Stable isotope compositions of pedogenic carbonates and soil organic matter in a temperate climate Vertisol with gilgai, southern Russia

Geoderma 136 (2006) 423 – 435 www.elsevier.com/locate/geoderma

Stable isotope compositions of pedogenic carbonates and soil organic matter in a temperate climate Vertisol with gilgai, southern Russia Irina Kovda a , Claudia I. Mora b,⁎, Larry P. Wilding c a Institute of Geography, Staromonetny 29, 109017, Moscow, Russia Department of Earth and Planetary Sciences, University of Tennessee, Knoxville, TN 37996, USA Department of Soil and Crop Sciences, Texas A and M University, College Station, TX 77843, USA

b c

Received 22 February 2005; received in revised form 30 March 2006; accepted 3 April 2006 Available online 15 June 2006

Abstract A Vertisol gilgai complex from the North Caucasus, Russia, was studied to evaluate the climatic and environmental conditions governing pedogenesis. The stable isotope compositions of soil organic matter (SOM) and carbonate pedofeatures in the Vertisol record a complex, but interpretable, pedogenic history. Variability in the isotopic composition of pedogenic carbonate as a function of morphology, position in gilgai microrelief, and soil depth emphasizes the importance of these parameters to interpreting pedogenic conditions. The isotopic compositions and age of SOM record evidence of pedogenesis in an earlier (N 5000 yr BP) environment that was warmer and drier than modern conditions, with a significant component of C4 flora (up to 30%). This portion of the soil history is preserved only in the deepest portions of the microlow (N140 cm) or in the central portion of the microhigh chimney. A change to cooler and wetter conditions was accompanied by a shift to an ecosystem dominated by C3 vegetation (N 95%) and resulted in a wetter soil environment, particularly in gilgai microlow positions. Most pedogenic carbonate preserves isotopic compositions reflecting only the more recent climate history. The wetter environment contributed to recrystallization of pedogenic carbonate nodules and formation of soft carbonate masses. There are systematic differences in the isotopic composition of hard nodules and soft masses that indicate the influence of greater soil water evaporation, a greater proportion of xerophytic plants, and/or lower rates of soil respiration in the gilgai microhighs. Self-consistent interpretation of the pedogenic and paleoclimate history of these Vertisols requires consideration of a full suite of information, including carbonate pedofeature micromorphology, the stable isotope composition and age (absolute and relative) of these features, and other soil characteristics. © 2006 Elsevier B.V. All rights reserved. Keywords: Vertisol; Pedogenic carbonates; Stable isotopes; Soil organic matter; Russia

1. Introduction Carbonate pedofeatures often occur in arid and semiarid Vertisols, including soft masses and hard nodules of ⁎ Corresponding author. Fax: +1 865 974 2366. E-mail addresses: [email protected] (I. Kovda), [email protected] (C.I. Mora), [email protected] (L.P. Wilding). 0016-7061/$ - see front matter © 2006 Elsevier B.V. All rights reserved. doi:10.1016/j.geoderma.2006.04.016

various morphology and microfabric (Blokhuis et al., 1990; Mermut and Dasog, 1996; Coulombe et al., 1996). The main factors determining carbonate morphology are the length of time the crystallization surface is moist or wet, rate of precipitation, ionic strength of the soil solution, and crystal surface interactions with organic and inorganic particulate matter (Chadwick et al., 1988). Pedogenic processes in Vertisols, such as shrink-

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swell and lateral shearing, result in a bowl-like subsurface structure and high- and low- microrelief (i.e., gilgai) of Vertisol soil cover (Wilding et al., 1990; Kovda et al., 1992). Spatial and morphological variability in the distribution of soft and hard carbonate pedofeatures between Vertisol microhigh and microlow topographic positions has been observed (Thompson and Beckmann, 1982; Wilding et al., 1990; Kovda et al., 1996a, Wilding et al., 2002). The variety of carbonate pedofeatures coexisting in Vertisols with gilgai may thus provide an avenue to examine pedogenic processes and soil evolution, in particular, whether they are polygenetic (Ryskov et al., 1996). A recent paper by Kovda et al. (2003) discussed aspects of pedogenic carbonate features in a temperate climate Vertisol gilgai soil complex that is further studied here. Carbonate pedofeatures include hard nodules, soft masses and soft masses with a hard carbonate core. Based on macro- and micromorphological observations, knowledge of the modern soil hydrology, and radiocarbon ages of nodules across the gilgai complex, a polygenetic formation was suggested for the carbonate pedofeatures (Kovda et al., 2003). An evolutionary model was hypothesized, from older to younger carbonate pedofeatures, where hard nodules represent an early pedogenic product initiated before or at the early stages of gilgai formation, soft masses represent more recent pedogenic material, and soft masses with hard cores are intermediate in formation between hard nodules and undifferentiated soft masses. Coexistence of carbonates of varying age and genesis has been established for several Vertisols (Rajan et al., 1972; Mermut and Dasog, 1986; Blokhuis, 1993; Pal et al., 2001).

