophg8. S, 177-224. G-ox G. D. (1964) Oxygen isotope ratios in coexisting minerals of regionally metamorphosed rooks. Ph.D. Thesis, California Institute of Technology. GARLICKG. D. and D~adoxn J. R. (1967) Oxygen-isotope exchange between volcanic glass and ocean water. (Abstract) Trans. Amer. Ueoplqp. Union 48, 236. GA~LICK G. D. and EPSTEIN S. (1967) Oxygen isotope ratios in coexisting minerals of regionally metamorphosed rooks. Qeochim. Coamochim. Actu 31, 181-214. GODFREYJ. D. (1962) The deuterium content of hydrous minerals from the East-Central Sierra Nevada and Yosemite National Park. Beochim. Coamochim. Acta 26, 1215-1245. GOLDBERUE. D. and GRIFFIN J. J. (1963) Clay-mineral distributions in the Pacific Ocean. In The Sea, (editor M. N. Hill), pp. 728-741. Interscience. HEROLD P. G. and HJZYLG. R. (1942) Kaolin deposits of southern Pike County. Arkaneae Geol. Swv. Bull. ‘7. KEITH M. L. and WEBER J. N. (1964) Carbon and oxygen isotopic composition of selected limestones and fossils. Geochim. Cosmochim. Acta 28, 1787-1816. KERR P. F., -TON P. K., PILL R. J., WHEELER G. V., LEWIS D. R., BURKHA~DT W. RENO D., TAYLOR G. L., Mmrxuz R. C., KINK M. E. and SC~IELTZN. C. (1960) Analytical data on reference clay minerals. API Project 49, Prelim. Rept. No. 7, Amer. Petrol. Inst. KERR P. F., KULP J. L. and HAB~ILTONP. K. (1949) Differential thermal analyses of reference clay mineral specimens. API Project 49, Prelim. Rept. No. 3, Amer. Petrol. Inst. KERR P. F., MAIN M. S. and HAMILTONP. K. (1950) Occurrence and microscopic examination of reference clay mineral specimens. API Project 49, Prelim. Rept. No. 6, Amer. Petrol. Inst. KONTA J. (1966) Der Rohhaolin von Sedlec bei Karlsbad. SIyrec&ruZ Q&3-12. KONTA J. (1961) Crystallization temperatures of clay minerals in the molybdenite and cassiterite Genese et Synthese des Argiles, Colloques wolfram&e ore veins of northern Bohemia. Internationaux de Centre National de la Recherche Scientiflque No. 106, 130-137. KONTA J. (1957) Jilovd ilfineraly &.skosbveneka, 319 pp. Prague. McDona~ I., DUNN P. R., COMPSTON W., WEBB A. W., RICEA~DS J. R. and BOB?NUEXV. M. (1965) Isotopic age determinations on Precambrian rocks of the Carpentaria region, Northern Territory, Australia. J. Qeol. Sot. Au&&a 12, 67-90. MCKINNJEYC. R., MCCREAJ. M., EPSTEINS., ALLEN H. A. and T_JnxyH. C. (1950) Improvements in mass spectrometers for the measurement of small differences in isotopic abundance ratios. Rev. Sci. In&r. 21, 724-730. NEDMANNF. R. (1927) Origin of the Cretaoeous clays of South Carolina. Ewn. Geol. 22,374-387. NLER A. 0. (1947) A mass spectrometer for isotope and gas analysis. Rev. Sci. Irastrum. 18, 398-411. NOTING P. G. (1943) Absorbent clays, their distribution, properties, production, and uses. U.S. cfeol. Suw. Bull. 928-C, 127-219. O’NEIL J. R. and TAYLOR H. P., JR. (1968) Oxygen isotope equilibrium between muscovite and water. In press. OWICNSJ. P., MINARD J. P. and BLAC~ON P. D. (1961) Distribution of clay-sized sediments in the coastal plain formations near Trenton, New Jersey. U.S. aeol. Sm. Prof. Paper 42d-C. 317-319. PERRY E. C., JR. (1968) The oxygen isotope chemistry of ancient oherts. Earth Planet. Sci. Lett. a, 62-66. PLUMB K. A. (1965) Personal communication to S. EPSTEIN. SAVIN S. M. (1967) Oxygen and hydrogen isotope ratios in sedimentary rooks and minerals. Unpublished Ph.D. Thesis, California Institute of Technology. SAVIN S. M. and EPSTEIN S. (1970a) The oxygen and hydrogen isotope geochemistry of ocean sediments and shales. Geochim. Coemochim. Acta 84, 43-63. SAVIN S. M. and EPSTEIN S. (1970b) The oxygen isotopic compositions of coarse grained rooks and minerals. Geochim. Cosmochim. Acta to be published. SIL~~R~~AN5. R. (1951) The isotopic geology of oxygen. &o&m. Coemochim. Acta 2,26-42.
