Natural and Human-Induced Factors on the Accumulation and Migration of Pedogenic Carbonate in Soil: A Review

Abstract

As a principal part of the atmosphere–lithosphere interface, soil plays a key role in regulating the atmospheric CO₂ concentration and global climate. Comprising two major pools (carbonate in soils and bicarbonate in groundwater), soil inorganic carbon (SIC) is deemed as the primary carbon (C) sink and source in areas with low mean annual rainfall. SIC may originate from soil parent material or from the formation of secondary carbonate when divalent cations from an extraneous source are supplied. The latter may result in pedogenic carbonate (PC) formation, increasing soil C content and sequestering atmospheric carbon. Since the sequestration of atmospheric CO₂ through formation of pedogenic carbonate is gaining popularity as a method to support climate change mitigation efforts and to claim carbon credits, the mechanisms influencing the formation and migration of pedogenic carbonate need to be well understood. The present review provides an overview of the available literature on potential natural and anthropogenic factors influencing the pedogenic carbonate pool in soils. Firstly, the overall mechanisms of pedogenic carbonate formation, as well as the control factors, are described. Secondly, the impact of various land-use changes on pedogenic carbon pool modification is discussed. Then, the potential of stabilizing atmospheric CO₂ through PC formation and the challenges and techniques of tracking the formation of PC through engineered pathways in soils are explored. Finally, isotopic signature as a technique for distinguishing neo-formed carbonate in soil is scrutinized.

1. Introduction

Containing more than 2500 PgC over the first meter of the vertical profile, soils are regarded as the largest terrestrial and the third-largest C reservoir on Earth, after fossil fuel reserves and the oceans [1,2]. As carbon storage in soil impacts diverse aspects of life on Earth, it has been subjected to scientific interest from actors working in various disciplines, such as earth scientists, agronomists, and even policymakers [3]. The dynamics of soil carbon are a function of several parameters that influence C flux exchange between the air and soil system, and thus contribute to moderating atmospheric CO₂ levels [4,5,6].

Soil C comprises two discrete segments, soil organic carbon (SOC) and soil inorganic carbon (SIC), with the latter typically considered a more stable sink [2,5,7]. Decomposition of plant and animal residues, synthetic reactions and microbial activities are the main origins of the SOC pool [8,9]. On the contrary, SIC refers to mineral-based carbon, generated through the weathering of bedrock materials [8,10]. Turnover of SIC is believed to be 70–400 times longer compared to that of SOC [11]. While the SOC is estimated to be the major C pool in shallow (<1 m) soil globally [1], the SIC contribution has been found significant in arid and semi-arid areas, where the SIC/SOC ratio is in the range of 2~17 [12,13]. SIC is usually found in the form of calcite (CaCO₃) [14], although a significant contribution of dolomite (CaMg(CO₃)₂) has been reported in a few studies [5,15,16]. Recently, there has been greater research interest in evaluating the SIC pool and its impact on the global C budget [10,17,18], and this trend overcomes the prior neglect of the SIC pool, in the context of climate change mitigation, as it had been traditionally viewed as being less dynamic compared to its organic counterpart [7,19].

The precipitation–dissolution–re-precipitation regime (CaCO₃(s) + H₂CO₃(aq) ⇌ Ca²⁺ + 2HCO₃^–) is a primary factor in the distribution of the SIC pool over the vertical soil profile [20]. Monger et al. [17] described the SIC pool in humid and arid soil profiles using the terms ‘flushing’ and ‘non-flushing’, respectively. Flushing refers to a wet climatic condition favoring the downward migration of dissolved carbonate (Ca²⁺ + 2HCO₃^–, at near-neutral pH) to the groundwater system, whereas solid carbonate is more likely to accumulate in the non-flushing regime [17]. The global rate of sequestrated C influx (as HCO₃^–) to groundwater is approximately 0.2~0.36 Pg annually, with a residence time in the order of 100′s to 1000′s of years [10,14].

The SIC pool is categorized into two components, lithogenic carbonate (LC) and pedogenic carbonate (PC), with the former originating from the parent material of soil (e.g., limestone or carbonate-rich unconsolidated sediment) [10,18]. The dissolution–reprecipitation of LC, as well as the mechanism of solvation of CO₂(g) in soil moisture (as CO₃^2–/HCO₃^–) followed by its precipitation with Ca- and Mg-ions, lead to the formation of PC [17]. The formation of PC may modify soil characteristics, such as soil porosity, soil water content, and gaseous diffusion (e.g., of O₂(g)) [4].

