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Isotopic evidence for dolomite formation in soils 

J.L. Díaz-Hernández, A. Sánchez-Navas, E. Reyes 

PII: 

S0009-2541(13)00141-1 

DOI: 

doi: 10.1016/j.chemgeo.2013.03.018 

Reference: CHEMGE 16850 

To appear in: 

Chemical Geology 

Received date: 19 November 2012 

Revised date: 25 March 2013 

Accepted date: 28 March 2013 

Please cite this article as: Díaz-Hernández, J.L., Sánchez-Navas, A., Reyes, E., 

Isotopic evidence for dolomite formation in soils, Chemical Geology (2013), doi: 

10.1016/j.chemgeo.2013.03.018 

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ACCEPTED MANUSCRIPT 

ISOTOPIC EVIDENCE FOR DOLOMITE FORMATION IN SOILS 

Díaz-Hernández, J.L. 1 , Sánchez-Navas, A. 2,3 , Reyes, E. 3 

1 IFAPA Camino de Purchil, Área de Recursos Naturales, Consejería de Agricultura, Pesca y Medio 

Ambiente, Junta de Andalucía, Apartado 2027, 18080 Granada, Spain. josel.diaz@juntadeandalucia.es 

2 Departamento de Mineralogía y Petrología, Universidad de Granada, Campus Fuentenueva, Granada, 

Spain. 

3 Instituto Andaluz de Ciencias de la Tierra IACT (CSIC-UGR), Avda. de las Palmeras, 18100 Armilla, 

Granada, Spain. 

Abstract 

Dolomite formation in soils constitute a particular challenge because of: 1) scant 

magnesium content in continental environments as opposed to the marine medium, 2) 

the kinetic problem related to the incorporation of magnesium into the carbonate, and 3) 

the unknown role of soil dolomites in the global carbon cycle. 

Pedogenic dolomite formed at deeper soil levels (subsoil) before the development of 

petrocalcic horizon barriers was investigated in a semiarid region of SE Spain (Guadix- 

Baza basin). Mineralogical characterization, textural relationships and isotopic data 

concerning soil dolomite, together with the results of a precipitation experiment, 

provided fuller knowledge of the processes and conditions governing neoformation of 

dolomite in these soils. 

In the study case, dolomite enrichment occurs beyond the limit of major biological 

activity, which coincides with the rooting depth of native perennial plants in the 

semiarid soils studied. Textural studies reveal the corrosion of inherited dolomite 

crystals in the upper soil horizons and the formation of dolomite in depth in relation to a 


clayey material, composed mainly of smectites. Stable isotope distribution in dolomites 

throughout the profiles indicates a fractionation with depth. This is explained by the 

formation of dolomites after the dissolution of the pedogenic calcite. The calcite 

detected in the subsoil is interpreted here as a precursor of the neoformed dolomites that 

transport the isotopic signal associated with biological activity of soils to deeper layers. 

Dolomite formation appears be favoured by the presence of clay minerals in the 

precipitation media. Clays retain water during evapotranspiration stages, which 

drastically change the transport properties of the media and promote the incorporation 

of Mg into the structure of the neoformed Ca,Mg-carbonate. As confirmed by 

laboratory experiments, diffusion-controlled crystal-growth processes lead to the 

1


formation a precursory “protodolomite” with disordered Ca,Mg distribution from a fluid 

locally supersaturated in dolomite. 

Keywords: Carbon-Oxygen isotopes, Dolomite problem, Inorganic carbon, Pedogenic 

processes, Semiarid soils, Soil carbon storage. 

1. INTRODUCTION 

The genesis of calcic and petrocalcic horizons in soils occurs as part of carbonate 

dissolution-precipitation cycles (Gile et al., 1966; Salomons and Mook, 1976). Detrital 

limestone and other calcareous materials undergoing decarbonation constitute a source 

for secondary carbonate accumulation in soils. Biogenic and atmogenic CO 2 are 

alternative sources for the carbon fixation in the soil under physico-chemical conditions 

stabilizing the carbonate ion. The pedogenic environment is strongly influenced by 

biological activity: root mats, in addition to associated lichens and microorganisms, 

contribute directly to mineral dissolution-precipitation in soils. In fact, pedogenesis 

begins with intense biological activity in the weathering zone (Wright and Tucker, 

1991; Wright, 1992; Cerling and Quade, 1993). Carbonic acid produced in association 

with biological activity (mainly roots) is the main acid involved in the dissolution of 

calcium carbonate. The leaching depth of the dissolved carbonate augments with 

increasing mean annual precipitation and temperature (Arkley, 1963; Stevenson et al., 

2005). Carbonate precipitation at the deeper layers results largely from the declining 

P CO2 in the soil and from the higher concentrations of solutes in the soil solution due to 

evapotranspiration. In this sense, the carbonation process is favoured in arid-semiarid 

regimes. Large amounts of soil inorganic carbon, stored in the form of calcium 

carbonate, has been reported in cultivated mollisols under moderately cold temperatures 

(5-5.5ºC), and with a mean annual precipitation between 581-397 mm (Mikhailova and 

Post, 2006; Mikhailova et al., 2009), suggesting an increase in the carbonation process 

with agricultural practices. 

Isotopic relationships found in pedogenic carbonates in arid-semiarid ecosystems 

indicate that carbonate dissolution-precipitation cycles in a pedogenic environment 

provoke changes in the isotopic composition of the precipitated carbonate, by inducing 

significant inputs of biogenic carbon (Cerling, 1984). Biogenic carbon carries a unique 

isotopic signature, and the biological fractionation processes deplete 13 C to negative 

values (Salomons et al., 1977; Magaritz and Amiel, 1980; Magaritz et al., 1981). 


2


Therefore, the pedogenic carbonate will usually have more negative 13 C than detrital 

carbonates. 

Modern dolomites are formed only in certain environments such as meteoric, marine, 

hypersaline, subsurface brines, and hydrothermal (Hardie, 1987). The precipitation of 

dolomite and dolomitization of previous Ca-rich carbonates are among the leastunderstood 

problems encountered in the geochemistry of carbonates: e.g. surface 

seawater is strongly oversaturated in dolomite, but little evidence is available on the 

widespread dolomite precipitation in modern, open-marine sediments. To explain the 

“dolomite problem”, models have emerged ascribing microbial activity to the mediation 

of dolomite precipitation. Various authors have focused their attention on the role 

played by the bio-mediation that leads to CO 2 enrichment and raises alkalinity, and 

presumably drives dolomite formation in Mg-rich environments (e.g. Wright, 1999). In 

any case, the occurrence of dolomite in the geochemical systems of the earth’s surface 

is often strongly controlled by reaction kinetics (Arvidson and Mackenzie, 1999; 

Deelman, 2005) and/or mass-transport processes. 

The formation of dolomite in soils is rarely reported, its origin being related to parentmaterial 

inheritance, to atmospheric Mg 2+ wet deposition, and to the addition of 

dolomite contained in atmospheric dust (Ghebre-Egziabhier and St. Arnaud, 1983; 

Ming, 2002; Goddard et al., 2007; Díaz-Hernández, et al., 2011). The occurrence of 

neoformed dolomite in soils is usually restricted to saline environments (Shermann et. 

al., 1962; Kohut et al., 1995). However, the presence of high-Mg water solutions in 

soils derived from weathering of serpentinites and basaltic rocks may lead to the 

formation of authigenic dolomite in non-saline environments (Podwojewski, 1995; 

Capo et al., 2000). Although the formation of dolomite in soils has been discussed 

(Drees and Wilding, 1987; Sobecki and Karathanasis, 1987; Bui et al., 1990; Whipkey 

and Hayob, 2008), no mechanism explaining the precipitation of dolomite in soils has 

been reported. The incorporation of magnesium into the calcite structure to form high 

magnesian-calcites and proto-dolomites under surface conditions is broadly described in 

the literature (Hardie, 1987 and references therein; Deleuze and Brantley, 1997). 

Sedimentary and diagenetic dolomites probably precipitated as protodolomites and their 

isotopic composition was controlled by the protodolomite-water fractionation (Fritz and 

Smith, 1970). In any case, the incorporation of magnesium to the rhombohedral 

carbonate structure is not easy and appears to be controlled by kinetic factors (Mucci 

and Morse, 1983; Putnis et al., 1995; Fernández-Díaz et al., 1996). 