Stable isotope analysis is a powerful tool for discerning different generations of pedogenic carbonate and estimating the paleoenvironmental conditions under which they formed (e.g., Amundson et al., 1989; Quade et al., 1989; Cerling et al., 1993; Mora et al., 1993, 1996; Nordt et al., 1998; Mora et al., 2002). It is also useful for examining carbon cycling, dynamics and behavior of soil organic matter (e.g., Mariotti, 1991). The goal of this paper is to use stable carbon and oxygen isotopic compositions of carbonate pedofeatures, combined with morphological, microfabric and SEM analyses, to discern soil processes governing pedogenesis within different gilgai domains and to reconstruct climatic/environmental conditions governing pedogenesis. Historically, Vertisols were considered soils form-ed in tropical and subtropical climate with strong seasonal differentiation of precipitation (Dudal, 1965), however, Vertisols in regions outside the tropical and subtropical belts (ex. Eastern Europe and Russia) are now estimated to com-prise about 32 million ha, or ∼10%, of known Vertisol area (Dudal and Eswaran, 1988). The results of this study help to elucidate the history of Vertisols formed under temperate climate, including how long they have been forming and whether they reflect soils predominantly formed under war-mer and/or more arid paleoenvironments. 2. Materials and methods 2.1. Study area Vertisols were described and sampled in southern Russia (44° 38′ 17″ N.L., 42° 15′ 04″ E.L.), in the North Caucasus (Fig. 1). The site was located on a large, nearly level,

Fig. 1. Location of the study site (star): Stavropol upland, North Caucasus, South Russia.

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erosional scarp with a gentle backslope, at an elevation about 470 m a.s.l. This is an area of temperate, continental, semiarid climate with mean annual temperature of 9 °C and mean annual precipitation about 500 mm. Soils are formed on Neogene marine clays under native steppe vegetation. 2.2. Soils Vertisols in the study area have normal gilgai microrelief with a surface amplitude of ∼35 cm. Significant variability in soil morphological, physical, chemical and biological properties accompany variability in hydrology across the microrelief (Kovda et al., 1992, 1996a,b, 1999). If each of the microrelief components are considered part of a gilgai soil complex, then according to Soil Taxonomy, the soils investigated on the microhighs would be classified as very fine, mixed, active, mesic, Sodic Haplusterts (similar to Chromic Haplusterts but with sodic conditions); on the microslopes as very fine, mixed, superactive, mesic, Sodic Haplusterts; and on microlows as very fine, mixed, superactive, mesic Typic Epiaquerts (Soil Survey Staff, 1999). Soil horizonation across gilgai microrelief, morphological forms of carbonates and their location are shown in Fig. 2. Radiocarbon ages of humic acids in the Vertisol gilgai complex range from modern in the A2 horizon of the microlow to ∼3200 yr to 5600 yr in the Bkss2 horizon of the microlow and microhigh, respectively (Kovda et al., 2005).

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2.3. Methods 2.3.1. Field methods Soils in a 2 m-deep and 5 m-long trench across the gilgai microrelief were described using routine vertical profile methods (Wilding et al., 2002) and additional information was obtained using the polygon method (Lynn and Williams, 1992). Bulk samples were collected by horizon for physical, chemical, isotopic and microfabric analyses. Carbonate pedofeatures were collected from different gilgai positions for microscopy, SEM and stable isotope analysis. 2.3.2. Stable isotopic analyses Stable isotopic compositions of carbon and oxygen were determined in carbonate nodules and soft masses from the Vertisol microlow, microhigh and microslope positions. Selected large nodules were cut, sampled with a microdrill, and isotopic composition was determined across the pedofeature. For soft masses with hard cores, samples of both core and surrounding carbonate matrix material were analyzed. Samples were ground to powder and heated for 3 h at 375 °C in order to remove volatile organic matter and 1–15 mg of sample was reacted with 100% orthophosphoric acid at 25 °C. Isotopic ratios of carbon and oxygen were measured on the CO2 produced by this reaction, using a Finnigan-MAT Delta plus mass spectrometer and are reported in standard δ-permil

Fig. 2. Morphological forms of carbonates and their distribution in the Vertisol gilgai soil complex. The numbers in the legend indicate: 1 — small (b6 mm diameter) hard nodules ; 2 — large (10–15 mm) hard nodules; 3 — aggregations of small hard nodules; 4 — undifferentiated soft masses; 5 — soft masses with hard cores; 6 — diffuse, undifferentiated soft masses; and 7 — depth to soil effervescence. Vertical lines mark the position of sampled pedons in the microlow, microslope and microhigh. Soil horizons are indicated according to Soil Taxonomy (Soil Survey Staff, 1999). Heavy lines across the gilgai complex indicate major vertical zones of differential pedogenic development, from top to bottom: (1) the dark A and AB horizons with maximal organic matter accumulation; (2) the Bk and Bssk horizons with major mixing of light and dark soil materials (high microzonality) just above the master slickenside; and (3) major slickenside development in the Bssk and Bss horizons with less intensively weathered Bssy horizons at depth.