The carbon exchange between soil and air is influenced by PC dynamics and hence is connected with the global climate [21]. Photosynthesis and root respiration are the major pathways for moving CO₂ from the atmosphere into the terrestrial ecosystem [10,12]. The main contributors to PC formation and accumulation include windblown dust [10,22], fertilizers [6], irrigation water [10,23], wet deposition [24,25], and bedrock weathering [10,22].

To accrue net SIC augmentation in the form of PC, the alkaline earth metal must be provided from an extraneous non-carbonate source; otherwise, the sequestered C equals the amount released to the atmosphere (e.g., during limestone calcination to produce quick lime) [17,26]. Hence, a narrower classification is needed to distinguish PC formation induced from an extraneous source and that resulting from recrystallization of pre-existing PC/LC; these PC sources are referred to as pedo-atmogenic and pedo-lithogenic carbonate, respectively [6,10,14]. A detailed description of the formation modes of both types can be found in Monger et al. [10].

The remainder of this paper reviews the mechanisms resulting in natural PC formation and the factors affecting its accumulation and migration. In addition, practices that modify carbonate pools (e.g., land-use change) and that improve PC formation (e.g., enhanced rock weathering) are discussed. Moreover, the application of isotopes as a tool for discriminating PC from LC is illustrated.

2. Pedogenic Carbonate Formation Mechanisms

Formation of PC involves a series of reactions: (1) dissolution of CO₂ and an alkaline earth metal-bearing mineral in the soil–water system; (2) migration of dissolved ions with soil water flow; and (3) precipitation of carbonate under favorable geochemical conditions (i.e., one that results in carbonate saturation and that depends on chemical properties of the aqueous medium) [4]. Depending on the soil water movement direction, this process may occur through several mechanisms, namely perdescendum (downward movement of solution) [27], perascendum (upward movement of the solution) [28,29], in-situ (absence of significant movement), and biologically induced mechanisms [4,30]. These processes have been explained in detail in published studies, including Zamanian et al. [4] and Li et al. [31].

Li et al. [31] argued that mechanisms, such as perdescendum or perascendum, may not fully explain the formation of carbonate nodules in loess sedimentary deposits. Instead, these authors attributed the pedogenesis of this carbonate to an “evapotransporative” mechanism, comprising evaporation at the surface and transpiration in the subsurface, with the former as the dominant process [31]. Díaz-Hernández et al. [15] identified various physiochemical mechanisms for the formation of surficial calcite (dominant over 0–180 cm) and deep dolomite (dominant deeper than 180 cm) owing to the weathering of basaltic rocks in a volcanic deposit site. Accordingly, it was proposed that calcite forms in an evaporative medium, whereas dolomite evolves in equilibrium with the groundwater system [15]. Laudicina et al. [32] described two pathways of PC formation involving biological factors: (i) biogenic pathways, where there is organic carbon enriched with a supplement of extraneous Ca²⁺ through rainfall; and (ii) dissolution and recrystallization of LC with a lower biogenic contribution.

PC occurs in several morphological forms, reflecting the dominance of certain abiotic/biotic processes [4,30]. Different morphological forms of carbonate may also be indicative of various sources (e.g., dissolution-recrystallization and atmospheric deposition), various PC formation mechanisms, as well as its emergence under certain humidity or vegetation regimes [16,29,30]. For instance, the dominance of biological activities may induce the formation of needle-fiber calcite crystals [33]. The distinctive morphological and physical characteristics of carbonate have been found useful for distinguishing several erosion-deposition sequences and soil type [34,35]. The formation of distinct types of PC may happen over different time scales, ranging between days and millennia [4].

3. Factors Affecting Pedogenic Carbonate Formation and Recrystallization

The formation of PC is a complex process, incorporating a wide variety of abiotic (e.g., climate, soil properties and characteristics, landscape) and biotic agents, as discussed and classified by Zamanian et al. [4] and illustrated in Figure 1. Each factor may affect morphology, leachability, and accumulation rate of PC as well as CO₂ solubility and Ca²⁺ availability in soil; all are influential aspects of PC dynamics in soil. Ferdush and Paul [10] reviewed how different biotic and abiotic factors may modify the SIC pool under elevated atmospheric CO₂ level conditions.