3


2. SCOPE OF THE INVESTIGATION 

In this paper, we study the occurrence of pedogenic dolomite in soils of an arid/semiarid 

region. Calcite is by far the most abundant and best-studied pedogenic carbonate. On 

the other hand, dolomite is not usually evaluated in soils and its role in the pedogenic 

processes is scarcely known. Here, we describe the formation of dolomite at soil depths 

below indurated layers delimitated by petrocalcic horizons or caliches. Deeper layers 

(subsoil) in these types of soils are scarcely studied because caliches are hard to 

excavate and their features at these depths are irrelevant for soil taxonomy or for 

judging the soil as a support for plants. 

Several reasons underlie this approach: 1) Knowledge of the distribution of carbon in 

the deep-soil-layer system in semiarid areas is essential to assess soil carbonates (calcite 

and dolomite) as a sink for atmospheric carbon. However, the carbon distribution in 

deeper layers (subsoil) has rarely been determined (Jobággy and Jackson, 2000; Díaz- 

Hernández et al., 2003). 2) Well-ordered dolomite, together with protodolomites, has 

recently been documented in soils (Whipkey and Hayob, 2008). 3) The soils of the 

study area are particularly rich in inherited dolomites and, therefore, could be suitable 

for the development of pedogenic dolomites in addition to calcites. 4) Periodic moisture 

movement to depths below solum horizons serves as a regulating factor in maintaining 

soluble Mg at levels favourable to the precipitation of low-magnesium calcites (St. 

Arnaud and Herbillon, 1973), and occasionally also leads to the formation of dolomites. 

5) The few studies available concerning pedogenic dolomites (St. Arnaud, 1979; Botha 

and Hughes, 1992; Capo et al., 2000) point to the occurrence of pedogenic dolomites in 

depth. 

The main objective of this research is to demonstrate the formation of dolomite in soils. 

For this, we firstly studied compositional, mineralogical, and textural variations 

throughout the soil profiles, finding that dolomite was dissolved in superficial horizons 

and precipitated at greater depths than was the pedogenic calcite. Also, we analysed the 

isotopes in calcite and dolomite of the soil samples. This second stage revealed that the 

biogenic isotopic signature of the pedogenic calcite was transferred to a neoformed 

dolomite through the dissolution-precipitation cycles that affected the soil carbonates. 

Finally, we discuss the physico-chemical constraints on the dolomite formation in 

subsoil. According to the model proposed for the precipitation of dolomite: 1) inherited 

dolomite constitutes the main source for the Mg 2+ ; 2) hydrodynamics in the soil studied 


4


controls moisture and solute transport properties in the subsoil; 3) the incorporation of 

magnesium to the carbonate structure in Mg-rich confined porous media to form 

dolomite is explained from the standpoint of kinetics. 

3. MATERIAL AND METHODS 

3.1. Geological setting and climate 

The Guadix-Baza basin is located in the Betic cordillera (SE Spain, Fig. 1A). The area 

studied was developed as an endorheic depression from the Late Neogene to the Late 

Pleistocene (Vera, 1970; Peña, 1985; García-Aguilar and Martín, 2000; Gibert et al., 

2007, among others), and occupies about 3500km 2 . During this period, the basin was 

filled with alluvial and lacustrine materials for which the composition varies depending 

on the nature of source materials. Those from the Subbetic ranges (North border) are 

rich in limestone, while those from the Betic ranges (South border) are rich in quartzite, 

micaschist, and dolomite (IGME, 1973-1999). The Guadix sub-basin is comprised of 

fluvial sediments (silts, sands and conglomerates) and the Baza sub-basin is formed 

mainly by lacustrine deposits (limestones, marly limestones, and gypsum). The rocky 

relief surrounding the Guadix-Baza depression coincides partially with the Atlantic- 

Mediterranean divide. The depression is characterized by having elevations ranging 

from 1500m a.s.l. (above sea level) on the edges to 650m a.s.l. at the Guadiana Menor 

River, the main drain of the area. 

The capture of the internal drainage network of the depression by the Guadiana Menor 

River (a tributary of the Guadalquivir River) modelled the landscape into two main 

geomorphological groups (Díaz-Hernández and Juliá, 2006): 

1) Old surfaces (350,000-205,000 yr): S1, S2, S3, consisting of dissected piedmont 

plains gently sloping towards the centre of the basin (Fig. 1B), capped by caliche soils 

with thick petrocalcic horizons (Petric Calcisols and Chromic Luvisols (FAO- 

UNESCO, 1988) (Calciorthids, Paleorthids, Paleargids, Haploxeralfs, Rodoxeralfs and 

Xerochrepts, following the Soil Taxonomy (SSS, 1999)). Dolomite enrichment 

observed under the petrocalcic horizon (see below) suggests that the formation of 

dolomite began within this broad time range. 

2) The youngest surfaces (115,000-48,000 yr): S4 (badlands, with Gypsic, Calcaric and 

Eutric Regosols (FAO-UNESCO, 1988) (Xerortents, SSS, 1999) and S5 (alluvials, with 

Calcaric Fluvisols (FAO-UNESCO, 1988) (Xerofluvents, SSS, 1999). 


5


The climate of the basin is semiarid, hot in summer, with frequent frosts in winter. The 

mean annual temperature is 15ºC. Mean annual precipitation is about 250-300mm in the 

central part and about 400mm in the higher fans located along the mountain fronts 

(MAPA, 1989). Rainfall occurs mainly during the autumn-winter months; in summer, 

rain is scarce, falling as infrequent torrential storms. 

3.2. Sampling 

Soil profiles were studied at 81 sites in the Guadix-Baza basin (Fig. 1A), where the 

depth and features of the caliches could be accurately determined. Natural or artificial 

pits reached an average depth of about 200cm (in several cases down to 400cm) and the 

sampling sites were distributed as evenly as possible, in 10 × 10km cells to simulate a 

probabilistic sampling, with one or two sites selected per cell. Data from the deepest 

layers were not considered here because of the statistical results from the few data 

(N


The fine fraction (


3.6. Electron microscopy 

Two samples corresponding to the fine fraction of soil material located below the 

petrocalcic horizon at 120 and 220cm depths (profile 46) were selected on the basis of 

their dolomite content. They were consolidated and prepared as polished sections and 

later carbon coated for the textural and mineralogical study by analytical scanning 

electron microscopy. Back-scattered electron (BSE) images and energy-dispersion X- 

ray (EDX) spot analyses (200nm of diameter) were performed with a high-resolution 

field-emission scanning electron microscope (FESEM) Auriga (Carl Zeiss) equipped 

with a LinK INCA 200 (Oxford Instruments) microanalysis system operated at an 

acceleration voltage of 10-20kV. 

For transmission electron microscopy (TEM), the same


4. RESULTS 

4.1. Compositional, mineralogical and particle-size variations in the soil profiles 

The distribution of organic and inorganic carbon with depth varied considerably in the 

soils of the study region (Figs. 2A-C; see also Fig. 1 in Díaz-Hernández, 2010), 

reflecting the heterogeneity of the area, in which several soil types are represented 

(Díaz-Hernández et al., 2003). Profiles of organic and inorganic carbon (calcite and 

dolomite) distribution with depth for alluvial soils, badlands, and piedmont units are 

depicted in Figs. 2A-C. Organic carbon decreases with depth in all soil types, and it is 

particularly marked in the case of piedmonts (Fig. 2A). Petrocalcic horizon, represented 

by the enrichment in calcite content in depth, is detected mostly in piedmonts. This 

horizon appears between 10cm and 90cm depth, with 30-130cm of thickness (Fig. 2B, 

see also Díaz-Hernández et al., 2003). An increase of the dolomite content with depth is 

also better appreciated in piedmont profiles, usually below the petrocalcic horizon 

(compare Fig. 2B with 2C). In the soils studied, dolomite enrichment can occur at great 

depth (below 200cm), as is the case of the profile PC-43 of the Fig. 2C. This can be 

explained by the presence of Mg-rich substrates, constituted by altered ophitic rocks in 

the subsoil. The development of the petrocalcic horizons and the occurrence of 

associated dolomite enrichment in depth are better observed in the case of piedmonts. 