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notation relative to PDB, with external precision of ± 0.1‰ (2σ) for both carbon and oxygen. Carbon isotope compositions of soil organic matter (SOM) were determined for bulk samples of soil matrix in the microhigh and microlow, collected at 10 cm intervals to a depth of ∼110–130 cm. Representative samples were also collected from depths of 130–190 cm. Additional samples were taken close to the microslope, and ∼50 cm

outside of the central part of the microhigh. Organic remnants of roots and debris were hand picked from soil samples prior to analysis. Dried soil samples were powdered and reacted with 10% HCl to remove inorganic carbon, washed with deionized water and centrifuged until neutral reaction, dried and powdered again. Soil samples (20–250 mg) were loaded in quartz tubes with CuO, pure Cu and Pt wire, combusted at 800 °C for 3 h, and the

Fig. 3. Thin-section photomicrographs of carbonate pedofeatures (plane-polarized light); scale bar is 1 mm. (A) Fragment of a large nodule including part of the central microsparitic aggregate coalescent with thick microsparitic joint cement, from microlow at depth of 100–120 cm; (B) Hard microsparitic nodule with thin micritic rim around embedded into the matrix, from microhigh at depth of 17–22 cm.; (C) Soft mass with internal hard cores. Soft mass exteriors strongly impregnated and aggregated into rounded pellets, from microhigh, in dark material adjacent to the chimney, at depth of 55 cm.

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evolved CO2 was collected, cryogenically purified, and analyzed for isotope ratios as described above. 3. Results 3.1. Field morphology and carbonate pedofeatures The three pedons analyzed show strong spatial variability of morphological attributes as a function of gilgai topographic position. Soils have clear horizonation, which is interrupted in the microhigh by a chimney (i.e., a diapiric protrusion of subjacent soil material that penetrates into the overlying horizons and approaches the surface; Coulombe et al., 1996). The dark, organicrich layer comprising the A- and B-horizons of typical Vertisols thins from ∼ 130 cm in the microlow to ∼ 40 cm over the central part of the microhigh chimney (Fig. 2). Carbonates are strongly leached in the microlow position, where soil effervescence of the slightly calcareous soil matrix starts at a depth of ∼90 cm. Hard carbonate nodules (2–3 mm in diameter) appear in the microlow at a depth of 90–100 cm. (Fig. 2) and are more abundant and larger (to 10–15 mm) at a depth of 109–125 cm. Hard nodules consist primarily of smaller microsparitic nodules coalesced into large nodules and aggregates by thick layers of microsparitic cements (Fig. 3A). The degree of nodule

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coalescence and aggregation is much smaller in the microhigh, in which hard nodules consist of a relatively large, older, microsparitic core surrounded by only a thin, micritic layer (Fig. 3B). A few powdery, soft, carbonate masses (1–2 cm in diameter) and a few with 1–3 mm hard cores, occur in the calcareous soil matrix at depths of 140–160 cm. Effervescence in the microhighs starts at the soil surface. Small, hard (1–6 mm) carbonate nodules appear at 40 cm. They occur mostly outside the chimney, in the dark soil matrix intermediate between the microhigh and microslope. Abundant white soft powdery masses, up to 3–5 cm in diameter, occur mainly within the chimney, but also in the black matrix adjacent to the chimney, where the carbonate segregations are more gray in color and many have hard cores (Fig. 3C). A detailed examination of carbonate micromorphology in this Vertisol was presented elsewhere (Kovda et al., 2003). 3.2. Stable isotopic composition of SOM The δ13C values of SOM in the microhigh and microlow gilgai elements ranged from −26.7‰ to −26.2‰ in surface horizons to −22.8‰ at the deepest parts of the pedons (Fig. 4). Modern roots at a depth of 0–10 cm had δ13C values of −27.3‰ in the microhigh and −26.1‰ in

Fig. 4. Stable isotope compositions of SOM (closed symbols) in different microtopographic positions within the Vertisol gilgai complex: microhigh (within chimney, outside of chimney) and microlow. Calculated proportion of C3 vegetation in ecosystem during pedogenesis and formation of SOM (open symbols).

