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