A total of 81 profiles were analysed, 8 corresponding to alluvial, 11 to badlands, and 62 

to piedmonts, in order to gain statistical information concerning carbon distribution in 

depth for the soils studied. Table 1 and Fig. 2D show the depth distribution of soil 

organic carbon content (from organic matter) and calcite and dolomite content (soil 

inorganic carbon), expressed as an average percentage. The average organic carbon 

decreased rapidly with depth, tending to near-zero values below 120cm (N = 81 

profiles), and continued to decline gradually to a depth of 220cm (N = 27 profiles). The 

calcite content initially increased with depth, showing minor amounts near the surface 

and a maximum between 40-60cm depth (43% in the bulge). Below this depth the 

calcite diminished gradually to 200cm in depth (around 30%). A second bulge at a 

depth of 200-280cm appeared from the profiles of the units of type S1. Polycyclic 

profiles showing this second maximum of calcite accumulation within this depth range 

were found for S1 unit (not shown). These minor maxima correspond to hard substrates 

or paleosols in those units, probably caused by regional sedimentological changes, and 

are not considered in this work. The average depth distribution of dolomite began with a 

low content (9% in the superficial horizons) which progressively increased from 40cm, 


9


reaching 17% at 140cm depth. However, the depth distribution of the dolomite and 

calcite contents proved somewhat irregular below 180cm due to the low number of 

samples for these depths (Table 1) Average particle-size distributions throughout the 

soil profiles studied show high contents of sand and gravels with respect clays and silts 

(Fig. 2E). 

The XRD study evidenced that the major minerals of soil samples were quartz, calcite, 

and dolomite, with minor amounts of clays, feldspars, and micas. For dolomite-rich 

samples, superstructure reflection (015) (the main reflection of well-ordered dolomites; 

Goldsmith and Graf, 1958; Reeder and Sheppard, 1984) was present. The depth 

distribution of quartz, calcite, and dolomite was shown in the XRD mapping of Fig. 3, 

corresponding to a depth sequence of diffractogram patterns for profile 46 (piedmont). 

This image shows a decrease in quartz with depth on the basis of the change in the 

intensity of the (101) reflection, and an interruption of the signal coinciding with the 

maximum of calcite enrichment (100cm depth). By contrast, the dolomite (104 peak) 

showed an inverse distribution to that of quartz (in-depth enrichment), with the same indepth 

interruption and lack in superficial horizon. Calcite enrichment between 60- 

130cm of depth, as indicated by the increase in the intensity of the (104) peak, 

coincided with the development of a petrocalcic horizon of 70cm thick in this profile. 

4.2. Isotopic data 

The frequency distribution of 

13 C and 

18 O of the calcite in the study samples are 

shown in Figs. 4A and B, respectively. These values were consistently negative, and 

13 C ranged from -13 to -1‰ and 

18 O from -10 to -2‰, with a normal distribution and 


modes around -7‰ in both cases. These negative values indicate a contribution of 

isotopically light carbonate. The histograms for the 

13 C and 

18 O of the dolomite 

showed bimodal distributions (Figs. 4C and D): the two modal values for 

13 C were -5 

and -1‰, and -6 and -3‰ for 

18 O. The scatter diagram relating both values for the 

dolomite samples also showed these two different tendencies in the histograms, and two 

clusters may be identified within the diagram (Fig. 4E). The average isotopic 

composition of the inherited dolomites from the coarse fraction of the studied soils is 

also represented in the bivariate plot (star in Fig. 4E). It plotted close to the analysed 

dolomite cluster with less negative and even positive values (cluster I in Fig. 4E). 

Samples of deeper layers (>70cm) defined the cluster with more negative 

13 C and 

18 O 

10


values (cluster II in Fig. 4E) that corresponded to the smallest left-hand maximum 

observed in the histograms (Figs. 4C and D). However, none of single profiles studied 

here showed a clear isotopic fractionation with the dolomite enrichment in depth. 

4.3. SEM and TEM study 

Electron microscopy confirmed the XRD observations of Fig. 3 regarding dolomite 

dissolution and calcite precipitation in relation to the formation of petrocalcic horizon. 

BSE images of Figs. 5A and 5B correspond to a sample below the petrocalcic horizon 

(120cm depth), where the corrosion affecting dolomite and quartz crystals and the 

precipitation of calcite are visible. In Fig. 5A the voids, after quartz dissolution, appear 

to be filled mainly with calcite. Frequently, calcification is widespread and calcite 

nodules occur with dolomite relicts within the concretions (Fig. 5B). The calcified 

nodules frequently had a highly porous texture. The dissolution process of inherited 

dolomites sometimes began along crystallographic planes and advanced from inner to 

outer regions with progressive corrosion, forming hollow dolomite crystals (Fig. 5C). In 

addition, corroded dolomite grains sometimes showed central voids (Fig. 5D). Calcite 

rhyzocretions also occurred close to the petrocalcic horizon (Fig. 5D). 

Samples from 220cm in depth showed corroded calcite grains surrounded by dolomite 

and clay minerals (Figs. 5E-F). Fig. 5F, corresponding to an inset of Fig. 5E, shows 

dolomite+clays secondary products located in a corrosion gulf which includes 

fragments of partially altered calcite relics. A detailed observation of precursory calcite 

evidenced the presence of a poorly crystalline material, basically clay minerals, close to 

irregular corrosion boundaries at the surface of calcite grains, in addition to inherited 

chlorites. Clay minerals also surrounded sub-idiomorphic dolomite crystals with sizes 

around 2 m in diameter, and sometimes localized preferentially between corroded 

calcite grains and rhombohedral dolomite crystals (Figs. 6A and 6B). Poorly defined 

crystalline material surrounding dolomite consisted of smectites plus Si-rich amorphous 

substance (Figs. 6C, 7 and Table 2). According to AEM analyses of Table 2 the 

smectite is intermediate between beidellite and montmorillonite, although several 

analyses correspond to mixtures of clay minerals and Si-rich amorphous substances. Si 

excess cannot be incorporated in the smectite structure and must therefore derive from 

adjacent non-crystalline material. Bubbles occurring in relation to clay minerals and 

amorphous substance in TEM image (Fig. 6C) resulted from the gradual release of 

volatile substances after pervasive electron bombardment. Clay minerals were also 


11


studied by XRD from separate fractions of the sample at 220cm depth (Fig. 8). This 

study confirms the occurrence of the following clay minerals: 1) An inherited chlorite is 

evidenced by (002) and (004) peaks (Fig. 8A; see also BSE image and qualitative 

microanalysis of Fig. 6B); 2) a detrital illite with peaks at 1.0 and 0.5nm d-spacing 

corresponding to (001) and (002) reflections; 3) abundant smectites and minor illitesmectite 

mixed-layer phases (probably R3-ordered I/S). The latter is evidenced by the 

broad basal reflection of the 1.0-1.2nm zone in the untreated sample (Fig. 8A). The 

smectite had a basal spacing of around 1.4nm in the air-dried state, which expands up to 

1.8nm upon ethylene glycol saturation (Fig. 8B). 

5. DISCUSSION 

5.1. Origin of compositional and mineralogical variations along the profiles 

Organic carbon shows the expected distribution (Figs. 2A and D) for these types of soils 

with a usual decrease with depth due to the more intense biogenic activity in the upper 

horizons. The combination of water from rainfall, Ca 2+ ions from diverse sources and 

CO 2 generated mostly from organic-matter decomposition and root respiration (Gile et 

al., 1966; Deines, 1980; Rech et al., 2003), the transient decrease in P CO2 and the 

increase in the concentration of solutes due to evapotranspiration lead to the 

accumulation in depth of pedogenic calcium carbonate in arid and semiarid regions. The 

formation of the petrocalcic horizon explains the occurrence of the bulge at 40-60cm 

depth in the calcite average regional profile (Fig. 2D), and the calcite enrichment found 

in the individual profiles shown in this work (Figs. 2B and 3). Although the dolomite 

content in the individual profiles was somewhat irregular (Fig. 2C), and in particular for 

depths below 180cm, it was evident that a regional decrease occurred near the surface 

with a marked enrichment in depth (Fig. 2D, Table1). Moreover, a smooth increase 

began at 40cm and continued to 140cm depth. Below 180cm depth, the values in the 

profile may have resulted from the interference of authigenic processes with the effects 

of the substratum on the mineralogy. Although the maximum in the case of the dolomite 

was less marked than in the calcite, the overall variation in dolomite (around 10%) was 

quite similar to the percentage change observed in calcite (around 12%). As occurred 

for calcite, dolomite dissolution/precipitation may explain the occurrence of the smooth 

maximum, just below the petrocalcic horizon, preceded by the change of sign in the 

curvature of the dolomite pattern in the Fig. 2D. Therefore, at the regional scale, 

dolomite underwent superficial dissolution and its precipitation occurred at greater 


12


depths than in the case of calcite. In short, the calcite-and dolomite-distribution patterns 

found here clearly indicate that they were affected by similar processes, but, in the case 

of dolomite, different dynamics seem to have occurred within the soil (see below). 