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the microlow. Isotopic compositions of SOM in the microhigh and microlow increase, in parallel, by ∼1.5‰ to a depth of ∼140 cm (Fig. 4). Below this, SOM values in the microlow are markedly more positive. In the microhigh, SOM values increase only below 190 cm depth, where they merge in value with microlow SOM compositions. Isotopic compositions of SOM collected just outside the microhigh chimney closely follow compositions within the chimney, even though the soil morphology (i.e., color and structure) is more similar to materials in the microlow. 3.3. Stable isotopic composition of pedogenic carbonates The carbon and oxygen isotope compositions of soil carbonates vary depending upon gilgai position (Fig. 5A) and carbonate morphology (Fig. 5B) and, less clearly, as a function of the depth of carbonate occurrence (Fig. 6A– D). The overall range of isotopic compositions over the gilgai soil complex varies by ∼2.3‰ for carbonate carbon (−12.7 to −10.4‰ PDB) and oxygen (−9.9 to −7.6‰ PDB) (Fig. 5A), with positive correlation between

the isotope values. Carbonate occurring in either the microhigh or microlow gilgai elements show nearly the whole range of carbon isotope compositions, varying over 1.8‰ in the microhigh and 1.9‰ in the microlow (Fig. 5A); oxygen isotope compositions are slightly less variable (1.5‰ microhigh, 1.5‰ microlow). Oxygen isotope compositions are enriched in the microhigh positions, relative to the microlow (δ18O = − 9.1 to −7.6‰ microhigh, −9.9 to −8.4‰ microlow). Carbon isotopes compositions show only a slight dependence on position in the gilgai complex, with δ13C = −12.3‰ to −10.4‰ (microhigh) and −12.6‰ to −10.7‰ (microlow). In each case, isotopic compositions in the microslope positions are intermediate in value between compositions noted in the microlow and microhigh. More clear distinctions can be drawn between isotopic compositions of pedogenic carbonate having different morphology, i.e., nodules, soft masses, or soft masses with hard cores (Fig. 5B). Carbonate nodules have the most negative carbon and oxygen isotopic compositions, with δ13C = −12.6‰ to −12.2‰ and δ18O = −9.9‰ to −8.3‰. Soft masses are more enriched, with δ13C = −12.0‰ to −10.4‰ and δ18O = −8.7‰ to −7.6‰. Hard

Fig. 5. (A) Stable isotope compositions for pedogenic carbonates according to their microtopographic position, i.e., in microhigh, microslope and microlow; (B) Stable isotope compositions for pedogenic carbonates according to their morphology, i.e., soft masses, hard cores and hard nodules.

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Fig. 6. (A) Carbon isotope compositions of pedogenic carbonate versus depth in the soil microhigh; (B) Oxygen isotope compositions of pedogenic carbonate versus depth in the soil microhigh; (C) Carbon isotope compositions of pedogenic carbonate versus depth in the soil microlow; (D) Oxygen isotope compositions of pedogenic carbonate versus depth in the soil microlow.

cores within soft masses have both oxygen and carbon isotope compositions that are intermediate between those of nodules and soft masses. Soft carbonate masses show much greater variability in carbon isotope compositions (total range of 1.6‰) compared to nodules (0.5‰) and hard cores within soft masses (0.7‰). For oxygen isotope compositions, the nodules are more variable (total range 1.6‰) than soft masses (1.2‰) or hard cores (0.7‰). The overall isotopic variability of a given morphology within the gilgai complex is greater than the variability observed across individual carbonate nodules or between hard cores and their surrounding soft crystic matrix. In the latter cases, isotopic compositions vary by up to 0.4‰ for carbon and 1.0‰ for O, although individual features in most samples show b 0.6‰ variability in δ18O values (Table 1). Most often, oxygen and carbon isotope compositions are relatively enriched in the outer portions of the nodule or the soft crystic mass surrounding a harder nodule, however, occasional exceptions occurred. The range of compositions observed across gilgai positions, carbonate morphology, and even across individual carbonate features, suggest the variability of factors and

mechanisms in effect during carbonate formation, which are discussed in the following section. The relationship between the isotopic compositions of carbonate features and their depth of occurrence within a specified part of the gilgai complex are shown in Fig. 6A– D. The overall depths trends in both the microhigh and microlow are less compelling than the clear differences in the compositions of the various types of carbonates within a given depth range. For example, carbon isotope compositions clearly increase between 110 and 160 cm depth in the microlow (Fig. 6C). More notably, however, nodule compositions are clearly distinct from the more enriched values of hard cores and soft masses. Isotopic compositions of soft masses are more variable in the microhigh than the microlow. 4. Discussion 4.1. A SOM isotope proxy record of climate and ecosystem evolution The SOM in gilgai microhigh and microlow positions has δ13C values that increase with depth. The

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Table 1 Stable carbon and oxygen isotope compositions of carbonate pedofeatures in a temperate climate Vertisol with gilgai Sample

Carbonate morphology

Description

G2-3

Soft mass with cores

External crystic matrix

G2-4

Soft mass with cores

External crystic matrix

G2-1

Soft mass with cores

Central core, diam. = 1–3 mm

G2-2

Soft mass with cores

Central core, diam. = 4 mm

G4-1

Soft mass with cores

External crystic matrix

G4-2

Soft mass with cores

Central core, diam. = 1–2 mm

G26-3

Small nodule

External layer

G26-2

Small nodule

Central part

G14-1 G14-2 G14-3 G14-4 G28-2 G28-1 G20-2 G20-3 G20-4 G20-5 G21-4 G21-5 G21-3

Gray soft mass with cores Gray soft mass with cores Gray soft mass with cores Gray soft mass with cores Soft mass with cores Soft mass with cores Large nodule Large nodule Large nodule Large nodule Soft mass with cores Soft mass with cores Soft mass with cores