H.
L&EWART
B.,
Jn.
(1962)
Oceanographic
Cruise
Report,
U.S.
Ex$orer-1960, 162 pp. U.S. Dept. (lommerce, Washington. TAYLOR H. P., JR. and EPSTEIN S. (1962) Relationship between minerals of igneous and metamorphic Bull. Beol. Sot. Amer. 73, 461- 450.
rocks.
Part
1. Principles
(‘onst
c1n.d GeotE. Suw.
#Ship
01s/016 ratios ill coexistllq and experimental results.
T~XJMBX~LL J. (1965) Personal communication. VEATCH 0. (1908) Kaolins of the Dry Branch region, Georgia. ,%~?a.Cr’eol.3, 109-117. WEBER J. N. (1965) The 01e/016ratio in ancient oceans. Beokhimi~o 674-680.
APPENDIX-SAMPLE
DESCRIPTIONS
AND GEOLOGICAL
BACKGROUND
A. Kaolinite 1. Murfreesboro, Arkansas (API Standard H 1). Upper Cretaceoue Tokio Fm. Waterlaid volcanic material altered in part after deposition (HEROLD and HEYL, 1942). 601s = +22*4 6D = -5.1. 2. Bath, S. Carolina (API Standard H 5). Cretaceous Upper Hamburg Fm. or Middendorf Fm. (VEATCE, 1908) Depositional environment non-marine or marine (NEUMANN, 1927). Some alteration since deposition (KERR et aE., 1960). 601* = +22-l, SD = -5.6. 3. Langley, 8. Carolina (Wards No. 492). A few miles from Bath, S. Carolina. 60’s = +21.2, 6D = -5.7. 4. Mesa Alta, N. Mexico (API Standard H 9). Cretaceous Dakota Sandstone. Post depositional alteration (DARTON, 1928). 601s = +18.‘7, SD = -7.3. 5. Sedlec, Czechoslovakia (provided by J. KONTA). Residual weathering kaolinite formed on granite. Oligocene and early Miocene age. 80’s = +19*3, 6D = -7.9. 6. Kadan, Czechoslovakia (provided by J. KONTA). Weathering origin, 33 km NE of the Sedlec kaolinite (KONTA, 1957). 60”’ = $19.4, 6D = -7.1. 7. Podlesi, Czechoslovakia (provided by J. KONTA). Similar to Sedlec and Kadan samples and of weathering origin (KONTA, 1957). 6018 = $19.0, SD = -7.2. B . Halloysite
1. Bedford, Indiana (API Standard H 12). Chester formation, possibly deposited from solution (CHAPMAN, SO’s = $22.0, 6D = -6-O.
1948).
C. Dickite 1. Horni Slavkov, Czechoslovakia (provided by J. KONTA). Hydrothermal origin. Temperature of formation estimated at approximately 350% (KONTA, 1957). 601* = +153, SD = -3.6. 2. Kladno Zapotocky, Czeohoslovakia (provided by J. KONTA). From a sedimentary deposit (KONTA, 1957). 601s = +15*6, 6D = -3.7.