Dissolution of the dolomite in the superficial horizon of the studied profiles (e.g. Fig. 3) 

in relation to the development of petrocalcic horizon was also evidenced by Figs. 5A-D, 

where detrital dolomite appeared partially or almost completely dissolved and calcite 

precipitates in voids and formed rhyzocretions. A general decline in the quartz content 

was noted in the profile in the Fig. 3, contradicting the assumption that quartz content 

remains constant throughout the soil profiles (Lelong, 1969; Souchier, 1971; Sohet et 

al., 1988). The appearance of a physical barrier represented by the petrocalcic horizon 

prevents the redistribution of the upward-added quartz from diverse sources, mainly 

aeolian (Díaz-Hernández et al., 2011), and is therefore responsible for the observed 

decrease of quartz below it. However, the quartz content may decrease below the 

petrocalcic horizon as a result of dissolution processes, as evidenced by Fig. 5A, where 

the precipitation of calcite after quartz dissolution is evident. Calcite concretions after 

the dissolution dolomite and quartz had a marked porous texture (Figs. 5A and 5B) 

which may be interpreted in relation to bio-mediated precipitation processes. 

As mentioned above, dolomite enrichment below 40cm in depth was observed in the 

depth-wise variation of dolomite at the regional scale (Fig. 2D) and in the representative 

profiles of the Figs. 2C and 3. One possible explanation would imply no dolomite 

precipitation, but only dissolution; thus, the increment in-depth would be due to more 

intense dissolution for superficial horizons, the solutes being leached out of the system. 

On the other hand, in dissolution/precipitation mechanism proposed here, dolomite 

dissolved above 40cm is potentially precipitated below this depth. Figs. 5E-F and 6, 

corresponding to a sample situated at 220cm in depth in the studied profile, show the 

occurrence of dolomite and clay overgrowth in relation to calcite. Textural relations 

indicate the dissolution of calcite and the precipitation of dolomite in relation to clays. 

The specific mechanisms explaining the observed spatial association of calcite, 

dolomite and clays in SEM and TEM images are discussed below. 


5.2. Pedogenic isotopic fractionation and neoformed dolomites 

The proposed dolomite dissolution/precipitation model is consistent with the 

heterogeneity observed in the isotopic data from the dolomite samples analysed, which 

showed two different signals (Figs. 4C, 4D, and 4E). Cluster I in Fig. 4E, with higher 

13


isotopic values, lies close to the parent material that supplies the dolomite to the system 

(star in Fig. 4E). The precipitation of a neoformed dolomite below the petrocalcic (Figs. 

5E, 5F and 6) as a result of dissolution of inherited dolomite and pedogenic calcite of 

upper horizons explains the high dispersion observed in the δ 13 C values of the samples 

of deeper layers (cluster II in Fig. 4E and smallest left-hand maximum in the histograms 

of Figs. 4C and D). These values define the MDL (Meteoric Dolomite Line), a term 

similar to that proposed by Lohmann (1988) for the Meteoric Calcite Line (MCL). The 

C dispersion is explained by a mixing of inherited carbon from dolomite-rich parent 

material, with positive values, and pedogenic carbon from neoformed dolomite, with 

very negative C values. In addition, O values close to 6‰ (V-PDB) observed for 

pedogenic dolomites (cluster II in Fig. 4E) are consistent with the O measured in 

surface waters, values ranging from -7.5 and -9‰ (V-SMOW), and temperatures 

between ~17 and 24ºC. 

The isotopic composition of the dissolved inorganic carbon (DIC) and oxygen in soil 

waters below the petrocalcic horizon is in equilibrium with the isotopic composition of 

neoformed dolomites. Carbon and oxygen are distributed among three chemical species 

with a pH-dependent relative abundance (CO 2 , HCO - 3 and CO 2- 3 ). Most of the carbonic 

acid derives from the vegetal cover and prompts an increase of negative C in DIC 

(Cerling, 1984, 1991; Usdowski and Hoefs, 1993; Reyes et al., 1998). In the pedogenic 

medium studied, this isotopic signature is transported towards deeper layers through 

carbonate dissolution (mostly calcite) above the petrocalcic horizon. The negative 

carbon isotope composition of the neoformed dolomite is determined mainly by the 

isotopic pedogenic calcite because pore fluids generally have low carbon contents so 


that the 13 C of the precursor is generally retained (Hoefs, 2004). Low 

13 C and 

values are expected in biomediated calcite and dolomite precipitates (Jones and 

Donnelly, 2004; Wacey et al., 2007). Multiple cyclic dissolution/precipitation processes 

that affected the carbonates in the soils explain not only the formation of an isotopically 

light calcite, related to pedogenic-biogenic processes above the petrocalcic horizon 

(Magaritz and Amiel, 1980; Magaritz et al., 1981), but also the formation of an 

isotopically light dolomite, occurring below the petrocalcic horizon, in which the 

isotopic composition differs from that observed for the inherited dolomites (Fig. 4). The 

isotopic values of the pedogenic calcite co-existing with neoformed dolomite may be 

assigned to the maximum observed in the histograms of Figures 4A-B. Calcite isotopic 

18 O 

14


values are more negative than those of dolomite, as expected, because most of the 

calcite in soil is pedogenic in origin. The pedogenic dolomite precipitated in depth has a 

marked isotopic contribution from inherited dolomite, and therefore lower negative 

values than those measured for the pedogenic calcite are expected. 

Dolomite isotopic composition is not usually determined in soils, and much less in deep 

layers (“subsoil”). Until now, the occurrence of a pedogenic 13 C (and 18 O) signal in soils 

is associated exclusively with the calcite genesis. In the case studied, the neoformed 

dolomite, characterized by relatively negatives values, forms after partial dissolution of 

the calcite, with a partial “transference” of the biogenic isotopic signature. Therefore, 

calcite precipitation is interpreted here as an intermediate stage for dolomite 

precipitation, below the petrocalcic horizon, because of the carbonate dissolutionprecipitation 

cycles forming pedogenic calcite are also responsible of the formation of 

the pedogenic dolomites. 

Kinetic effects, such as diffusion transport, may also induce a certain fractionation in 

soils. The carbon isotope fractionation in soils caused by diffusion processes has been 

estimated to be around 4‰ (Cerling, 1984; Hesterberg and Siegenthaler, 1991). 

According to these authors 13 CO 2 diffuses into the atmosphere at lower rates than the 

lighter 12 CO 2 molecule. However this effect is clearly counteracted by the major 12 CO 2 

production related to root respiration and organic-matter decomposition in the upper 

horizons of soil, as evidenced clearly by the lower 

13 C values recorded for the 

pedogenic calcite of soil in relation to the pedogenic dolomite occurring in the subsoil. 

The kinetic effect discussed above resulting in 13 C enrichment would partly explain the 

shift from the very negative 

13 C at the C3 vegetal cover (~ -27‰ vs. V-PDB, Deines, 


1980) due to CO 2 from organic matter decomposition, toward values around -12‰ in 

the soil (Cerling, 1984; 1991; Reyes et al., 1998). Other kinetic processes such as 

intracrystalline diffusion of C and O within the dolomite structure might explain isotope 

exchange between the solid and the pore fluid in soils, and the observed fractionation in 

Fig. 4, without the need for an explanation involving the dolomite dissolutionprecipitation 

process. However, alternative explanations of these isotopic data based on 

intracrystalline diffusion processes should be discarded in our case because of: 1) 

According to the Arrhenius relationship between the diffusion rate and the temperature, 

intracrystalline diffusion is negligible at low temperature (Putnis, 2003). 2) Volume 

diffusion of C within the structure of the rhombohedral carbonates occurs by diffusion 

of carbonate anions, whereas the lattice diffusion of O is due to vacancy migration 

15


(Labotka et al., 2000); this explains why the diffusivity of carbon is 4 orders of 

magnitude smaller than the diffusivity of oxygen in calcite (Ferry et al., 2010). The 

latter point disagrees with our observations, where a broader range of variation is found 

for 

13 C compared to that of 

dolomite located at deeper levels. 

