A paleoclimate interpretation derived from pedogenic clay minerals from the Piedmont Province, Virginia

CHEMICAL GEOLOGY I?/rLUI~G

ISOTOPE GEOSCIENCE

ELSEVIER

Chemical Geology 142 (1997) 201-211

A paleoclimate interpretation derived from pedogenic clay minerals from the Piedmont Province, Virginia W. Crawford Elliott a,*, Samuel M. Savin a, Hailiang Dong b,1, Donald R. Peacor b a Department of Geological Sciences, Case Western Reserve University, Cleveland, OH 44106-7216, USA b Department of Geological Sciences, University of Michigan, C.C. Little Building, Ann Arbor, MI 48106-1063, USA Received 27 February 1997; accepted 20 June 1997

Abstract

Oxygen isotope ratios of pedogenic kaolinites were measured to estimate the ambient climate conditions during the formation of a saprolite in the Piedmont Province (Woodbridge, VA). Kaolinite is the predominant pedogenic clay mineral in this saprolite. It is formed mostly from weathering of muscovite and to a lesser extent from weathering of microcline. A dioctahedral smectite-vermiculite phase formed by weathering muscovite, is present in minor amounts ( < 10%). The 180/560 ratios indicate that most, if not all, the kaolinite formed in a cooler climatic regime in which meteoric waters were more depleted in 580 than most present meteoric waters. It is possible that the pedogenic kaolinite formed shortly after, or during, the Pleistocene glacial stage in North America. This interpretation is consistent with the age of this saprolite derived from 5°Be dating methods. The oxygen isotopic signatures of these kaolinites are not consistent with an older middle or early Miocene age for this saprolite. © 1997 Elsevier Science B.V. Keywords: paleoclimatology; Pleistocene; oxygen; isotopes; kaolinite; Piedmont; Virginia

1. Introduction

Faunal and stable isotopic analyses of fossils from marine carbonates have provided much quantitative knowledge about Cenozoic climates (Savin, 1977; Shackleton, 1982; Miller et al., 1987; Woodruff and Savin, 1991; and others). Information about continental paleoclimates can be inferred from the pres-

* Corresponding author. Present address: Department of Geology, Georgia State University, Atlanta, GA 30303, USA. Fax: + 1 404 651 1376. E-mail: [email protected] J Present address: Department of Geological and Geophysical Sciences, Princeton University, Princeton, NJ 08544, USA.

ence of certain pedogenic minerals (e.g., calcite, siderite, iron oxides and clay minerals), and some microstructures, as well as from l a c / 1 2 C ratios of organic carbon in paleosols (e.g. Birkeland, 1984; Ceding, 1991; Retallack, 1994; Rye et al., 1995; Mora et al., 1996). Quantitative estimates of past temperature, humidity and precipitation on the continents have been inferred from stable hydrogen ( D / H ) , and oxygen (18 O//16 O) isotope ratios of pedogenic clay minerals and iron oxides (Lawrence and Taylor, 1971, Lawrence and Taylor, 1972; Bird and Chivas, 1988; Yapp, 1993; Bird et al., 1993; Giral et al., 1993; Lawrence and Rashkes-Meaux, 1993; Giral, 1994; and others). Stable isotope ratios of clay minerals

0009-2541/97/$17.00 © 1997 Elsevier Science B.V. All rights reserved. PH S 0 0 0 9 - 2 5 4 1 ( 9 7 ) 0 0 0 8 9 - 2