Optimized moisture content, suitable pH range (i.e., greater acidity is conducive to silicate mineral weathering, and near-neutral conditions are required for stability of solid carbonate), CO₂ supply, and availability of divalent cations (particularly Ca²⁺) are four essential conditions for PC formation [17,36]. The recrystallization rate of PC increases with CO₂ concentration elevation [37,38,39]. Zhao et al. [39] found that crystallization rate is a function of specific soil properties; the rate proportionally increases/decreases with salinity/pH, respectively. The latter effect is related to a negative correlation between CaCO₃ solubility with pH, wherein lower pH eases the dissolution, and thus recrystallization, of CaCO₃ [39]. They attributed the positive correlation between salinity and PC formation to a higher accumulation of exchangeable cations (e.g., Ca²⁺) in saline soil [39]. These authors also demonstrated that the increase of CO₂ partial pressure from 0.04 vol% (atmospheric) to 4 vol% results in a 100-fold increase in the carbonate recrystallization rate. Although the same trend was observed in the experiment conducted by Gocke et al. [37], several controls (e.g., fluctuation of CO₂ concentration in soil, temperature, root density, native crop regime, and soil texture) were also important as regulators of the rate of re-precipitation in natural soil systems. These factors, investigated in several studies as key aspects, are discussed next.

Hasinger et al. [36] demonstrated that the presence of a macroporous layer in soil facilitates CO₂ gas diffusion over the vertical and lateral profile of soil, acting as feedstock for PC formation. Investigating inter-seasonal variability of PC formation and seasonality of rainfall, Gallagher et al. [40] determined hydrological parameters (rainfall, evaporation, and soil water content) as the primary factors, rather than the mean annual temperature that was believed to previously be the main control. However, the vital role of temperature in PC dynamics in forest-steppe soils (alongside the mentioned hydrological parameters and biological activities in the growing season) has also been documented [41].

Select studies have highlighted the role of fine particles (nominally, clay) in the formation and accumulation of secondary carbonate [15,42]. According to Díaz-Hernández et al. [42], clay particles take up water that has migrated to the subsoil and therefore are conducive to the precipitation of carbonate. In contrast, the presence of gravel and sand particles can influence the hydrological regime and PC dynamics by facilitating the transport of water to subsoil layers [42]. Li et al. [13] also observed the same trend of SIC accumulation (with a PC isotopic signature) in a clay layer in the Yellow River Delta.

Microbial and root respiration, including the mineralization of organic carbon in the rhizosphere, are the main factors in supplying CO₂ in a subsoil medium [12,37]. Recent laboratory [43] and field-scale [44] studies showed that biotic processes (microbial activity) in deserts could induce biogenic PC formation through biomineralization with rates higher than abiotic pathways. Accordingly, the presence of bacteria may enhance the adsorption of Ca²⁺, by its attachment to the bacterial membrane, cell wall, and extracellular polymeric substances (EPS) layer, as highlighted by Liu et al. [43]. Furthermore, an increase in pH through metabolic activities (e.g., ureolysis, denitrification, ammonification, and sulfate reduction) can facilitate the precipitation of CaCO₃ [44]. A genus of termites, macrotermes can contribute to secondary carbonate formation through supplying Ca²⁺ that originates from plant tissues and biomineralization [45,46]. On the other hand, in the absence of biotic agents, Fa et al. [47] showed that in semi-arid deserts carbon remains in the dissolved inorganic carbon (DIC) phase, and migration of CO₂-originating carbon to groundwater is observed. Microorganisms have also been shown by Zhao et al. [38] to retard the recrystallization of PC due to partial surficial growth of microorganisms on CaCO₃ crystals, stabilizing the formed PC.

