Quantitative paleoenvironmental and paleoclimatic reconstruction ...

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ARTICLE IN PRESS 6 N.D. Sheldon, N.J. Tabor / Earth-Science Reviews xxx (2009) xxx–xxx 76% of original thickness given independent constraints on burial depth. The primary disadvantages to the approach of Sheldon and Retallack (2001) are that it requires accurate soil taxonomy and that it requires a fairly accurate estimate of the burial overburden. Because the differences in compactibility among different soil orders can be fairly large (Fig. 4), accurate taxonomic identification is very important. Similarly, most of the compaction predicted by Eq. (4) occurs within burial depths of 4 km or less, with very little change among deeply buried paleosols. Thus, uncertainty in burial depth of even 500 m for shallowly buried paleosols would have a large effect on the calculated compaction. Fig. 4. Compactibility of different soil orders. To give one specific example, Retallack (1986) used ptygmatically folded clastic dikes in the ~2.2 Ga Hekpoort paleosol to estimate that it had been compacted to 67–73% of its original thickness. Sheldon and Retallack (2001) used Eq. (4) and constants for a Vertisol (Retallack, 1986; Driese et al., 2005) fromTable 2 to estimate compaction to 72– 2.2.2. Ichnology A number of recent workers have begun integrating continental ichnology with studies of paleosols (e.g., Genise et al., 2004; Kirkland, 2006; Kraus and Hasiotis, 2006; Laza, 2006; Hamer et al., 2007a,b), and trace fossils have been key to understanding some early ecosystems (e.g., Feakes et al., 1989; Retallack and Feakes, 1987). A wide variety of animal and plant traces have been identified (e.g., Hasiotis, 2004) that are useful in paleoenvironmental reconstructions using paleosols. In particular, rhizolith or root trace density (Fig. 1D) and penetration depth have been used to understand vegetation type and density (O'Geen and Busacca, 2001; Retallack, 2007) and to characterize the paleohydrological setting (Kraus and Hasiotis, 2006). In marine settings, a semi-quantitative ichnofabric index has been created (Droser and Bottjer, 1986) that looks at total bioturbation density. Attempts to apply this method to paleosols have been relatively rare and as Genise et al. (2004) point out, well-developed Fig. 5. Bioturbation index versus root size in paleosols. Data are from N.D. Sheldon and J.M.M. Hamer (unpublished). See text for discussion. 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
ARTICLE IN PRESS N.D. Sheldon, N.J. Tabor / Earth-Science Reviews xxx (2009) xxx–xxx 7 paleosols may be nearly devoid of bioturbation and poorly developed paleosols may be very heavily bioturbated, so the definitions appropriate to a marine setting do not necessarily apply to a continental setting. In part, this is due to fluctuating water tables and redox conditions in continental settings. Whereas hydrological and chemical conditions in a marine setting may be relatively uniform over long periods of time, they may fluctuate on very short time scales in continental settings (i.e., seasonally), which affects both the creation and preservation of ichnofabric. Furthermore, bioturbation in paleosols may be the result of either animals or plants, so care is needed to delineate both types. For example, while most animals in continental settings (exceptions are bivalves or gastropods) are unlikely to burrow down past the level of the water table, plant roots are very likely to extend to that depth beneath the surface. An example of an attempt to use ichnology semi-quantitatively rather than qualitatively may be found in Fig. 5 using previously unpublished data collected from Miocene paleosols in Montana formed on finegrained overbank flood deposits. Working independently, one half of the team (N.D. Sheldon and J.M.M. Hamer) measured various root trace parameters (mean diameter, maximum diameter, range of root diameter at a given stratigraphic level) and the other half of the team made semi-quantitative assessments of the degree of bioturbation made by animals throughout a 60 m section (Fig. 5). This approach yielded a couple of interesting results. The highest mean root diameters, maximum root diameter (and by extension maximum plant size), and maximum range of root diameters are associated with moderate degrees of bioturbation, suggesting that while some bioturbation helps plant communities to become established, heavy bioturbation may impede growth of high stature plants, instead favoring low stature or fast-growing plants. The data also skew toward mid-range root diameters in heavily bioturbated paleosols and only small roots (with a small size range) are found in weakly bioturbated paleosols (Fig. 5). While more work is needed to confirm these results and to extend them to other depositional settings, these results suggest that there is considerable scope for integrating semiquantitative as well as qualitative ichnological data with studies of paleosols. 