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W. Crawford Elliott et al./ Chemical Geology 142 (1997) 201-211

formed by weathering in soil zones are useful in paleoclimate studies because: (1) they generally form in isotopic equilibrium with ambient waters; (2) the variation in 180/160 and D / H of meteoric waters with climate and geography are known well (Craig, 1961; Dansgaard, 1964; Rozanski et al., 1993); (3) the 180/160 ratios of soil water primarily reflect that of local meteoric waters, although the 180/160 ratios of soil water may be modified to some extent by evaporation and evapo-transpiration, especially in the uppermost part of a soil profile (Allison et al., 1983; Wenner et al., 1991; Hsieh et al., 1993, Hsieh et al., 1997; Savin and Hsieh, 1997; and others); (4) oxygen isotope fractionations between clay minerals and water are known with reasonable reliability (Savin and Lee, 1988; Sheppard and Gilg, 1996); and (5) once formed, clay minerals are resistant to subsequent isotopic exchange in the absence of mineralogic alteration of the aluminosilicate lattice (Savin and Epstein, 1970; Yeh and Savin, 1976; Yeh and Epstein, 1978; Yeh and Eslinger, 1986). Most previous isotopic studies of pedogenic clays have focused on tropical soil profiles (Bird et al., 1993; Giral et al., 1993; Giral, 1994). Here, we present the results of an isotopic and mineralogic study of pedogenic clay minerals (predominantly kaolinite) from an actively forming saprolite developed in a temperate climate within the Piedmont Province of Virginia (Pavich et al., 1989). These results indicate that most, if not all, of the kaolinite formed in climatic regime cooler than today' s, probably during the Pleistocene Epoch.

Organic-rich soil

Loess(?)

~=2-----"-

0cln

Stone line

Argillic Horizon

__-ta ~ - - E -_ Saprolite

/ /.

/

~'% / ~

, , ,oo.

E5 E3 E2

E 0 ~

Fig. 1. Sketch of the Woodbridge Saprolite showing locations of samples E0-E5,

The saprolite shows relict textures of the parent Occoquan Adamellite, the minimum age of emplacement of which is 560 Ma (Seiders et al., 1975). The base of the saprolitic zone was not seen at this site, but observation at nearby localities suggest that it probably extends 10-15 m below the exposed outcrop. A stone line composed of 1 cm pebbles occurs near the top of the argillic horizon. Resistant quartz veins are clearly visible transecting all three horizons. Samples E0, E2, E3, and E5 were collected from this profile at locations shown in Fig. 1.

2. Samples

3. Methods

Samples were collected from an abandoned construction site on the north side of State Route 641, approximately 1.5 km west of Woodbridge High School in Prince William County, VA, about 35 km south of Washington, DC. The pedology, clay mineralogy and chemistry of this profile have been described extensively and denoted Pedon VA3 by Feldman and Zelazny (1995). Clearly recognizable horizons of the weathering profile are shown in Fig. 1. This study focused on the saprolite horizon, which is separated from the overlying argillic horizon by a sharp lithologic break.

3.1. Sample preparation and mineral separations Whole-rock samples (ca. 30 g) were treated to remove non-silicate phases, according to the procedures of Jackson (1979). Carbonates were dissolved with a 1 M Na acetate-acetic acid buffer, pH = 4.5), organic matter and Mn oxides were removed with 30% H202, and ferric oxides and Fe- and Al-containing amorphous phases were removed with a Nacitrate, NaHCO 3, Na dithionite solution. These treatments do not affect the 180/160 ratios of kaolinite or smectite (Eslinger, 1971; Yeh, 1980). The prod-