18 O for the cluster II (Fig. 4E) corresponding to the 

5.3. Physico-chemical constraints on the dolomite formation in the subsoil 

As shown, dissolution of inherited dolomite above petrocalcic constitutes the main 

source for the Mg 2+ enrichment in the soil solution necessary for dolomite 

neoformation. An exception to this is provided by the profiles with altered ophitic rocks 

in their subsoil, where minor dolomite was formed at great depth (below 200cm, Fig. 

2C). The montmorillonitic component of the smectites localized preferentially between 

corroded calcite grains and rhombohedral dolomite crystals (Table 2 and Figs. 6A-B) 

may constitute an additional source for Mg 2+ . Despite the occurrence of high Mg:Ca 

mole ratio, Mg rich calcite form only at very high salinity (Stanley and Hardie, 1998). 

However, for medium salinities, a low Mg content was expected in the calcite formed 

from solutions in soils. This is because Mg 2+ ions tend to remain in solution during 

carbonate precipitation, where they are thought to be heavily solvated (de Leeuw and 

Parker, 2001), due to strongly attractive Mg-water interactions. 

The formation of the petrocalcic horizon prevents rain-water infiltration into the subsoil. 

The observed dolomite enrichment in the subsoil probably occurred during the initial 

stages of the formation of the petrocalcic horizon (soft caliches), and before its 

hardening that led to the consolidation of the impermeable barrier. The high content of 

gravels and sands throughout the profile (Fig. 2E) favoured rainwater infiltration into 

the subsoil. Under these physical conditions, evapotranspiration processes affected not 

only the soil, but also deeper levels during the less rainy climatic episodes. The 

decreased water content in the subsoil diminished bulk flow of the grain-boundary fluid 

containing dissolved solutes. Under this regime, the fluid film is restricted to a 

monolayer of molecules adsorbed onto the smectite particles. To understand the role of 

the water molecules adsorbed onto smectites in the incorporation of Mg to the structure 

of the calcite, we dealt here with the transport of solute in a viscous solid media (Putnis 

et al., 1995; Fernández-Díaz et al., 1996; Sánchez-Navas et al., 2009). It can be deduced 

from Table 3 that the quantity of Mg incorporated into the structure of the calcite 

increased in viscous medium and that the incorporation of Mg into magnesian-calcite 


16


produced a lattice contraction of the rhombohedral carbonate, as well as a reduction of 

crystallinity. In addition, metastable phases also appeared when a viscous medium was 

used for the precipitation experiment (Fig. 9). Micro-textural data (Fig. 6) indicate that 

the precipitation of neoformed dolomite in depth after partial dissolution of pedogenic 

calcite was mediated by the smectites of the soils studied (Figs. 7 and 8). Clay minerals 

surrounding neoformed dolomite crystals acted as a catalytic material in the process of 

incorporation of Mg into the structure of the rhombohedral carbonates. This was 

because the transport of reacting material from dissolution sites to those of growth was 

the rate-determining process. Therefore, as suggested by our laboratory experiments, it 

is necessary to consider the constraints that transport properties may make on the 

incorporation of an element during crystal growth. Fig. 10A shows the concentration 

profile of Mg at the crystal-water interface during interface-controlled calcite growth. 

As suggested at the beginning of this section, the partition coefficient between the 

concentration of magnesium in carbonate crystal and the concentration of magnesium in 

bulk medium, [Mg] carb /[Mg] m , must be clearly less than 1. For a diffusion-controlled 

growth of dolomite, the concentration of Mg rises above [Mg] m at the crystal-medium 

interface ([Mg]i m ). If the partition coefficient remains constant, [Mg] carb increases with 

decreasing diffusion rate until, in the limit, [Mg] carb =[Mg] m . As a result the effective 

partition coefficient equals 1, and therefore most of the Mg of the medium is then 

incorporated into the structure of the rhombohedral carbonate (Fig. 10B). 

Previous experimental data and the occurrence in soils of high-magnesium calcite 

together with well-ordered dolomite (Whipkey and Hayob, 2008) suggest that our 

neoformed dolomite that precipitated in relation to smectites in the “subsoil” should be, 

at least initially, a “protodolomite” or “Mg-kutnohorite” with essentially disordered Ca- 

Mg distribution. 


6. CONCLUSIONS 

The present study detected the presence of a clear enrichment of dolomite beyond the 

lower limit of biological activity in semiarid soils. This limit coincides with the rooting 

depth of native perennial plants and, in the case study, is defined by the occurrence of 

the petrocalcic horizon. 

Isotopic data establish a bimodal distribution of 

13 C and 

18 O for dolomites, indicating 

the occurrence of an inherited and a neoformed dolomite. Isotopically light dolomite 

formed below the petrocalcic horizon after partial dissolution of the pedogenic calcite, 

17


which transferred the pedogenic (biological) 13 C and 18 O signature from the soil (s. str.) 

to subsoil. 

Before the formation of nonpermeable petrocalcic horizon, the high contents in gravel 

and sand fractions in the studied soils favoured infiltration and evapotranspiration 

processes in the subsoil. The moisture in the subsoil was a determining factor 

controlling the precipitation of Ca and Mg carbonates, the precursors of dolomite. The 

precipitation of the neoformed dolomite below the petrocalcic horizon was related to the 

presence of smectites. Smectite retained soil water and acted as a catalyst for the 

incorporation of Mg into the carbonate. 

The geochemical model proposed here for the incorporation of Mg to the carbonate in 

slow diffusive transport medium, such as a clayey-rich and water deficient 

environments, would also explain the close association of Mg-carbonates with clays 

observed on the surface of Mars (Ehlmann et al., 2008). 

This study documents the occurrence of changes in carbonate mineralogy in depth. 

Levels below the petrocalcic are traditionally considered as zones not affected by the 

pedogenic processes. 

In the soils studied, the dolomite content reached 43 kg m -2 (Díaz-Hernández, 2010), 

and therefore dolomite should be considered in the quantification of carbon stored in 

arid-semiarid soils. Pedogenic processes occurring in the subsoil, such as the dolomite 

formation, should be included in the carbon-soil balances. 

ACKNOWLEDGMENTS 

We acknowledge support from grant P11-RNM-7067 of the Junta de Andalucía 

(C.E.I.C.-S.G.U.I.T.) and the projects CGL2007-65572-C02-01/BTE, CGL-2009- 

09249, CGL2010-21257-C02-01 and CGL2012-32169 of the Ministerio de Ciencia e 

Innovación (MICINN, Spain). We would like to thank Drs. D. Martín-Ramos and F. 

Nieto for help with the XRD analyses, A. González, M.M. Abad and M.J. Guerrero for 

guidance with SEM and TEM studies and Dr. A. Delgado for his assistance with the 

stable isotope data. We thank the anonymous reviewers for their suggestions, which 

contributed to improve the manuscript. Mr. David Nesbitt revised the English 

manuscript. 


REFERENCES 

Al-Aasm, I.S., Taylor, B.E., South, B., 1990. Stable isotope analysis of multiple 

18


carbonate samples using selective acid extraction. Chemical Geology 80, 119-125. 

Arkley, R.J., 1963. Calculation of carbonate and water movement in soil from climatic 

data. Soil Science 96, 239-248. 

Arvidson, R.S., Mackenzie, F.T., 1999. The dolomite problem: control of precipitation 

kinetics by temperature and saturation state. American Journal of Science 299, 

257-288. 

Botha, G.A., Hughes, J.C., 1992. Pedogenic palygorskite and dolomite in a late 

Neogene sedimentary succession, nothwestern Transvaal, South Africa. Geoderma 

53, 139-154. 

Bui, E.N., Loeppert, R.H., Wilding L.P., 1990. Carbonate phases in calcareous soils of 

the western United States. Soil Science Society of America Journal 54, 520-529. 

Capo, R.C., Whipkey, C.E., Blachère, J.R., Chadwick, O.A., 2000. Pedogenic origin of 

dolomite in a basaltic weathering profile, Kohala peninsula, Hawaii. Geology 28, 

271-274. 

Cerling, T.E., 1984. The estable isotopic composition of modern soil carbonates and its 

relationships to climate. Earth and Planetary Science Letters 71, 229-240. 