4. Pedogenic Carbonate Pool Modification under Various Management Practices/Land-Use Change

As PC formation is one of the main factors modifying the SIC pool, it is essential to comprehend its dynamics under different land uses [48,49]. Land-use change may come with changes in soil thermal and moisture regimes, consequently influencing the pool of PC in soil [14,28]. More than 30% of the anthropogenic contributions to atmospheric CO₂ concentration rise is due to land-use changes in the last 150 years [49]. A meta-analysis conducted by An et al. [49] denoted that various internal (e.g., pH, soil moisture, organic matter content) and environmental (e.g., mean annual temperature, rainfall) factors affect (positively or negatively) the SIC pool in the event of various land-use changes. In that study, SIC accumulation was observed in the conversion from grassland to farmland, and from sandy soil to forest soil, while conversion from farmland to grassland has reportedly led to an opposite correlation under certain scenarios [49]. Recently, Li et al. [50] assessed which of the key soil properties (e.g., bulk density, soil texture, pH, and nutrients availability) are influential on SIC and SOC dynamics through forestation and vegetation cover restoration, highlighting nutrients as the most relevant factor. It is important to note that the relationship between SIC and SOC is complex, and at times uncorrelated [49] (i.e., one can increase, while the other decreases); thus, land-use changes can have a different effect on total soil carbon than what changes in SIC or SOC alone imply. Practices at farmlands also vary substantially from crop to crop and region to region, and certain practices, such as nitrogen fertilization, have been shown to be measurably detrimental to the SIC pool [51].

Studies investigating PC pool modification in agricultural fields are abundant in the literature. Sanderman [14] reviewed three possible management practices influencing the SIC pool in the agricultural sector: irrigation, soil acidification, and liming. While the latter was identified as the undisputable C emission source (through dissolving in the soil–water system and efflux of CO₂ to the atmosphere), it remained unclear if the first two activities can act as either atmospheric C source or sink [14]. Depending on the objective, the studies investigating neo-formed carbonate in agricultural fields can be classified into two categories: (i) alteration of native soil in land-use change to agriculture (e.g., [23,29,52]) and (ii) investigating the impact of various agricultural management practices on the PC pool (e.g., [51,53,54]).

When establishing new agricultural fields, Nyachoti et al. [23] indicated a higher accumulation rate of pedogenic carbonate in agricultural land compared to natural soil in an arid environment. These authors identified the abundance of Ca²⁺ in irrigated water, as well as the provision of more DIC, as the main factors for a higher accumulation rate of carbonate in the cropland [23]. This study highlights the role of cation supplement from an extraneous source as a contributing factor in pedogenic carbonate formation. Conversion of shrubland to cropland has contributed to 0.87 t C·ha⁻¹·yr⁻¹ uptake in semi-arid croplands of the Yanqi Basin as reported by Wang et al. [27]. Kim et al. [54] observed a loss in the SIC pool due to land-use change to farmland, likely because of the leaching of carbonate in rain-fed and irrigated agricultural lands.

Focusing on agricultural management practices, Wang et al. [53] noted enhancement of PC and SIC pools over the vertical soil profile of arid and semi-arid croplands under various fertilizer management practices, with higher accumulation in soil amended with considerable amounts of organic matter. According to Bughio et al. [11], in deep layers of soils, an increase in plant biomass due to irrigation and fertilization practices leads to an increase in soil CO₂ levels (due to root and microbial respiration), providing an opportunity for PC formation in cases of Ca²⁺ and Mg²⁺ supplement. The biomass also contributes to evaporation, which is favorable for carbonate precipitation [11]. In another study [55], tillage and crop straw were effective for PC formation due to the rise in SOC level.

In addition to agriculture-related conversion/management, PC dynamics have been investigated in other forms of land-use change/rehabilitation, including forestation [56] and vegetation restoration [50]. In this respect, Gao et al. [56] noted that afforestation can enhance the accumulation of PC in arid soils over a 30-year period and within a 100 cm vertical soil profile.

SIC and SOC Correlation

There is a large body of studies reporting a correlation between SOC, SIC, and soil CO₂ level (e.g., [27,48,56]). When SOC level is high in the soil, the CO₂ concentration will increase due to organic matter decomposition, leading to the formation of HCO₃^– and H⁺ [57]. Accordingly, higher SOC stock leads to more PC formation, as was observed in arid lands of the Yanqi Basin [27], in northwest China [56], in arid cropland of China [55], in the Yellow River Delta [13], and in the Mu Us desert [50]. In contrast, progressive desertification and shifting to a less humid climate have been found to lead to a decrease in SOC and SIC in topsoil (e.g., 0–30 cm), contributing to the loss of PC stock [3,58].

While this SOC/SIC positive correlation has been largely detected in arid and semi-arid regions, the negative correlation of SOC and SIC (and PC) in regions with higher mean annual precipitation may be indicative of the leaching of carbonate to the subsoil and groundwater at depth [2,3]. Besides this, other environmental factors, such as soil properties and management practices, as well as native plant community, may lead to the negative correlation between SIC and SOC [9,59].