W. Crawford Elliott et al. / Chemical Geology 142 (1997) 201-211

ucts of these treatments consist primarily of clay minerals, quartz and feldspar. The chemically treated samples were dispersed and separated into sand ( > 50 lxm), silt (20-50 t~m), fine silt-coarse clay (2-20 Ixm) and clay-size ( < 2 ~m) fractions by screening and timed centrifugation. Muscovite was separated from the > 50 ~m fractions by hand picking coarser quartz and feldspar and rolling the remaining mica-rich samples in a polystyrene vial. Mica grains adhere to the vial and can be removed with an artist's brush. The particle size distributions of the suspended, Na-saturated, < 2 Ixm fractions were measured using a NICOMP TM Model 370 particle size analyzer. Most of the < 2 ~m fractions have bimodal size distributions, and further particle size separations at the submicrometer scale using the Sharpies TM Continuous Flow Supercentrifuge were designed to isolate the individual modes. After centrifugation, all fractions were cleaned by dialysis in deionized water for 3 days, or by washing in deionized water. The submieromelzr size separates are predominantly kaolinite ( ~ 90%) but contain minor amounts ( ~ 10%) of a smectite-vermiculite phase. (See Section 4.) Attempts to further concentrate either of those phases by physical means, either further particle size separation or magnetic separation of aqueous suspensions using the Frantz TM high magnetic field gradient separator, were unsuccessful. Chemical purification of the smeetite-vermiculite phase was carried out using a reagent described by Rich (1966) and termed the F-solution (a solution containing NH4F, NH4CI, and 1 N HC1, pH = 3). Aliquots of 0.1-0.2 g were treated with F-solution at room temperature for 48 h. Typically, a sample lost 85% of its weight by this treatment, and residues showed no XRD peaks corresponding to kaolinite.

3.2. Mineralogic and petrologic methods Thin sections for petrographic study, 5 × 7.5 cm, were prepared from epoxy-impregnated samples. Clay minerals were identified using a Philips TM Model 3100 X-ray Diffrac.tometer with monochromatic CuK~ radiation. Oriented mounts were made from K-saturated (1 M KCI), Mg-saturated (1 M MgC12), and Na-saturated (1 M sodium acetate, acetic acid

203

buffer pH = 4.5) fractions. The Mg-saturated samples were solvated in glycerol vapor for 3 - 4 days at 60°C. The Hinckley Index was measured on randomly oriented mounts of the < 2 Ixm kaolinite-rich fractions (Plan~on et al., 1989). The < 2 p~m fractions were treated with n-formamide for identification of haUoysite following the procedure of Churchman et al. (1984). The kaolinite contents of separates were measured using a Seiko TM Model 210 Differential Scanning Calorimeter (DSC) at The Ohio State University. The N 2 flow rate was 50 cm3/min. Heating rates were 50°C/min from 90°C to 200°C and 20°C/min from 200°C to 625°C. A calibration curve was constructed from the integrated area of kaolinite endotherms of a series of mixtures (5-60 wt%) of poorly ordered KGA-2 standard kaolinite and corundum powder. The calibration curve had a regression coefficient (r 2) of 0.99. Samples were diluted with equal weights of corundum powder to enhance the precision of the analyses. Clay mineralogy of one sample (E5) was further studied using electron microscopic techniques. Thin sections of samples impregnated with L.R. White TM resin were prepared normal to relict gneissic layers, so that the (001) lattice fringe image could be identified and measured easily. Representative samples of areas rich in phyllosilicate minerals were identified using optical microscopy and removed from the thin sections, ion-milled and carbon coated. Scanning electron microscopy (SEM) was used to identify areas for transmission electron microscopy (TEM) study. TEM observations were made using a Phillips TM CM 12 instrument fitted with a Kevex Quantum solid state detector (operating voltage 120 kV, beam current 20 jzA, objective aperture 20 Ixm). Most of the lattice fringes were obtained at × 75,000 magnification. To minimize electron beam damage, most observations were made with minimum beam intensity, and focusing was optimized in the areas next to those of interest. A camera length of 770 mm and a selected area aperture of 10 Ixm were used to obtain selected area electron diffraction (SAED) patterns. Quantitative chemical analyses were made in the STEM mode with a beam diameter of 50 ,~ and a scanning area 200 ,~ on a side: The conditions for analysis and data processing are described in Jiang et al. (1990).

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W. Crawford Elliott et al. / Chemical Geology 142 (1997) 201-2 l I

3.3. Isotopic techniques

The 180/160 ratios

of clay minerals were measured using the BrF5 extraction technique of Clayton and Mayeda (1963). Results are expressed in ~-notation as per mil (%0) deviation of the 180/160 ratios from that of Standard Mean Ocean Water (SMOW). The ~180 values of soil waters at this site were not measured. Most of the kaolinites were not analyzed in replicate. Replicate oxygen isotope analyses of clay minerals during the study period yielded a reproducibility better than 0.3 %0 (average deviation).