Cerling, T.E., 1991. Carbon dioxide in the atmosphere: evidence from cenozoic and 

mesozoic paleosols. American Journal of Science 291, 377-400. 

Cerling, T.E., Quade, J., 1993. Stable carbon and oxygen isotopes in soil carbonates. In: 

Swart, P.K., Lohmann, K.C., Mckenzie, J., Savin, S. (Eds.), Climate Change in 

Continental Isotopic Records. Geophysical Monograph, vol. 78. American 

Geophysical Union, Washington, pp. 217-231. 

de Leeuw, N.H., Parker, S.C., 2001. Surface-water interactions in the dolomite problem. 

Physical Chemistry and Chemical Physics 3, 3217-3221. 

Deelman, J.C. 2005. Low temperature formation of dolomite and magnesite. Geology 

Series, version 2.1. Eindhoven: Compact Disc Publication. 

Deines, P., 1980. The isotopic composition of reduced carbon. In: A. Fritz and P. Fontes 

(Eds.). The terrestrial environment. Handbook of environmental isotope 

geochemistry: Elsevier Scientific Press, 329-434. 

Deleuze, M., Brantley, S.L., 1997. Inhibition of calcite crystal growth by Mg 2+ at 100°C 

and 100 bars: Influence of growth regime. Geochimica et Cosmochimica Acta 61, 

1475-1485. 

Díaz-Hernández, J.L. 2010. Is carbon storage in soils understimated. Chemosphere 80, 

346–349. 


19


Díaz-Hernández, J.L., Barahona-Fernández, E., Linares-González, J., 2003. Organic and 

Inorganic Carbon in soils of semiarid regions: a case study from the Guadix-Baza 

basin (Southeast Spain). Geoderma 114, 65-80. 

Díaz-Hernández, J.L., Juliá, R., 2006. Geochronological position of badlands and 

geomorphological patterns in the Guadix-Baza basin (SE Spain). Quaternary 

Research 65, 467-477. 

Díaz-Hernández, J.L., Martín-Ramos, J.D., López-Galindo, A., 2011. Quantitative 

analysis of mineral phases in atmospheric dust deposited in the south-eastern 

Iberian Peninsula. Atmospheric Environment 45, 3015-3024. 

Drees L.R., Wilding, L.P., 1987. Micromorphic record and interpretations of carbonate 

forms in the rolling plains of Texas. Geoderma 40, 157-175. 

Ehlmann, B.L., Mustard, J.F., Murchie, S.L., Poulet, F., Bishop, J.L., Brown, A.J., 

Calvin, W.M., Clark, R.N., Marais, D.J., Milliken, R.E., Roach, L.H., Roush, T.L., 

Swayze, G.A., Wray, J.J., 2008. Orbital Identification of Carbonate-Bearing 

Rocks on Mars. Science 322, 1828-1832. 

FAO-UNESCO, 1988. Soil Map of the World. Revised Legend. FAO, Roma. 

Fernández-Díaz, L., Putnis, A., Prieto, M., Putnis, C.V., 1996. The role of magnesium 

in the crystallization of calcite and aragonite in a porous medium. Journal of 

Sedimentary Research 66, 482-491. 

Ferry, J.M., Ushikubo, T., Kita, N.T., Valley, J.W., 2010. Assessment of grain-scale 

homogeneity and equilibration of carbon and oxygen isotope compositions of 

minerals in carbonate-bearing metamorphic rocks by ion microprobe. Geochimica 

et Cosmochimica Acta 74, 6517-6540. 

Fritz, P., Smith D.G.W., 1970. The isotopic composition of secondary dolomites: 

Geochimica et Cosmochimica Acta 34, 1161-1173. 

García-Aguilar, J.M., Martín J.M., 2000. Late Neogene to recent continental history and 

evolution of the Guadix-Baza basin (SE Spain). Revista de la Sociedad Geológica 

de España 13, 65-77. 

Ghebre-Egziabhier, K., St. Arnaud, R.J., 1983. Carbonate mineralogy of lake sediments 

and surrounding soils. 2. The Qu’Appelle Lakes. Canadian Journal of Soil Science 

63, 259-268. 

Gibert, L., Ortí, F., Rosell, L., 2007. Plio-Pleistocene lacustrine evaporates of the Baza 

Basin (Betic Chain, SE Spain). Sedimentary Geology 200, 89-116. 

Gile, L.H., Peterson, F.F., Grossman, R.B., 1966. Morphological and genetic sequences 


20


of carbonate accumulation in desert soils. Soil Science 1015, 347-360. 

Goddard, M.A., E.A. Mikhailova, C.J. Post, and M.A. Schlautman. 2007. Atmospheric 

Mg 2+ wet deposition within the continental United States and implications for soil 

inorganic carbon sequestration. Tellus B 59, 50-56. 

Goldsmith, J.R., Graf, D.L., 1958. Structural and compositional variations in some 

natural dolomites. Journal of Geology 66, 678-693. 

Hardie, L.A., 1987. Perspectives dolomitization: a critical view of some current views. 

Journal of Sedimentary Petrology 57, 166-183. 

Henderson, P., 1986. Inorganic geochemistry. Pergamon Press, Oxford, 356 pp. 

Hesterberg, R., Siegenthaler, U., 1991. Production and stable isotopic composition of 

CO 2 in a soil near Bern. Switzerland. Tellus 43B, 197-205. 

Hoefs, J., 2004. Stable isotope geochemistry, Springer-Verlag, Berlin Heidelberg, 244 

pp. 

IGME (1973-1999) Mapa Geológico de España, escala 1:50.000, Hojas 929, 930, 950, 

951, 972, 973, 993, 994, 1011 and 1012. 

Irwin, H., Curtis, C., Coleman, M., 1977. Isotopic evidence for source of diagenetic 

carbonates formed during burial of organic-rich sediments. Nature. 269, 209-213. 

Jobággy, E.G., Jackson, R.B., 2000. The vertical distribution of soil organic carbon and 

its relation to climate and vegetation. Ecological Applications 10, 423-436. 

Jones, M.B., Donnelly, A., 2004. Carbon sequestration in temperate grassland 

ecosystems and the influence of management, climate and elevated CO 2 . New 

Phytologist 164, 423-439. 

Kim, S.T., O´Neil, J.R., 1997. Equilibrium and nonequilibrium oxygen isotope effects 

in synthetic carbonates. Geochimica et Cosmochimica Acta 61, 3461-3475. 

Kohut, C., Dudas, M.J., Muehlenbachs, K., 1995. Authigenic Dolomite in a Saline Soil 

in Alberta, Canada. Soil Science Society of American Journal 59, 1499-1504. 

Labotka, T.C., Cole, D.R., Riciputi, L.R., 2000. Diffusion of C and O in calcite at 100 

MPa. American Mineralogist 85, 488-494. 

Lelong, F., 1969. Nature et gènèse des produits d'altération de roches cristallines sous 

climat tropical humide (Guyane Française). Mèmoires Science de la Terre, Nancy, 

14, 187 pp. 

Loeppert, R.H., Suárez, D.L., 1996. Carbonate and gypsum. In: Sparks, D.L. (Ed), 

Methods of Soil Analysis, part 3 Chemical Analysis, SSSA Book series, No 5, 

ASA & SSSA, Madison, WI, 437-474. 


21


Lohmann, K.Z., 1988. Geochemical patterns of meteoric diagenetic systems and their 

application to studies of paleokarst. In: James N.P., Choquette P.W. (Eds.), 

Paleokarst: Springer-Verlag, 58-80, Berlin. 

Lorimer, G.W., Cliff, G., 1976. Analytical electron microscopy of minerals. In: Electron 

microscopy in Mineralogy (Ed. H.R. Wenk), pp. 506-519. Springer-Verlag, 

Berlin. 

Magaritz, M., Amiel, A.J., 1980. Calcium carbonate in a calcareous soil from the Jordan 

Valley, Israel: Its origin as revealed by the stable carbon isotope method. Soil 

Science Society of American Journal 44, 1059-1062. 

Magaritz, M., Kaufman, A., Yaalon, D.H., 1981. Calcium carbonate nodules in soils: 

18 O/ 16 O and 13 C/ 12 C ratios and 14 C contents. Geoderma 25, 157-172. 

MAPA, 1989. Caracterización agroclimática de la Provincia de Granada. Ed. Secretaría 

General Técnica del MAPA. 197 pp. 