5. Sequestration of Atmospheric CO₂ as Pedogenic Carbonate through Engineered Pathways

The availability of cations, particularly Ca²⁺ and Mg²⁺, is believed to be the main limiting factor for the PC formation process [53,55]. Hence, the addition of Ca- and Mg-rich silicate minerals to soils (a.k.a. ‘engineered soils’) is among the popular techniques proposed to enhance the PC pool in agricultural [60,61], artificial [62], and urban soils [63]. These methods generally seek to reduce the atmospheric CO₂ level by enhancing the SIC pool, which under drier climates (and other amenable conditions) means augmenting the PC pool.

To accelerate the weathering of Ca- and Mg-rich silicate minerals in the soil-water system, the crushed form of these minerals is applied [60,64,65]. This process, known as enhanced rock weathering (ERW), consists of three steps: (1) atmospheric CO₂ solvation (e.g., in rainwater); (2) mineral dissolution (releasing base cations in the soil–water system); and (3) precipitation of carbonate (under favorable geochemical conditions) [60,66]. However, when new water is introduced, the newly formed carbonate might redissolve leading to the liberation of cations that speciate with bicarbonate and carbonate anions in the soil–water system as DIC [64,66,67]. This may result in either re-precipitation of PC in deeper layers of the soil or migration of DIC downward to groundwater [64,66]. In the latter scenario, the cations and DIC find their way to greater depths in the subsurface, discharge into freshwater systems, and eventually find the oceans as the ultimate sink [64,65].

Basalt (characteristically labradorite-rich (Ca,Na)(Al,Si)₄O₈) [61], olivine (characteristically forsterite-rich (Mg₂SiO₄)) [68] and wollastonite (CaSiO₃) [60,66] are Mg- and Ca-bearing minerals proposed for enhanced rock weathering in croplands. Several studies have highlighted the accumulation of SIC in field-scale experiments through the amendment of agricultural soils with these minerals [60,66]. In this respect, Haque et al. [60] demonstrated that the rate of SIC accumulation and sequestrated CO₂ increases over time when higher doses of wollastonite are spread onto the soil of commercial farms. In a controlled environment, Kelland et al. [61] showed the possibility of carbonate formation in soil columns (up to 4 t CO₂ after 5 years) due to the application of basaltic dust. ERW has also been found effective in enhancing plant growth apart from CO₂ sequestration [61,64,69]. From an agricultural and environmental perspective, the practice of ERW can serve as an alternative for liming agents, which are conventionally applied as a Ca source for soil structural improvement [6,14,60] but is known to generate atmospheric CO₂ emissions (0.273 PgC annually) globally [51,70].

Focusing on urban and artificial soils, Ca-mineral bearing materials, such as cement (bearing portlandite (Ca(OH)₂)) and plaster (bearing gypsum (CaSO₄·2H₂O)) are the main sources of calcium in construction and demolition (C&D) waste [71]. Synthetic soils, a combination of composted waste material and quarry fines enriched in silicate minerals, were found to be effective for trapping atmospheric CO₂ and enhancing the PC pool [62]. A summary of studies that have reported SIC increase due to ERW is shown in

Table 1. CO₂ fixation uptakes (on areal and total basis) reported in ERW studies conducted at field-scale.

Location	Land Use	Duration (y)	Depth of Collected Sample (cm)	SIC Content (net) (t CO2/ha)	Total Sequestrated C in the Study (t CO2)	Reference
Ontario, Canada	Agricultural field	0.41	15 cm	0.32	2.88	[60]
Ontario, Canada	Agriculture field	0.375	15 cm	0.40	0.20	[60]
Ontario, Canada	Agriculture field	1.25	15 cm	6.96	0.034	[66]
Newcastle, United Kingdom	C&D waste	1.5	10 cm	85.0	850	[63]
Newcastle, United Kingdom	Artificial soil	1	30 cm	17.6	0.53	[62]
Newcastle, United Kingdom	Brownfield	12	20 cm	59.0	151	[72]
Newcastle, United Kingdom	Natural soil	2	20 cm	123.6	0.15	[73]

.