4. Results

4.1. Petrography and mineralogy The primary framework minerals of the saprolite are quartz, microcline and minor muscovite. Relict gneissic texture is evident without magnification. Quartz and microcline form elongate bands 0.3-1.5 mm wide and constituting 50-75% of the thin section. These bands are composed primarily of aggregates of angular quartz with small amounts of microcline. Microcline has a bimodal size distribution. Coarser microcline grains are tabular with corroded interior cleavage surfaces and grain boundaries, and range up to 2 mm in size. The finer grains (0.1-0.4 mm) occur within quartz aggregates. Channels cutting the quartz/microcline bands are, in some cases, filled with clay and translocated quartz. The bands are separated by argillaceous matrix containing variable amounts of muscovite. Two types of microstructures characterize the argillaceous matrix. A vuggy micaceous groundmass contains minor amounts of quartz, variable amounts of muscovite, and scattered opaque minerals. The micaceous groundmass grades into a vuggy nonmicaceous groundmass, which also occurs adjacent to interior cleavage and grain boundaries of large microcline grains. Euhedral muscovite, in some cases with splayed edges, occurs in variable abundance in the matrix. A few grains of mica have undergone complete pseudomorphic replacement by an unidentified pleochroic phase that has slightly lower relief and less distinct grain boundaries than muscovite.

Fig. 2. (a) Scanning electron microscope image showing kaolinite (kaol) as a weathering product of muscovite (ms) in saprolite E0. (b) Lattice fringe image of kaolinite in the argillaceous matrix of saprolite E0. (c) TEM lattice fringe image showing the formation of halloysite from muscovite in saprolite E0. The (001) d-spacing measured for halloysite is 7.6 .~.

W. Crawford Elliott et al. / Chemical Geology 142 (1997) 201-211

The predominant pedogenic clay mineral is kaolinite, identified in thin section by a platy or tabular habit and first order interference colors. It occurs in both the micaceous and non-micaceous matrix. Petrographic relationships indicate that the kaolinite formed by weathering of muscovite and, to a lesser extent, microcline, and that weathering of microcline to kaolinite predates; the weathering of muscovite to kaolinite. SEM observations indicate typical compositions of the argillaceous matrix in the range: kaolinite, 60%; muscovi~Le, 30%; and quartz and microcline, 10%. XRD aJaalysis of the < 2 Ixm fraction and of subfractions separated from it also indicate that poorly ordered kaolinite (Hinckley Index values between 0.12 and 0.19) is the predominant clay mineral in all samples. DSC analyses show kaolinite abundances between 73% and 96%. SEM observations provide further information about the kaolinite. The ends of the muscovite grains are splayed and altered to clay (Fig. 2a). TEM images show two al~Lerationproducts. One occurs as packets up to 140 A thick at the ends of heavily altered muscovite grains (Fig. 2b). It gives lattice fringes with d0oI of 7 A. AEM analyses indicate a formula of K0.06Mg0.16Fe 0.25A1 3.22Si 4.31 (based on 10 tetrahedral oxygens and 8 OH groups). While the chemical composition and the d-spacing suggest that the predominant ( > 80%) phase in this alteration