Martín-Ramos, J.D., 2004. Using XPowder®, a software package for powder X-Ray 

diffraction analysis. D.L.GR-1001/04. Spain. ISBN: 84-609-1497-6. 

Mikhailova, E.A., Post, C.J., 2006. Effects of land use on soil inorganic carbon stocks in 

the russian Chernozem. Journal of Environmental Quality 35, 1384-1388. 

Mikhailova, E.A., Post, C.J., Cihacek, L., Ulmer, M., 2009. Soil inorganic carbon 

sequestration as a result of cultivation in the mollisols, in Carbon sequestration 

and its role in the global carbon cycle, Geophysical Monograph Series 183, edited 

by B.J. McPherson and E.T. Sundquist, 129-133, A.G.U., Washington, D.C. 

Ming, D.W. 2002. In: Encyclopedia of Soil Science (eds R. Lal). Marcel Dekker, New 

York, 139-141. 

McCrea, J.M., 1950. On the isotopic chemistry of carbonates and a paleotemperature 

scale. Journal of Chemical Physics 18, 849-857. 

Mucci, A., Morse, J.W., 1983. The incorporation of Mg 2+ and Sr 2+ into calcite 

overgrowths: influences of growth rate and solution composition Geochimica et 

Cosmochimica Acta 47, 217-233. 

Peña, J.A., 1985. La Depresión de Guadix-Baza. Estudios Geológicos 41, 33-46. 

Podwojewski, P., 1995. The occurrence and interpretation of carbonate and sulphate 

minerals in a sequence of vertisols in New Caledonia. Geoderma 65, 223-248. 

Putnis, A., 2003. Introduction to Mineral Sciences. Cambridge University Press, 457 pp. 

Putnis, A., Prieto, M., Fernández-Díaz, L., 1995. Fluid supersaturation and 

crystallization in porous media. Geological Magazine 132, 1-13. 


22


Reeder, R.J., Sheppard, C.E., 1984. Variation in lattice parameters in some sedimentary 

dolomites. American Mineralogist 69, 520-527. 

Rech, J.A., Quade, J., Hart, W.S., 2003. Isotopic evidence for the source of Ca and S in 

soil gypsum, anhydrite and calcite in the Atacama Desert, Chile. Geochimica et 

Cosmochimica Acta 67, 575-586. 

Reyes, E., Pérez del Villar, L., Delgado, A., Cortecci, G., Núñez, R., Pelayo, M., Cózar, J., 

1998. Carbonatation processes at the El Berrocal analogue granitic system (Spain): 

mineralogical and isotopic study. Chemical Geology 150, 293-315. 

Rosenbaum, J., Sheppard, S.M.F., 1986. An isotopic study of siderites, dolomites and 

ankerites at high temperatures. Geochimica et Cosmochimica Acta 50, 1147-1150. 

Salomons, W., Mook, W.G., 1976. Isotope geochemistry of carbonate dissolution and 

reprecipitation in soils. Soil Science 122, 15-24. 

Salomons, W., Goudie, A., Mook, W.G., 1977. Isotopic composition of calcrete 

deposits from Europe, Africa and India. Earth Surface Processes and Landforms 3, 

43-57. 

Sánchez-Navas, A., Martín-Algarra, A., Rivadeneyra, M.A., Melchor, S., Martín-Ramos 

J.D., 2009. Crystal-growth behaviour in Ca-Mg carbonate bacterial spherulites. 

Crystal Growth and Design 9, 2690-2699. 

Shermann, D.G., Shultz, F., Always, F.J., 1962. Dolomitization of soils of the Red 

River Valley, Minnesota. Soil Science 94, 304-313. 

Sobecki, T.M., Karathanasis, A.D., 1987. Quantification and compositional 

characterization of pedogenic calcite and dolomite in calcic horizons of selected 

Aquolls. Soil Science Society of American Journal 51, 683-690. 

Sohet, K., Herbauts, J., Gruber, W., 1988. Changes caused by Norway spruce in 

ochreous brown earth, assessed by isoquartz method. Journal of Soil Science 39, 

549-561. 

Soil Survey Staff, 1999. Soil taxonomy, a basic system of soil classification for making 

and interpreting soil surveys. U.S. Department of Agriculture, Soil Conservation 

Service, Agriculture Handbook no. 436, Washington, D.C., 754 pp. 

Souchier, B., 1971. Evolution des sols sur roches cristallines à l'étage montagnard 

(Vosges). Mémoires du Service de la Carte Géologique d’Alsace et de Lorraine 

33, 134 pp. 

St. Arnaud, R.J., 1979. Nature and distribution of secondary soil carbonates within 

landscapes in relation to soluble Mg 2+ /Ca 2+ ratios. Canadian Journal of Soil 


23


Science 59, 87-98. 

St. Arnaud, R.J., Herbillon, A.J., 1973. Occurrence and genesis of magnesium-bearing 

calcites in soils. Geoderma 9, 279-298. 

Stanley, S.M., Hardie, L.A., 1998. Secular oscillations in the carbonate mineralogy of 

reef-building and sediment-producing organisms driven by tectonically forced 

shifts in seawater chemistry. Palaeogeography, Palaeoclimatology, Palaeoecology 

144, 3-19. 

Stevenson, B.A., Kelly, E.F., McDonald, E.V., Busacca, A.J., 2005. The stable carbon 

isotope composition of soil organic carbon and pedogenic carbonates along a 

bioclimatic gradient in the Palouse region, Washington State, USA. Geoderma 

124, 37-47. 

USDA-NRCS, 1996. Soil Survey Laboratory Methods Manual. S.S. Investigation 

Report No. 42, Version 3.0, 716 pp. 

Usdowski, E., Hoefs, J., 1993. Oxygen isotope exchange between carbonic acid, 

bicarbonate, carbonate, and water: A re-examination of the data of McCrea (1950) 

and an expression for the overall partitioning of oxygen isotopes between the 

carbonate species and water. Geochimica et Cosmochimica Acta 57, 3815-3818. 

Vera, J.A., 1970. Estudio estratigráfico de la Depresión de Guadix-Baza. Boletín 

Geológico y Minero 81, 429-462. 

Wacey, D., Wright, D.T., Boyce, A.J., 2007. A stable isotope study of microbial 

dolomite formation in the Coorong region, South Australia. Chemical Geology, 

244, 155-174. 

Whipkey, Ch.E., Hayob J.L., 2008. Textural and compositional evidence for the 

evolution of pedogenic calcite and dolomite in a weathering profile on the Kohala 

Peninsula, Hawaii. Carbonate Evaporite 23, 104-112. 

Wright, D.T., 1999. The role of sulphate-reducing bacteria and cyanobacteria in 

dolomite formation in distal ephemeral lakes of the Coorong region, South 

Australia. Sedimentary Geology 126, 147-157. 

Wright, V.P., 1992. Paleosol recognition: a guide to early diagenesis in terrestrial 

settings. In: Wolf K.H. and Chilingarian G.V. (Eds.), Diagenesis, III: Amsterdam 

(Elsevier), Developments in Sedimentol. 47, 591-619. 

Wright, V.P., Tucker, M.E., 1991. Calcretes. Ed. Blackwell, London, 352 pp. 


24


FIGURE CAPTIONS 

Fig. 1. A) Location of the study area showing the position of 81 profiles sampled. B) 

Cross section indicated in A showing the spatial relationship between the 

geomorphological units of the area. 

Fig. 2. A-C) Vertical distribution of organic and inorganic (calcite and dolomite) carbon 

in alluvial soils (PA, Fluvisols), badlands (PB, Regosols) and piedmonts (PC, 

Calcisols). Geomorphological units are also indicated in brackets. Arrows in Fig. C 

mark dolomite enrichment in depth for the diverse type of soils. D) Depth-wise (y-axis) 

variation of the organic carbon, calcite, and dolomite (inorganic carbon) content, 

expressed as an average percentage, in the study profiles of Guadix-Baza basin, drawn 

from the data of Table 1. E) Distribution with depth of diverse granulometric fractions 

of soils; only mean values calculated from the studied profiles has been represented 

(gravel + fraction


the dolomite-water system (Irwin et al., 1977) with the range of local meteoric water (-8 

to -9‰ vs. V-SMOW) and surface temperatures from 17 to 24ºC. 