External crystic matrix Central core, diam. = 1–2 mm Central core, diam. = 3 mm Central core, diam. = 4 mm External crystic matrix Central core, diam. = 1–3 mm External layer Central part Crystals from septaric void Crystals from septaric void External crystic matrix Central core, diam. = 4 mm Central core, diam. = 2 mm

radiocarbon ages of humic acids in these pedons also increase with depth, ranging from modern to 3170 ± 110 yr BP (depth of 160 cm) in the microlow and 5610 ± 110 yr BP (depth of 1 m) in the microhigh (Kovda et al., 2005). Commonly invoked explanations for changes in the isotopic compositions of SOM include vegetation changes in the soil ecosystem, isotopic fractionation of SOM during decomposition by microorganisms, or differential mineralization or humification of SOM (O'Brien and Stout, 1978; Stout et al., 1981; Martin et al., 1990; Balesdent, 1991; Cerling et al., 1993; Nordt et al., 1998). In this study, the SOM isotopic compositions are closely paralleled by isotopic changes in pedogenic carbonate, suggesting an ecosystem influence on isotopic compositions. The soil ecosystem, dominated by vegetation, can be approximated as a binary mixture of C3 (including most woody species and cool season grasses, having an average δ13C of − 27‰) and C4 (primarily warm season grasses favored by warm, water-stressed conditions; average δ13C = − 12‰) vege-

Position, depth

Microhigh, outside chimney, 70 cm Microhigh, outside chimney, 70 cm Microhigh, outside chimney, 70 cm Microhigh, outside chimney, 70 cm Microhigh, outside chimney, 70 cm Microhigh, outside chimney, 70 cm Microhigh, outside chimney, 70 cm Microhigh, outside chimney, 70 cm Microslope, 119 cm Microslope, 119 cm Microslope, 119 cm Microslope, 119 cm Microlow, 130–140 cm Microlow, 130–140 cm Microlow Microlow Microlow Microlow Microlow Microlow Microlow

δ13C (PDB)

δ18O (PDB)

− 11.96

− 8.44

− 11.84

− 8.03

− 11.91

− 8.60

− 12.26

− 8.64

−11.34

− 8.26

− 11.56

− 8.37

− 12.19

− 8.82

− 12.55

− 9.12

− 11.94 − 11.93 − 12.17 − 12.21 − 11.71 − 11.59 − 12.24 − 12.32 −12.62 − 12.43 − 11.65 − 11.65 − 11.72

− 8.63 − 8.62 − 8.68 − 8.47 − 8.72 − 8.60 − 9.28 − 9.88 − 9.42 − 8.85 − 8.40 − 8.72 − 8.71

Maximum intra-sample variability δ13C(‰)

δ18O(‰)

0.42

0.61

0.22

0.11

0.36

0.3

0.28

0.21

0.12

0.12

0.38

1.03

0.07

0.32

tation. The proportion of C3 vegetation in the ecosystem (X, where X = C3 / (C3 + C4)) can thus be approximated: (− 27)X + (1 − X)(− 12) = δ13C(SOM). Results are shown in Fig. 4, expressed in percentage C3 or X(100). Throughout the gilgai complex, the source of SOM was predominantly C3 vegetation. The proportion of C3 decreased from 95–98% in the surface to ∼ 71% in the lowest horizon (190 cm in the microlow; Fig. 4), with a marked shift towards lower percentages of C3 vegetation at about 115–140 cm. This abrupt shift coincides with the master slickenside (second heavy line in Fig. 2), that marks the boundary between mottled SOM-rich layers and lighter-colored, SOM-depleted, clayey parent material. The SOM distribution and isotopic compositions in the Russian Vertisol gilgai complex are thus consistent with hotter and drier climate conditions at the beginning of Vertisol pedogenesis, during formation of deeper portions of the profile, and cooler, wetter conditions with a shift towards a more productive, C3 dominated ecosystem for at least the past 3000–3500 years of