205

product is kaolinite, it is enriched in Si and K and depleted in AI relative to stoichiometric kaolinite (AI4Si4010(OH)s). The compositional data require that an additional phase must be present. A second phase associated with the altered mica is halloysite, which occurs as packets of layers (1030 layers thick) directly interlayered with packets of unaltered muscovite (Fig. 2c). It was identified as halloysite on the basis of layer spacings and analytical electron microscopy (AEM) data. SAED patterns and lattice fringe images showed that the layer spacing is 7.6 ,~ (d001 = 7.6 ,~), as consistent only with a member of the kaolinite family of minerals expanded relative to kaolinite (d0o1 = 7.0 ,~). AEM data could not be obtained from the thin packets of layers with d0oI = 7.6 A. However, data largely from such layers and adjacent muscovite layers were consistent with a mixture of a kaolinite family mineral and muscovite. Halloysite was not observed by XRD after treatment with n-formamide, and this is consistent with the small proportions of halloysite observed by TEM only in direct association with partially altered muscovite. In addition to kaolinite, dioctahedral vermiculite and an interstratified mica-vermiculite (2:1 clays) are present together in small amounts (typically < 15%). Vermiculite is identified on the basis of a basal spacing of 12.6 ,~ for the Na-saturated clay, expano

Fig. 3. Lattice fringe image of the F-solution treatment residue of saprolite E5 showing the formation of smectite (16 ,~) from weathering muscovite (10 ,~).

206

W. Crawford Elliott et al. / Chemical Geology 142 (1997) 201-211

sion of the clay when° glycerol-solvated and Mgsaturated at 60°C to 14 A, and collapse to 10 ,~ upon K-saturation at room temperature (Douglas, 1982). The mica-vermiculite is identified on the basis of a basal spacing of 11 ,~ for Na-saturated clay, expansion upon glyocerol-solvation and Mg-saturation at 60°C to 12.6 A and collapse to 10 A upon K-saturation at room temperature (Reynolds, 1980). Further information about the 2:1 layer clay comes from TEM and X-ray examination of the residue of the F-solution treatment of sample E5. This residue is primarily muscovite ( > 80%) with a small proportion of a mineral having a d001 spacing of 14-16 A (Fig. 3). The residue also contains small amounts (10-20%) of Fe-Ti oxides, and 5-10% of an amorphous phase containing Si, A1, K and Fe and minor amounts of oxygen. The amorphous phase may be a product of the treatment with the F-solution. XRD, lattice fringe images and SAED patterns provide no evidence for kaolinite or any other 7 A clay. AEM data indicate a high degree of substitution of F - for O a n d / o r O H - in the 2:1 structure, undoubtedly a consequence of treatment with F-solution. In some analyses, the F - content is comparable to that of OH-. AEM and lattice fringe data suggest that this mineral is a high-charge sm e c t i t e , Ca0.14(Mn0.13Fe0.13A13.98)(A10.58Si7.42020XOH)2. The XRD peak of the glycerol-solvated sample at 14 A collapsed to 10 A upon K-saturation, and it shifted to approximately 10.5-11 ~, upon NH~ saturation by the F-solution. The NH~--saturated residue also exhibited a broad prominent peak from 10.6 to 10.8 ,~. These data suggest that the residue is composed largely of an interstratified vermiculite/mica, or an expanded mica (Reynolds, 1980). This is consistent with the results of XRD analysis of the untreated fractions described above. It is inconsistent with the results of STEM analysis of the residue, which indicate the presence of a high-charge smectite. The discrepancy, however, may be an artifact of experimental procedures. Smectite and vermiculite lie at opposite ends of a compositional continuum, differing in the amount of lattice charge on the 2:1 aluminosilicate layers. We solvated our samples for XRD in glycerol vapor. S.B. Feldman (pers. commun., 1993) pointed out that solvation by dropping glycerol directly on the slide results in a better intercalation complex and a higher d-spacing. A1-

though the effect on this phase of treatment with F-solution is uncertain, XRD data are consistent with the interpretation that dioctahedral vermiculite, and interstratified vermiculite-mica (possibly a hydrated mica) are present in the untreated sample (Douglas, 1982).