300 to
The oxygen and hydrogen isotope geochemistry of clay minerals
41
D. Mc&morillmdte (jrom ocean cora) I. RIS 81 (provided by E. GOLDBERU and J. J. GR~FZ?XN). Pacific Ocean floor (14”08’S, 138OO6’W) probably altered volcanic m&terisl (GOLDBERG and G~FFIN, 1963). Purified by J. J. GRWIN. 60’s = $26.1, 6D = -7.8. 2. EM-7 Run 1 (provided by J. J. GRI~BIN). From experimental Mohole test site 28”58’N, 117”28’W, approximately 163 m below sediment surface. Volcanic fragments are found in nearby sediment and basalt Iayer is 15 m below sample. J. J. GRIFFINsepamted the finest gmined, very pure montmorilloni~ fraction. 601s = +28*5, 6D = -6.8. E. Mot&norzXonite(from bentonites) 1. Polkeville, Mississippi (API Standard H 19). Ohgoeene Vicksburg formation. Highly altered volcanic ash. Masine origin (Kxnx et al., 1950). 6018 = +285, 6D = -6.4. 2. Chambers, Arizona (API Standard H 23). Altered volcanic tuff, non-marine (NUTTING,1943). so’8 = -+X9.1, 6D = -7.9. 3. Otay, California (API Standard H 24). Interbedded with Pliocene masine ssnds. Formed from wemtheredvolcanic ash (CLEvnLAxD, 1989). 601s = +19.0, 6D = -4.1. 4. Clay Spur, Wyoming (API Standard H 20). Highly crystallinemontmorillonitefrom the marine Upper CretaceousMowry formation. Altered volcanic ash. Seoondary alteration present (UN, in KERR et al., 1950). 6018 = -l-17-6, 6D = -99. 5. Little Rock, Arkansas (API Standard H 28). Marine (P) Eocene Wilcox formation (KERRet al., 1950). 6018 = +2&o, 6D = -6.0. 0. Blaok Hills, S. Dakota. The exact source is not known, nor was its mineralogy studied in detail. Bentonites in this area, of the United St&es are generally formed by altemtion of volcmic material. 6018 = + 16.9, 6D = --I26 7. Manito, W~hin~on (API Standard H 33b). Nontronite rather than a bentonite, occurs as fissure fillings in volcanic material (KERR et al., 1950). 6018 = +13.2, 6D = -10.9.
1. Morris, Illinois (API Standsrd H 36). From the non-marine underclay of a Pennsylvanian cyolothem. SOi = $20.5 96D = -51. G.
Glaworaite (from ocean sedimenta) I. Strt 1752 (provided by J-S THUMBS). From a surface sample taken by the Woods Hole Oa~o~&phic ~titution at the inner edge of the Blake Plateau in the Atlantic Ocean f31°28*7’N, 79*29.O’W). This glaueonite may be detrital, origixmtingin Tertiary rocks in the Carohnas (3. Scmmr, personal communication) or it may be forming st the present time (R. PRArT, personal communication). 6018 = +26.3, 6D = -7.4.
42
2. Sta 1957 (provided by JAMES TRUIVIBULL). From a surface sample taken in shallow water between Long Island, New York anti Block Island, Rhode Island (40”21’N, 75”57’W); probably detrital in origin. X-ra) analysis has not indicated the presence of any impurities. 6018 = +21*8 6D = -9.0. 3. Sta 6941 (pro;ided by H. A. LOWENSTAM). From glauconite rich sediments in Santa Monica Bay, California. It may have remained in marine waters since formation or may be detrital from some of the glauconite bearing rocks of the surrounding land area. 601s = 238, 6D = -8-R. H. Gkxuconite (porn New Jersey Coastal Plain povided by JOHN SCHLEE)
Upper Cretaceoua and Early l’ertky
Rock+--
Next four samples were deposited in shallow water marine conditions (OVENS et al., 1961). R-8, Merchantville Fm.- (Upper Cretaceous), Kinkora, N. J. Poorly consolidated, fine grained sandstone. Glauconite is fine sand or silt sized. Analyzed sample may be impure. 601* = +20*3, 6D = -7% R-12, Marshalltown Fm. (Upper Cretaceous), Wallenford, N.J. Similar in appearance to R-8. 601* = $23.0, 6D = -8.0. R-13, Navesink Fm. (Upper Cretaceous), New Egypt, N.J. Similar in appearance to R-8 and R-12 but glauconite is coarse sand sized, bright green. 601* = +23-O, SD = -8.6. R-17 Hornerstown Fm. (Paleocene), north of New Egypt, N.J. Almost entirely coarse glauconite pellets in fine glauconite matrix. The coarse glauconite was analyzed. 60** = +22.5, 6D = -8.5. I.
Ghwmite
(from the Precambrian Crawford Fm. of Au&ral~ovided
@ J. RIU~~ARDS)
All samples were taken from the same area on the west side of the Gulf of Carpentaria. All contain glauconite dated by K/A at 1390 million years (MCDOUGIALLet al., 1966). 1. E53/3/14 601* = +14.6; 6D = -6.3. 2. E53/3/16 601* = $13.7; 6D = -6.2. 3. E63/3/16 Differs from E53/3/14and E63/3/16in that much of the glauconite appears to be altered to a brownish material. The brown grains were hand picked and discarded rtnd the nnaltered glauconite was analyzed. Sol8 = + 18.5; 6D = -4-l. 4. E53/3/18 601* = 1-17.0; 6D = -6.4. 5. E53/3/19 6018 = +16*4; 6D = -6.4.