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pedogenesis. Several recent papers suggest the climate in the region was marked by a relatively dry and warm middle Holocene, changing to cooler and wetter conditions about 3200–4000 years ago (Velichko and Starkel, 1994; Alexandrovskiy and Chichagova, 1998; Khokhlova et al., 2001a,b). 4.2. Microvariability of pedoenvironment reflected in SOM stable isotopic composition Soil microrelief influences the modern soil pedoenvironment and hydrology, which, in turn determines the isotopic composition of SOM. More positive δ13C values are associated with C4 plants, which grow preferentially under warm, water stressed environments (Ehleringer, 1988). In addition to SOM values, modern root specimens collected at a depth 0–10 cm have δ13C values that are ∼1.2‰ more negative than roots in the microhigh. Previous studies of this same site have shown that the vegetative composition across the gilgai complex had more xerophytic plants in the microhigh, and more hydrophytic plants in the microlows (Kovda et al., 1999). Thus, in addition to the regional climate history, variability of stable isotope compositions across the gilgai complex, with somewhat lower values of δ13C in the microlow compared to the adjacent microhigh (Fig. 5), reflect local differences in the hydrology of the soil, with a wetter pedoenvironment in the microlow and the differentiation of plant species across the gilgai complex. Similar, slight differentiation of modern vegetation and SOM isotopic compositions have been observed in a tropical gilgai soil complex in Texas (Miller, 2000). 4.3. Pedogenic processes inferred from the stable isotopic composition of carbonate pedofeatures The controls on pedogenic carbonate formation can be complex and one would expect differences in carbonate isotopic compositions as a result of their age, mechanisms of formation, and morphology. Carbonate initially precipitated as soft masses in soils may undergo dissolution and reprecipitation, in situ or accompanied by translocation, from less stable micritic forms to more stable microspar or sparic crystal forms, with consequent induration or cementation to form hard nodules. These features may themselves undergo subsequent alteration, including nodule growth or dissolution, depending on superimposed climatic or soil hydrologic changes. Thus, carbonate pedofeatures may represent several generations of pedogenesis. As any of these processes may alter their isotopic composition, interpretations of soil carbonate isotope compositions must draw on all available, relevant

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observations. The following section considers interpretations of the isotopic compositions of carbonate pedofeatures with regard to their depth in the pedons, position in gilgai microrelief, age, and microfabric morphology. 4.3.1. Establishing the pedogenic nature of carbonate pedofeatures It is important to establish that carbonate pedofeatures are indeed pedogenic, and have not been inherited in the soil parent material. We note that the carbon isotope composition of pedogenic carbonates is normally enriched by ∼ 14–16‰ relative to coexisting SOM (Cerling, 1991). This is the range of isotopic fractionation noted for all carbonate morphologies and SOM at the same level within the Vertisol gilgai complex. The oxygen isotope composition of soil carbonates reflects the isotopic composition of the soil pore water composition, derived from precipitation, although typically enriched by ∼ 1–4‰ compared to precipitation due to evaporation of soil water (Cerling, 1984; Quade et al., 1989). The modern day δ18O of meteoric water for the study area is around − 10.3‰ SMOW (Brezgunov et al., 1998); pedogenic carbonate oxygen isotope compositions in the range − 7.6‰ to − 9.9‰ PDB (23.1 to 20.7‰ SMOW) are consistent with precipitation from, or recrystallization and exchange with, meteoric water similar in composition to modern values at ∼ 5–13 °C (Friedman and O'Neil, 1977). Both carbon and oxygen isotope compositions of pedogenic carbonate are very different from typical marine carbonate (both δ18O and δ13C close to 0‰ PDB), indicating carbonate pedofeatures are not inherited from the marine parent material, but have been formed during Vertisol pedogenesis. Systematic variability in the carbon and oxygen isotopic compositions of carbonate pedofeatures with depth or pedofeature morphology may reflect their detailed history of formation and is further explored below. 4.3.2. Variation of carbonate isotopic compositions with depth Carbonate pedofeatures show a slight enrichment in 13 C and 18O with depth (Fig. 6A–D), however, this trend is largely defined by a few samples of soft masses measured at the deepest levels (N 120 cm). The 13C enrichment trend is neither as systematic or as large as is noted in SOM isotopic compositions (Fig. 4), in fact, almost the entire range of δ13C values (− 10.5‰ to − 12.4‰) is noted in the top 70 cm of the microhigh. The variability of the carbon isotopic compositions, especially near the surface of the microhigh, may reflect several factors, including addition of carbonates — via

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lateral shearing or migration of carbonate along slickensides — from portions of the soil that formed under different pedogenic conditions, or an increased component of isotopically-heavy atmospheric CO2 in the upper horizons of the soil during carbonate precipitation. In the microlow, leaching of carbonates above 110 cm depth has resulted in a relatively thin carbonate horizon (∼40 cm thick), which varies over the entire range of δ13C values noted. Carbonate δ18O values in the microlow are ∼ 1‰ lower than carbonates sampled in the microhigh. Rather than preserving a simple relationship to depth, variability in both carbon and oxygen isotope compositions appear to more closely relate to pedogenic carbonate morphology and the conditions under which these carbonate morphologies formed, as discussed in Section 4.3.4, below. 4.3.3. Isotopic composition and the position in the microrelief The isotopic compositions of carbonate pedofeatures vary, with considerable overlap, as a function of gilgai position. Carbonate pedofeatures have more negative δ13C and δ18O values in the microlow, intermediate values in the microslope, and relatively enriched ratios in the microhigh (Fig. 6A–D). These trends are consistent with better drainage (drier soil conditions) along the gilgai complex from microlow to microhigh. The wetter environment in the microlows leads to a larger component of C3 vegetation, higher plant productivity and CO2 production, lower gas diffusivity through the soil matrix, and thus more negative carbon isotopic compositions for carbonates precipitated in this portion of the gilgai complex compared to the microhigh (Cerling, 1984). Wetter soil conditions in the microlow also reduce soil water evaporation, thus δ18O values in microlow carbonates, particularly close to the soil surface, may tend to be more negative than in the microhigh. The better drained conditions of the microhigh supports plants compatible with slightly more xeric conditions, generally lower productivity, higher soil gas diffusivities, more effective evaporation of soil water and, consequently, relatively enriched carbon and oxygen isotope ratios in the microhigh compared to the microlow. Carbonate isotopic compositions in the microslope are most similar to the microlow, suggesting that the soil water regime during carbonate precipitation is similar in the microslope and microlow, but differs from the water regime in the microhigh. The same result was obtained via direct seasonal measurements of water content in these soils (Kovda et al., 1996a). The range of isotopic values is smallest in the microslope, compared to the microlow and microhigh. The