4.2. Oxygen isotope ratios All isotopic data are listed in Table 1. The 8180 values of muscovite from the lower three saprolite samples range from 9.6 %o to 10.1%0. We did not analyze any mica from the Occoquan Adamellite. However, the 8180 values of muscovite from the saprolite are high relative to muscovite from typical unaltered plutonic rocks (i.e., 6-8 %0) (Taylor and Epstein, 1961), suggesting that the micas have undergone some post-crystallization oxygen isotopic exchange, as a result of either metamorphism or weathering. The t5 ~80 values of the residues of the F-solution treatments ranged from 9.9 %o to 11.5 %0. However, the significance of these values is uncertain. Oxygen yields were very low (2.5-12.3 lzmol/mg), compared to values of 12-15 ixmol/mg expected for stoichiometric smectites due to exchange of O by F or OH. Oxygen yields are not correlated with 8180 values of the residues. The measured 8180 values are similar to those of muscovite (the dominant 2:1 phase in the residues based on STEM examination). The < 2 ~m fraction of sample E0 was reacted with F-solution prepared using 18O-enriched (8180 + 300 %0) water. Its 8180 value (14.4 %0) was 2.7%0 higher than that of the 0.1-0.5 ~ m fraction of the same saprolite treated with F-solution made with laboratory deionized water. Although the comparison was made between different size fractions of the same sample, the results suggest that no more than a small amount (less than 1%) of isotopic exchange occurred during treatment with F-solution. Some or all of the observed isotopic difference between the fractions may have been the result of substitution of F for OH during treatment. As a result of their uncertain significance, we use the 8~80 values of the residues with caution. The 8180 values of the kaolinite-rich mixtures (kaolinite and 2:1 clay minerals) range from 18.3 %o

41

66

86

129

E5

E3

E2

EO

1.2 0.2

0.41

0.42 0.10

0.6 0.2 <0.1

Mode size

86 14

100

85 15

56 43 1

Mode abundance (~

83

81

91

78

83 73 84

Kaolinite ~'/e~

K, M, M-V K, M-V K, M-V

0.1-0.5 < 0.1

85

83 f

85 f

M K,M,Q,F,M-V,V(tr) K, V, M-V, M 88

K, V-M

M K, V, V-M, M

K, M, M-V

M K, M, Q, F, M-V K, V, V-M

K, M-V, V, M K, M, M-V K, M-V

Mineralogy b

0.5-2.0

> 50 2-5 < 2

< 1

> 50 < 2

0.25-1.0 < 0.25

> 50 2-5 < 2

< 2 1-2 < 0.1

Size range ..... ~,~,, ~ , u j

Mixed clays Residue Mixed clays

Mixed clays Residue

0.19

0.12

Mixed clays Residue Mixed clays Mixed clays Residue Muscovite Mixed clays Residue Mixed clays Residue Muscovite

Mixed clays Mixed clays Mixed clays Residue Muscovite

Phase(s) analyzed c

0.18

0.16

Hinckley index

15.56 10.00 12.65

15.22 3.92

15.55 12.30 15.92 15.20 4.69 15.00 16.85 7.02 15.00 8.27 14.80

18.08 15.69 14.59 2.45 15.00

02 yield (Ixmoi/mg)

19.3 11.5 18.3

19.0 14.3 ~

19.1 9.8 18.8 19.2 11.3 10.1 19.5 10.7 19.1 10,1 9.8

19.9 19.3 18.4 12.6 9.6

Measured o'°t)

20.7

19.6

20.7

21.3

20.4

21.4

19.4

Calculated d kaolinite ~180

a Depths are measured from the top of the saprolite zone. b Mineralogic abbreviations are F = feldspar, K = kaolinite, M = mica, M-V = mica-vermiculite, Q = quartz, v = vermiculite, tr = trace. Mixed clays = kaolinite-rich clay fraction, Residue = product of the treatment of kaolinite-containing mixed clays with F-solution to remove kaolinite. d The 81So value of kaolinite is calculated using mass balance, the amount of kaolinite, a kaolinite 02 content of 17.4 ~ m o l / m g , a non-kaolinite 02 content of 15.0 I~mol/mg, and the measured 8J80 values of the mixed clays and the residue. e The F-solution with which this sample was treated was made with water with a ~ 180 value of + 300%~. f These kaolinite abundances were deternined by XRD. All others were determined by DSC.