As is the case with natural PC formation, various climatic and environmental factors affect SIC formation in engineered soils [60,72]. This includes pH [64,67,72], fineness of particles [72,74], hydrological parameters (redistributing SIC content over the soil profile) [60,72], and additive mineral characteristics [62].

Several prospective studies have assessed how ERW could contribute to carbon sequestration and climate change mitigation, globally. Using reactive transport modeling, Beerling et al. [67] have estimated the long-term (up to the year 2050) potential capacity of CO₂ removal in several countries through ERW of basalt in cropland areas, considering mineral demand and process costs. Recently, Haque et al. [75] propounded extending the ERW technique to urban roofs and balconies, labeled as ‘urban farming’. These authors estimated more than 13.6 Mt C could be sequestrated by applying this method to 1% of urban areas worldwide [75]. Manning et al. [76] proposed two techniques for enhancing PC formation at a global scale: (1) afforestation on lands with basaltic igneous rock subsurface structure; and (2) investment in the management of historically industrial lands that are enriched in materials whose composition is suitable for enhanced weathering. In addition to C&D waste, other industrial by-products, such as fly and bottom ash [71] and mining residues [77], can be suitable feedstocks for enhancing SIC stocks since they are rich in alkaline earth metals.

The practice of mining, crushing, and transporting minerals to the targeted land for performing ERW could be a source of CO₂ emissions into the atmosphere [78]. The life cycle assessment (LCA) study conducted by Lefebvre et al. [78] denoted that the application of basalt from quarries to cropland of Sao Paulo, Brazil could release, on average, 110 kg of CO₂(eq) per tonne of CO₂(eq) removed from the atmosphere. The potential unintended impact of minerals on ecosystems (including native land and water) is another factor that should be considered when ERW is applied on a terrestrial environment [66,79].

The current challenge of ERW implementation is the uncertainty with respect to the fate of neo-formed PC and its permanence. This points to the importance of developing predictive tools, such as a reactive transport modeling framework [67,80], in order to predict the long-term fate of PC in soil over decades to millennial time scales. It is also recognized that such complex carbon removal approaches require the participation of transdisciplinary actors to enable wide dissemination of its potential and best practices, and equitable and sustainable implementation, which in turn is critical for the inclusion of this approach into relevant global climate policies [81].

The application of biochar is another engineered management practice recently found to be effective in increasing PC stocks by capturing atmospheric CO₂ [82,83]. According to these studies, biochar can enhance PC formation through an increase in Ca²⁺ and pH as well as soil porosity (thus facilitating labile SOC decomposition and increasing soil CO₂ concentration) [82,83]. In a 10-year field experiment by Shi et al. [83], biochar application was found to be effective in increasing SIC over the 0–15 cm layer of soil.

A summary of the impact of engineered management practices and land-use change on carbonate formation based on the results of the reviewed studies is illustrated in Figure 2.

6. Isotopic Signatures: A Tool for Detecting and Quantifying Pedogenic Carbonate in Soil

There is a large body of literature demonstrating the success of the isotope technique in estimating PC for various objectives, such as paleoenvironmental reconstruction [84,85], neo-formed PC detection [62,86], land-use change [55,56], and labeling experiments [87]. The isotope signature of different types of carbonate varies in a soil horizon, reflecting the parent material, dissolution-reprecipitation cycle, and climatic factors [85,88,89]. Various isotope signatures including δ¹³C [62], ¹⁴C [29], δ¹⁸O [42,62], clumped isotopes [40], and U [23] have been used to study PC formation and accumulation. However, the focus of this section is on the first three signatures (δ¹³C, ¹⁴C, and δ¹⁸O) that are usually concerned with detecting and quantifying modern PC in the literature.

The δ¹³C signature of PC is influenced by processes occurring during the growing season in agricultural soil, including the decomposition of organic matter and the respiration of roots and microorganisms [4]. Accordingly, the δ¹³C value of PC is determined based on the contribution of C₃ and C₄ (photosynthesis carbon fixation classifications) plant endmembers, which have a mean value of –27‰ and –13‰, respectively [58]. Comparatively, the isotope value of the soil’s CO₂ is 4.5‰ enriched relative to reported biomass values (–27‰ and –13‰) in the literature [15]. As the PC precipitated from soil solution will be in equilibrium with soil CO₂ in terms of isotopic exchange, several studies have investigated the dynamics of PC in the atmosphere–soil–water system employing C-isotopic fractionation [15,36]. For instance, Hasinger et al. [36] described the evolution of atmospheric CO₂ to PC (through photosynthesis → SOM decay → root respiration → dissolution in meteoric water → precipitation) using C-fractionation.