more consistent isotopic values may reflect, in part, relatively small pedoclimatic and vegetative vectors in this microtopographic position during pedogenesis. In the microhigh and microlow positions, nearly two-times larger isotopic variability may reflect the more complicated soil evolution or stronger changes of pedoenvironment and vegetation during pedogenesis due to, and in part responsible for, microrelief formation. Alternatively, the variability of carbonate isotopic compositions may simply reflect the polygenetic nature of the carbonate pedofeatures for which isotopic signatures were obtained; a greater variety of pedofeatures is noted in the microhighs and microlows. 4.3.4. Carbonate morphology and isotopic composition Interpretation of the stable isotope composition of pedogenic carbonate in the Vertisol gilgai complex is best made in the context of the carbonate morphology, as there are clear differences in the compositions of different phases (nodules, soft masses; Fig. 5A,B) and of these phases in different microtopographic positions (i.e., soft masses, Fig. 7). In a previous micromorphology study of the carbonate pedofeatures (Kovda et al., 2003), carbonate nodules were found to be compound, septaric, nucleic and aggregated, consisting of a microsparite core surrounded by a concentric layer of micrite, and having sharp, well-defined boundaries with the soil matrix. Hard nodules in the Vertisol microlow pedons were larger in size and morphologically more complex than nodules in microhighs, reflecting the influence of shrink-swell, lateral shearing, an increased degree of aggregation of small nodules and modification by superimposed hydrologic conditions. Soft carbonate masses have micritic microfabric and vary with soil depth and microtopographic position in their degree of impreg-

Fig. 7. Carbon and oxygen isotope compositions of soft carbonate masses (SM) according to their position in the soil microrelief.

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nation, aggregation into pellets, and the presence of hard nodular cores. Current calcite deposition is indicated by variable development of a halo around many soft masses and the presence of calcite hypocoatings in internal soil voids. An evolutionary model for carbonate pedofeatures in these Vertisols (Kovda et al., 2003) suggests that hard nodules represent a relatively early pedogenic product, initiated before or in the early stages of gilgai formation, with soft carbonate masses comprising more recent pedogenic material, and soft masses with hard cores intermediate between the two. Micromorphological details of the carbonate provide critical observations for discerning superimposed pedogenic processes or events and belie a more complex history of carbonate formation and recrystallization. For example, radiocarbon dating of hard nodules (Kovda et al., 2003) indicates that nodules from the microlow are younger (1300 ± 130 yr BP) than the nodules from the microhigh (1880 ± 120 yr BP). In both cases, the nodules are markedly younger than organic matter in the same position of the soil (2290–3720 yr BP). In the microlow, nodules consist primarily of smaller nodules coalesced into larger nodules by thick layers of micritic cements. Thus, a significant amount of the “hard nodule” is more recent microsparitic cement, indurating the nodule (Fig. 3A) and giving it, in bulk, a younger age. Nodules in the microhigh consist of a relatively large, older microsparitic core surrounded by only a thin, abraded, micritic layer (Fig. 3B), and these yield an older average radiocarbon age. Clearly, all nodules experienced the addition of micritic pedogenic carbonate subsequent to their formation, and their coarser-grained microsparitic cores suggests some post-depositional dissolution and reprecipitation; this observation proves critical to interpretation of their stable isotopic compositions. To reconcile the paragenesis of carbonate pedofeatures inferred from micromorphology, stable isotopic compositions of the pedofeatures, ages determined for different microtopographic positions and horizons within the gilgai complex, and the climate history interpreted for this region, it is important to recognize the complexity of the carbonate–SOM–age relationships. Hard nodules, interpreted to be the oldest carbonate pedofeature on the basis of micromorphological relationships with other carbonates, generally yield the most negative values of δ18O and δ13C measured in the various pedofeatures at a given depth of the soil, and are also consistently more negative in the microlow compared to the microhigh. These isotopic results seem at odds with interpretation of coexisting SOM carbon isotope compositions, and other studies, that suggest deeper soil material preserves evidence of a warmer and drier