Depth ~

Sample

Table 1 Mineralogical and isotopic data

I

t~

b~

-x.

t~

¢

208

W. Crawford Elliott et al. / Chemical Geology 142 (1997) 201-211

to 19.9 %0. The 8180 values of the kaolinite end number of each sample was calculated from the 6180 values of the kaolinite-rich concentrate, the 8180 value of the F-solution residue of the same sample and the amount of kaolinite in the mixture, In spite of the caveat above regarding the 8180 values of the 2:1 phase, the small concentrations of that phase in the kaolinite concentrates increased the uncertainty of 0.2-0.3 %o for the calculated 8180 values of kaolinite end-members. The calculated 6180 values of the kaolinite end-members range from 19.4 %o to 21.3 %0.

5. Interpretation and discussion Kaolinite is the predominant pedogenic clay mineral in the Woodbridge saprolite. Petrographic observations indicate it formed primarily from muscovite and to a lesser extent from microcline. Hinckley Index values indicate a high degree of disorder. Small amounts of halloysite, which may be an intermediate product in the formation of kaolinite from muscovite, were identified only through SEM and STEM observations. S T E M / A E M observations indicate that the dioctahedral smectite-vermiculite is an alteration product of muscovite. Its paragenesis relative to kaolinite could not be ascertained. No data on the isotopic composition of precipitation in the vicinity of Woodbridge, VA are available. The International Atomic Energy Agency (IAEA) and the World Meteorologic Organization (WMO), through the Global Network for Isotopes in Precipitation (GNIP), have developed an extensive database of mean monthly temperature and precipitation 8 JSO values at Hatteras, NC, approximately 440 km (275 miles) southeast of Woodbridge. This database (IAEA, 1996) includes data for 81 months between 1962 and 1976. The same database includes mean monthly temperature (but not isotopic) data for Washington, D.C., near the Woodbridge locality. The data for Hatteras were grouped into those for winter months (December through March), spring and fall months (April, May, October and November) and summer months (June through September). The relationships between precipitation 8180 value and temperature were determined by linear regression through the data for each of these periods

.*2t,Q .~ -4

o

~-6 -8

-10

o

5

10

15

20

25

30

Temperature (°C) Fig. 4. The 8tSo of meteoric waters at Woodbridge, VA, calculated from meteorological and isotopic data for Hatteras, North Carolina, and Washington, DC, as described in text. Each curved trajectory within the shaded band represents the possible conditions under which one of the kaolinites at Woodbridge formed. The shaded band encompasses the conditions of formation of all the kaolinites. The arrows drawn from a winter, a spring/fall and a summer data point illustrate the expected effect on that point of a 4°C cooling (solid arrow) and 4°C warming (dashed arrow). Similar arrows could be drawn for each point on the diagram.

(0.49%o/°C in winter, r 2 = 0 . 4 5 ; 0.11%o/°C in spring, r z -- 0.10; and fall and 0.0%o/°C in summer, r 2 = 0.11). The 6180 value of precipitation for each month at Woodbridge was then estimated from the 6180 value of precipitation at Hatteras, the temperature difference between Hatteras and Washington, D.C. for the same month, and the relationship between temperature and the 6180 value of precipitation for that month. Results of this calculation are plotted in Fig. 4. There is significant scatter in the relationships between 6180 values and temperature at Hatteras for each of the seasons. While this creates uncertainties in the estimation the 6180 value of precipitation at Woodbridge for any particular month, the distribution of calculated 8180 values and the relationship between 8180 value and temperature shown in Fig. 4 probably reflect the range of conditions at Woodbridge over the 81 months for which the Hatteras data are available. Using the kaolinite-water fractionation equation of Savin and Lee (1988) 2 and each measured kaoli-

2 The kaolinite-water oxygen isotope fractionation factor is given as: 1000 In a = ( 0 . 4 2 × 1 0 6 / T 2 ) + ( 1 0 . 6 × 1 0 3 ) / T 15.337, where T is absolute temperature.

