Quantitative paleoenvironmental and paleoclimatic reconstruction ...

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36 N.D. Sheldon, N.J. Tabor / Earth-Science Reviews xxx (2009) xxx–xxx
From the perspectives of field-based morphological observations
and laboratory-based mineralogical inspection, goethites that form in
the presence of either 2- or 3-component soil CO 2 mixing are indistinguishable,
and they are only discerned from one another by
determination of Fe(CO 3 )OH δ 13 C values. That is, goethites with Fe
(CO 3 )OH δ 13 C values that are too positive to be explained by twocomponent
soil CO 2 mixing are interpreted to have formed by three
component soil CO 2 mixing which included dissolution of pre-existing
carbonate with very positive δ 13 C values (Hsieh and Yapp, 1999; Tabor
et al., 2004a,b; Tabor and Yapp, 2005a,c).
7.5. δ 18 O and δD of hydroxylated minerals
Pedogenic chemical weathering products are dominantly clay-size
(b2 μm), of which a large proportion are hydroxylated (e.g., Wilson,
1999). The most common hydroxylated pedogenic minerals include
mica-like minerals, vermiculites, smectites, chlorites and interlayered
minerals, kaolinite, halloysite, iron oxyhydroxides, aluminum oxyhydroxides,
and manganese oxyhydroxides (Dixon and Weed, 1989;
Wilson, 1999). If these hydroxylated minerals form in chemical equilibrium
with ambient soil waters, and the oxygen and hydrogen
isotope fractionation factors of these minerals are known (see Table 5
for examples pertinent to this work), then oxygen and hydrogen
isotope analysis of both soil and paleosol hydroxylated minerals may
provide insights into environmental and climatic conditions at the
time of mineral crystallization.
The earliest isotope geochemical studies of soils focused upon the
δ 18 O and δD values of silicates and oxyhydroxides in weathering
profiles that developed upon igneous and metamorphic protolith
(Taylor and Epstein, 1964; Savin and Epstein, 1970a; Lawrence and
Taylor, 1971, 1972). These early studies recognized an upwards
enrichment of bulk silicate 18 O through the soil, as concentrations of
clay-size alteration products become greater (Taylor and Epstein,
1964). These results were interpreted to reflect an increasing contribution
of oxygen from low-temperature, pedogenically formed,
minerals, and destruction of pre-existing igneous and metamorphic
minerals, toward the soil surface. Lawrence and Taylor (1971, 1972)
noted that although the concentration of hydroxylated minerals
increases upwards through the soil profiles, there is little or no profilescale
variation in silicate δD values because the protolith from which
the soils are derived contained little or no primary hydroxylated
minerals (i.e., soil alteration products dominate the δD value). Furthermore,
the δ 18 O and δD values of the soil alteration products
defined an array that roughly paralleled the meteoric water line
(Fig. 21); clay minerals exhibited offset from the meteoric water line,
in order from greatest to smallest: kaolinite, goethite, gibbsite and
amorphous aluminosilicates (Savin and Epstein, 1970a,b; Lawrence
and Taylor, 1971; Yapp, 1987a,b; note that this agrees with the isotope
fractionation equations in Table 5). Based upon the observed relationships,
those early studies hypothesized that δ 18 O and δD values
of soil-formed hydroxylated minerals approach isotope
equilibrium with local meteoric waters at temperatures very near
that of the mean annual surface atmosphere. Subsequent studies
of soil-formed, hydroxylated- mineral δ 18 O and δD values have utilized
the basic observations outlined in those earlier studies in order
to (1) differentiate supergene (low-temperature) from hypogene
Fig. 21. Cross plot of δD versus δ 18 O values. The blue line (upper left) represents the isotope composition of waters that lie upon the meteoric water line as defined by Craig (1961).
Black lines depict 5 °C isotherms between 0 °C (right-most black line) and 35 °C (left-most black line) that represent isotope equilibrium between kaolinite and meteoric water. The
red line is the 40 °C isotherm that depicts isotope equilibrium between kaolinite and meteoric water, and is the “supergene/hypogene” reference line (Sheppard et al., 1969). Hightemperature
hypogene processes result in kaolinite δD versus δ 18 O values that lie to the left, whereas low-temperature supergene processes result in kaolinite δD versus δ 18 O values
that lie to the right, of the supergene/hypogene line. The gray polygon represents the modern surface domain (MSD) for kaolinite (Tabor and Montañez, 2005; see also Yapp, 1993a,
2001a,b; Savin and Hsieh, 1998), which is calculated using oxygen and hydrogen isotope fractionation equations for kaolinite (Table 5) in conjunction with the δD and δ 18 O values of
meteoric precipitation and temperature data from the IAEA data base (Rozanski et al., 1993). The orange polygon represents the Warm Earth Surface Domain (WESD) which
considers global temperatures that are 5 °C warmer-than modern, and an ice-free, Earth. Open circles represent δD and δ 18 O values for kaolinite samples that have been interpreted
to originate from hydrothermal and deep-burial diagenetic processes (Sheppard et al., 1969; Sheppard and Taylor, 1974; Marumo, 1989; Macaulay et al., 1993; Stewart et al., 1994;
Osborne et al., 1994; Decher et al., 1996; Hedenquist et al., 1998; Whelan et al., 1998; Matthews et al., 1999; Harris et al., 2000; Bethke et al., 2000; Parnell et al., 2000, 2004; Uysal
et al., 2000; Marfil et al., 2005; Simeone et al., 2005). Open squares represent δD and δ 18 O values for kaolinite samples that were collected from soil and paleosol profiles (Sheppard,
1977; Hassanipak and Eslinger, 1985; Bird and Chivas, 1989 [Cenozoic samples only]; Lawrence and Rashkes Meaux, 1993; Mizota and Longstaffe, 1996; Boulvais et al., 2000; Girard
et al., 2000, Tabor and Montañez, 2005). For pure kaolinite samples, analytical uncertainty for δD and δ 18 O values are ±3‰ and ±0.2‰, respectively, which corresponds to
uncertainties in paleotemperature estimates of ±3 °C. Several kaolinite δ 18 O values, however, are calculated end-member values from mixtures of kaolinite and some other silicate
(e.g., quartz; Bird and Chivas, 1989; Lawrence and Rashkes Meaux, 1993), and have larger uncertainties (b±1.1‰). With the exception of a single kaolinite from the Oligocene of
Australia (Bird and Chivas, 1989), all pedogenic and paleopedogenic kaolinite samples reside within the MSD or WESD. (For interpretation of the references to colour in this figure
legend, the reader is referred to the web version of this article.)
Please cite this article as: Sheldon, N.D., Tabor, N.J., Quantitative paleoenvironmental and paleoclimatic reconstruction using paleosols, Earth-
Science Reviews (2009), doi:10.1016/j.earscirev.2009.03.004