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climate conditions during the mid-Holocene, until about 4000 yrs ago (Velichko and Starkel, 1994; Alexandrovskiy and Chichagova, 1998; Khokhlova et al., 2001a,b). Evidence of nodule recrystallization (microsparitic texture), the abundance of micritic rims and septarian cements in the nodules, and relatively young carbonate 14C ages compared to coexisting SOM, suggest the nodules have experienced a complex history of recrystallization and exchange throughout their pedogenic history and their isotopic compositions reflect more recent soil conditions than the SOM, including a predominantly C3 ecosystem and a cooler and/or wetter environment. The small, but systematic, isotopic differences between microlow and microhigh nodules most likely reflect the better drained microhigh environment. Soft masses show surprising variability in their isotopic compositions, which can be related to their microtopographic position and the pedogenic horizon in which they occur (Figs. 6 and 7), but do not strongly correlate with the observed differences in their microfabric (i.e., aggregated or undifferentiated; dense or weak degree of impregnation). The δ13C values of soft masses are up to ∼ 2‰ more positive, and δ18O values of soft masses are up to ∼ 1.5‰ more positive, than nodules or hard cores at the same level of the soil (Fig. 6A,C), although the largest shifts are mostly associated with deep, soft masses (discussed separately below). There is greater differentiation in the compositions of the different morphologies in the microhighs. The small shift to more positive compositions suggests a temporal change towards slightly cooler temperatures (given meteoric water similar to modern values). More positive oxygen isotope values in the better-drained microhigh likely reflect the increased influence of soil water evaporation. Although more positive carbon isotope compositions in the soft masses, particularly in the microhigh, may be interpreted to indicate a recent shift in ecosystem to include a greater proportion of C4 plants, an interpretation more consistent with the overall climate history, and modern soil vegetation determined in other studies (Kovda et al., 1999) is that the carbon isotope values reflect a slightly lower rate of soil CO2 respiration during formation of the soft masses, and therefore admixture of a greater proportion of isotopically-heavy atmospheric CO2 in the soil gas (Cerling, 1991). Soft masses occurring in the oldest and deepest portions of the soil (i.e., in the yellowish matrix of the deepest horizons within the microlow, 125–150 cm depth, and within the microhigh chimney, 50–180 cm; Fig. 7; Fig. 3C) call for closer examination. These horizons retain the oldest SOM (N 5600 yr), based on 14C ages of SOM (Kovda et al., 2005), and the carbonate

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masses have the most positive values of δ13C and δ18O for all carbonates we measured, particularly where the deep matrix material penetrates into the microhigh chimney (Fig. 2; Fig. 6B,D; Fig. 7, open circles). The 13 C- and 18O-enriched deep soft masses seem, on first consideration, to be consistent with an earlier, warmer, more arid climate and C4-richer ecosystem inferred from SOM isotopic compositions. But, at isotopic equilibrium in a soil, the carbon isotope composition of pedogenic carbonates is normally enriched by ∼ 14–16‰ relative to coexisting SOM (Cerling, 1991). Given SOM with δ13C values in the range − 23.5‰ to − 22.7‰, pedogenic carbonate formed in equilibrium with the SOM is expected to have a composition of ∼ − 6.5 to − 9.5. By comparison, the deep soft masses have δ13C in the range − 10.4‰ to − 11.7‰. The lack of isotopic equilibration with the SOM, and similar isotope compositions of soft masses at higher levels in the microhigh, point to modest isotopic resetting of the deep, soft carbonate during more recent pedogenesis, and compositions that, which still enriched relative to other carbonate pedofeatures, reflect more recent climate conditions than the coexisting SOM. Failure to recognize this resetting would lead to underestimation of the magnitude of ecosystem and climate change. Similarly, relatively enriched oxygen isotope compositions of deep, soft masses would at first glance support the SOM-based interpretation of early pedogenic conditions that were warmer or drier than modern conditions. The δ18O values, however, are also likely shifted from their original compositions. These isotopic results help to elucidate the generation of soft carbonate masses not clearly understood by macro- and micromorphological evidence alone (Kovda et al., 2003), and points to the value of detailed analysis combining morphological and isotopic approaches in discerning the details of sometimes complicated pedogenic histories. 5. Summary The stable isotope compositions of SOM and carbonate pedofeatures record a complex, but interpretable, pedogenic history. Variability in the isotopic composition of pedogenic carbonates as a function of morphology, position in gilgai microrelief, and soil depth emphasizes the importance of these parameters in reconstructions of Vertisol pedogenesis or paleoclimate. The soil records evidence of pedogenesis in an earlier (N3200 yr BP) environment that was warmer and drier than modern conditions, consistent with the regional climate history suggested for this area. A change to cooler and wetter conditions resulted in a shift to an ecosystem that is N95% C3 vegetation. The wetter environment drove dissolution

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