3. Quantitative methods overview In recent years, paleopedology has shifted from a largely qualitative field based on comparisons with modern analogues to an increasingly quantitative endeavor. Some of this change has been a result of applying existing techniques to new materials (Sections 4 and 5 below, clay mineralogy and whole rock geochemistry), but many of the innovations have been the result of applying new techniques to new materials, including thermodynamic modeling of soil formation, isotope geochemistry, and applications of empirical relationships derived from modern soils. Paleoclimatic and environmental properties that may reconstructed using the new proxies include provenance, weathering intensity, mean annual precipitation and temperature during pedogenesis, nutrient fluxes into and out of the paleosols, the atmospheric composition of important gases including CO 2 and O 2 , reconstructed vegetative covering, and paleoaltitude. 4. Clay mineralogy of soils and paleosols The clay size fraction in soils is operationally defined as the population of grains and crystals that are b2 μm equivalent spherical diameter (e.s.d.; e.g., Soil Survey Staff, 1975). In a strict sense, therefore, clay minerals can be any mineral in the soil that is b2 μm e.s.d. In practice, however, the term clay minerals usually refers to finely crystalline minerals that result from pedogenic alteration, and most commonly is applied to phyllosilicates, oxides, and oxyhydroxides. It has long been recognized that soil clay minerals are correlated with climate factors, such as temperature and water availability in the soil, because these factors strongly affect the amount and effectiveness of chemical weathering in the soil profile. The correlation of soil clay mineralogy with climate conditions has been reviewed extensively in numerous studies over the past five decades (e.g., Jackson, 1964; Pedro, 1982; Wilson, 1999), and this contribution makes no attempt to provide a thorough review of this subject. Earlier reviews suggest generally that clay mineralogy follows a weathering pattern, from hot and humid to cool and dry, in the order kaolinite→smectite→vermiculite→chlorite and mixed-layer phyllosilicates →Illite and mica. We provide below a brief summary of methods for clay mineral identification, and the occurrence of kaolinites and smectites in soil and paleosol profiles. This reflects that these two phyllosilicates are among the most abundant minerals that have been identified in paleosols, and therefore also most commonly have undergone quantitative geochemical analysis. 4.1. Occurrence of clay minerals Kandite minerals, a group of 1:1 phyllosilicates, are recognized by a basal d(001) spacing of ~7 Å that expands to 10 Å upon solvation with hydrazine and Dimethylsulfoxide (DMSO; Jackson and Abdel-Kader, 1978; Calvert, 1984). Kaolinite and Halloysite are common pedogenic minerals in soil profiles (Wilson, 1999), although halloysite is not a common constituent of paleosol profiles (Srodon, 1999). Pedogenic kaolinites form in well-drained and acidic soils with moderate silica activity and very low base cation activity in soil solutions (Dixon, 1989). This corresponds generally with soil formation in warm and humid climate. Kaolinite-dominated soil profiles are frequently associated with oxide and oxyhydroxide minerals such as goethite, hematite and gibbsite (e.g., Bárdossy et al., 1977; Tardy et al., 1990a,b). Kaolinite may also occur as a minor constituent in soils dominated by 2:1 phyllosilicates (Nash et al., 1988). Kaolinite is an abundant clay mineral in paleosol profiles that formed in paleotropical sites (e.g., Retallack and Germán-Heins, 1994; Gill and Yemane, 1996; Tabor and Montañez, 2004; Jacobs et al., 2005). Notable dissimilarities between modern pedogenic kaolinites and geological deposits include finer grain size, greater degree of disorder, and small amounts of isomophous substitution of Fe for Al in modern soil kaolinites (e.g., Herbillon et al., 1976). This may indicate that at least some diagenetic transformations occur in kaolinites. Smectites, a large group of 2:1 phyllosilicate minerals, are recognized by a basal d(001) spacing of ~12–14 Å which expands to ~18 Å in response to Mg-saturation and glycerol solvation, or expansion to ~20 Å upon glycol solvation (Moore and Reynolds, 1996). Montmorillonite and bedellite are the most common smectite minerals in soils and paleosols, and both minerals permit a substantial amount of isomorphous substitution in the tetrahedral and octahedral layers (Wilson, 1999; Tabor et al., 2004a,b; Tabor and Montañez, 2005). Pedogenic smectite forms in poorly drained soils characterized by high pH, with high chemical activity of silica and basic cations (Borchardt, 1989). This corresponds generally to intermittently poorly drained environments, including monsoonal and xeric climates characterized by strongly seasonal precipitation. Smectite-dominated soil profiles are frequently associated with hematite and calcite. The clay size fraction in soils and paleosols is seldom composed of a single mineral. Rather, it typically includes mixtures of phyllosilicate, Fe- and Al- (oxyhydr) oxides, quartz and occasionally feldspars that cannot easily be separated from each other. As a result, mineralogical and geochemical characterization of discrete minerals in the soil (or paleosol) clay size fraction is complicated. Three approaches have been used to address complex clay mineral mixtures: (1) Adopt assumptions about the composition of minerals in the mixtures that are not under study in order to calculate an end-member composition for the mineral of interest (e.g., Bird and Chivas, 1988a,b; Lawrence and Rashkes Meaux, 1993), (2) ignore complications that may arise 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
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