Fig. 5. A-B) Back-scattered electron images showing precipitation of pedogenic calcite 

at 120cm depth (profile 46): Grain of quartz partly corroded and filled with calcite and 

probably siderite (A). Calcite nodule including dolomite relicts (B). Note the porous 

texture of the calcite precipitates in these images. C) Back-scattered electron 

micrograph of a large idiomorphic crystal of dolomite in an advanced stage of corrosion 

located at 120cm in depth. Note that this corrosion occurs in preferential plains, and 

progresses from the centre to outer edges. D) Small dolomite crystals with central voids; 

in addition, there are pedogenic calcites as rhyzocretions. E) Back-scattered electron 

image of a partly corroded large calcite grain surrounded by a mixture of clays and 

carbonates in a sample from 220cm in depth (profile 46). F) Inset of E in which we can 

see some dolomite grains surrounded by clays and corroded calcite grains. 

Fig 6. Electron images of a sample from 220cm depth (profile 46). A-B) Back-scattered 

electron images of rhombohedral dolomite crystals adjacent to corroded calcite grains 

and surrounded by clays. Detrital clays (chlorites) also occur in the area shown in Fig. 

B. C) Transmission electron microscopy image of a small subhedral dolomite crystal 

surrounded by smectite plus Si-rich amorphous substance. 

Fig. 7. A) General TEM image of clays and poorly crystalline material surrounding 

dolomite crystals in a sample from 220cm depth (profile 46). In addition to well-defined 

smectite packets (S), with a composition shown in Table 2, amorphous substance (A) 

with a mottled appearance and marked dark contrast also occurs in the poorly crystalline 

material located around the neoformed dolomite crystal (Dol). Other areas in the image 

lacking fringe contrast and of mottled aspect may correspond to smectites with a slight 

disorientation of the layers (Sd) or to a mixture of poorly crystalline smectites and 

amorphous substance (Table 2). B) Structural features of the small smectite packets, 

some of them composed of a few basal layers. The packets appear to be part of a (gellike) 

matrix formed by disoriented smectite-like flakes and a substance of relatively low 

atomic number lacking fringe contrast. 


26


Fig. 8 A) Powder X-ray diagram obtained from oriented aggregated (fraction 2) of the 

sample at 220cm depth (profile 46), with an indication of d values corresponding to the 

reflections of the observed mineral phases. (B) Comparison of air-dried and ethyleneglycol-treated 

diagrams, where the basal spacing of smectites is displaced from 1.46nm 

to 1.76nm after glycolation. 

Fig. 9. X-ray diffraction diagrams for magnesian calcites obtained in batch precipitation 

experiments (see Section 3.7 and Table 3) by using viscous (A) and liquid media (B). 

d 104 spacing of calcite varies from 0.301nm for the liquid medium to 0.295nm for the 

viscous media. In the viscous medium monohydrocalcite also precipitates together with 

the magnesian calcite. 

Fig. 10. Concentration profiles of Mg at the crystal-medium interface. A) Concentration 

distribution at an infinitely fast bulk-diffusion rate. B) Steady-state distribution at very 

slow surface-diffusion rate along the grain boundaries of clay minerals. ([Mg] carb , 

concentration of magnesium in carbonate crystal; [Mg] m , concentration of magnesium 

in bulk medium; [Mg] i m, concentration of magnesium at the carbonate crystal-medium 

interface; thickness of the diffusion boundary layer). Based on Henderson (1986). 

TABLE CAPTIONS 

Table 1: Arithmetic mean percentage (Mean in %) of the organic and inorganic (calcite 

and dolomite) carbon content and of the diverse particle-size fractions for the specified 

depths in the studied profiles (number of samples, N, for each depth is also indicated). 


Table 2. Analytical electron microscopy chemical analyses, expressed in atoms per 

formula unit, of smectite packets and mixture of smectites and amorphous substance. 

Normalization to 22 charges. 

Table 3. X-ray diffraction results for magnesian calcites obtained in batch precipitation 

experiments by using two types of media. In both cases the concentrations of the mixed 

soluble salts are the same: 50mM Na 2 CO 3 , 50 mM CaCl 2·6H 2 O, and 0.1 M 

MgCl 2·6H 2 O (starting pH 7.4). 

27


SPECIFIC ABBREVIATIONS 

a.s.l.: Above Sea Level 

AEM: Analytical Electron Microscopy 

apfu: Atoms Per Formula Unit 

BSE: Back Scattered Electrons 

DIC: Dissolved Inorganic Carbon 

EDX: Energy-Dispersion X-ray 

FESEM: Field-Emission Scanning Electron Microscope 

MCL: Meteoric Calcite Line 

MDL: Meteoric Dolomite Line 

OA: Oriented Aggregated 

TEM: Transmission Electron Microscopy 

XRD: X-ray diffraction 


28


Table 1 

Organic Carbon Inorganic Carbon Particle-Size (2mm) 

Depth Calcite Dolomite Clay Silt Sand Gravels 

(cm) N Mean (%) Mean (%) Mean (%) Mean (%) N Mean (%) N 

0 81 1.82 29.35 9.19 19.48 31.26 42.26 69 34.44 81 

20 81 1.13 37.60 8.42 19.82 29.59 50.60 55 34.83 68 

40 81 0.67 42.59 10.34 19.38 25.58 55.04 43 37.85 60 

60 81 0.46 42.23 12.95 15.51 28.27 56.22 43 43.14 57 

80 81 0.34 39.35 14.68 13.62 26.45 59.93 46 39.89 65 

100 81 0.25 35.90 15.24 12.04 23.43 64.53 43 37.73 65 

120 81 0.19 32.64 16.16 11.88 23.96 64.16 50 39.05 67 

140 75 0.16 31.41 16.83 11.46 21.49 67.05 42 35.55 65 

160 64 0.15 29.78 15.81 11.08 19.66 69.26 37 41.63 59 

180 55 0.12 31.25 13.28 10.94 22.82 66.23 35 41.14 48 

200 44 0.13 28.09 16.97 12.72 20.46 66.82 31 45.99 40 

220 27 0.11 33.63 18.33 12.12 17.99 69.89 17 53.06 24 

240 19 0.12 38.25 13.57 13.76 25.17 61.08 14 52.15 17 

260 12 0.14 37.31 16.26 10.18 15.63 74.19 6 51.91 9 

280 9 0.15 28.81 12.73 14.57 17.92 67.51 5 56.14 7 

300 8 0.18 26.78 12.77 18.89 19.18 61.94 4 59.46 6 

320 8 0.22 25.03 13.51 11.72 29.10 59.18 4 55.71 6 

340 4 0.22 31.14 14.00 9.20 15.66 75.14 2 58.12 3 

Total 892 546 747 

Table 2 

Smectite analyses 

Mixed analyses 

Si 3.79 3.88 3.83 3.78 4.16 4.17 4.12 4.41 

Al IV 0.21 0.12 0.17 0.22 -0.16 -0.17 -0.12 -0.41 


Al VI 1.55 0.84 1.57 1.69 1.55 1.28 1.43 1.53 

Fe (3+) 0.24 0.62 0.17 0.14 0.05 0.28 0.24 0.16 

Mg 0.34 0.53 0.22 0.14 0.31 0.48 0.26 0.21 

Ʃ VI 2.14 2.00 1.97 1.97 1.91 2.03 1.93 1.90 

Ca 0.03 0.03 0.00 0.02 0.05 0.05 0.05 0.03 

K 0.07 0.36 0.34 0.31 0.28 0.16 0.22 0.21 

Ʃ XII 0.10 0.40 0.34 0.33 0.33 0.21 0.28 0.24 

29


Table 3 

Type of medium d 104 (Å) FWHM (º2θ)* % Mg 

Liquid 3.01 0.3 10 

Viscous 2.95 1.26 25 

*FWHM: full width at half maximum (peak broadening measured for 104 reflection in the XRD 

diagrams of the Fig. 10 ). 


30


Fig. 1 


31



Fig. 2 

32


Fig. 3 


33


Fig. 4 


34


Fig. 5 


35



Fig. 6 

36


Fig. 7 


37


Fig. 8 


38


Fig. 9 


39


Fig. 10 


40


RESEARCH HIGHLIGHTS 

1. Neoformed dolomite appears in semiarid soils below the petrocalcic horizon 

2. An isotopically light dolomite forms after dissolution of pedogenic calcite 

3. Smectite acts as a catalyst for dolomite/protodolomite precipitation 

4. Pedogenic dolomite must be taken into account in carbon-soil balances 


41

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