W. Crawford Elliott et al. / Chemical Geology 142 (1997) 201-211

nite ~180 value, we calculated a trajectory describing the range of temperature and water 6180 values under which that kaolinite could have formed. These are shown in Fig. 4. The shaded band encompasses the range of all possible conditions of formation of the Woodbridge kaolinites. Although temperature and precipitation 6180 values for some months plot within the shaded band, most plot above the band. That is, the isotopic compositions of the kaolinites are consistent with formation from meteoric waters which are more depleted in 180 a n d / o r cooler than most present meteoric waters. The 1SO/160 ratio of soil waters typically is similar to that of local meteoric water or is enriched in ~s{3, by up to a few per mil by evaporation. Greatest enrichments in temperate climates generally occur in the upper meter of the water column (Darling and Bath, 1988; Wenner et al., 1991; Hsieh et al., 1997). While we cannot rule out the possibility that the Woodbridge kaolinites formed in equilibrium with some modem soil waters precipitated in the 'winter months, the relationships seen in Fig. 4 increase the likelihood that they formed (at least in part, since a pedogenic phase may form over an extended period of time) when temperatures were cooler and the 6180 values of precipitation were lower than at present. It is not possible to calculate with any precision the effect of a climate change on the isotopic composition of precipitation. However, the relationship between temperature mad modem 81SO values provides some guidance in ?lacing limits on the kinds of isotopic effects that might have occurred in the past. Examples for summer, spring/fall, and winter datum points are shown in Fig. 4. In each case, the expected effect of a 4°C warming and cooling are shown. Almost all winter points move closer to the shaded band upon cooling. Spring/fall points also move closer, but by very small amounts. Summer points move closer to the shaded band during warming, but most remain far from the band at climatically realistic temperatures. We therefore conclude that a significant fraction of the Woodbridge kaolinite formed during times that were cooler than the present. This interpretation is consistent with young Pleistocene age estimates for this section of the profile; for example,, the estimate of < 1 Ma from 1°Be analyses (Pavich, 1989). It is not consistent

209

with formation of much of the kaolinite during warmer intervals, for example, the early or middle Miocene, as might be expected from Cleaves (1993) model for the formation of Piedmont saprolites.

6. Conclusions Kaolinite is the predominant pedogenic clay mineral formed from weathering muscovite in the micaceous matrix in the saprolite at Woodbridge, VA. We conclude that it formed at lower temperature a n d / o r from solutions depleted in 180 relative to most present day meteoric waters in the region. It is likely that much of this segment of the Piedmont Province saprolite formed under conditions cooler than present conditions, perhaps during the Pleistocene Epoch. This interpretation is consistent with 1°Be analyses, which indicate a young age ( < 1 Ma) for this saprolite (Pavich, 1989). The oxygen isotopic ratios of kaolinite are not be consistent with an older (i.e. Miocene) age for this saprolite.

Acknowledgements We thank Steven Feldman (VPI and SU) and Milan Pavich for providing much information on the saprolite examined in this study. Jean-Pierre Girard collected these saprolites and shared his insights with us. WCE thanks Professor Bruno Boulang6 and colleagues (University Aix-Marseille, France) for instruction on the rudiments of petrographic examination of soil thin sections. Thin sections were prepared by Spectrum Petrographics (South Jordan, UT). We thank Linda Abel for the 180 analyses and Professor Jerry Bigham and Randy Fowler (The Ohio State University) for making their Differential Scanning Calorimeter available for this study. We thank M.I. Bird, J.R. Lawrence, and J.I. Drever for their comments. This study was supported by National Science Foundation grants EAR91-05112 and EAR94-18790 to S.M.S. CWRU Geology Contribution 203.

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