Due to different origins, LC and PC demonstrate higher (near zero) and lower (negative values reflecting soil respiration) values, respectively [58]. Given the distinct isotope values of C₃ and C₄ biomass, many efforts have been made to relate the contribution of secondary carbonate to different vegetation regimes as well as environmental factors (e.g., [53,56]). As reported by Gao et al. [56], forestation leads to a progressive reduction in δ¹³C values (towards the C₃ and C₄ signature, aforementioned) over a 30-year time scale, reflecting neo-carbonate accumulation in the profile of the soil.

¹⁴C can also serve as a good proxy for distinguishing carbonate formed over the vertical soil profile and under different formation mechanisms (e.g., per-ascendum/descendum, in-situ) [19,29]. Using a ¹⁴C labeling experiment, Gocke et al. [87] noted most carbonate is leached from the 0–15 cm layer of soil, whereas the highest accumulation was observed at a depth of 20–35 cm, corresponding to the zone dominated by root biomass.

The δ¹⁸O isotope signature of local meteoric water is another useful isotope signature that has been used in high evaporative systems [42,89]. The PC formed under such conditions is correlated with enriched δ¹⁸O values [42].

Usually, a combination of the above-mentioned isotopic techniques is employed to investigate processes influencing PC formation and accumulation. Some key examples follow. Using isotopic signatures of δ¹³C and δ¹⁸O, it has been estimated that approximately 40% of neo-formed PC is formed through a photosynthetic process in synthetic [62] and urban soils [86]. These findings underscore the usefulness of using isotopic signatures to describe processes (e.g., photosynthesis) that lead to uptake of atmospheric CO₂ and their relative contributions to PC formation in soils amended with Ca and Mg bearing minerals. Mikhailova et al. [19] used both δ¹³C and ¹⁴C to quantify the PC fraction of total soil carbon and the percentage of modern carbon (pMC), respectively. They showed that pMC can vary between 73% at 30–40 cm and 25% for the 160–200 cm horizon of croplands in Russia [19]. Díaz-Hernández et al. [42] showed that tracking δ¹³C and δ¹⁸O is useful for describing pedogenesis of inherited and neo-formed dolomites formed under various geochemical conditions within the vertical soil profile. In another study, these authors explained the formation of deep dolomite vs. surficial calcite through distinguished δ¹⁸O values (with more negative values attributed to calcite) [15].

Nevertheless, estimating the PC pool using isotope techniques usually comes with a degree of uncertainty, arising from the decoupling between CO₂ production and carbonate precipitation spatially and temporally [53]. For instance, an increase in aridity leads to a decrease in the respiration rate of roots, which in turn results in the dominance of atmospheric CO₂ in soil pores [52,88]. In such a case, intrusion of atmospheric CO₂ (with a discreet isotopic signature) into soil pores is predictable, modifying the isotopic signature of soil CO₂ and soil carbonate (shifting to more positive values), particularly in topsoil [27,52,88]. Da et al. [88] noted that the contribution of atmospheric CO₂ in soil CO₂ could be up to 60% in a desert area.

7. Conclusions and Prospect

Detecting and quantifying PC is of significance as it is a dynamic pool in soils and could contribute to sequestrating atmospheric carbon and mitigating global warming. As discussed in this review, the occurrence and migration of PC in soil is a function of several hydrological and climatic factors. Moreover, soil physiochemical properties, microorganisms, and vegetation cover are the other factors that have been the focus of researchers investigating PC formation in the soil medium. All the above-mentioned factors control the morphological structure of secondary carbonate as well as the timing of PC formation.

Exploring the impact of land-use change on the PC pool has been a matter of interest to several researchers. This reflects the importance of tracking the PC reservoir as a part of the SIC pool particularly in arid and semi-arid environments where SIC is the dominant form of soil carbon. A change in the cation budget of soil and the presence of SOC, which affects soil CO₂ concentration, are the main driving factors governing PC modification in the case of land-use change practices.

As a technique to address the limitation of Ca²⁺ and Mg²⁺ cations in PC formation, enhanced rock weathering has been highlighted as an engineered pathway to increase the pool of secondary carbonate in agricultural and urban soils. Such a passive method has demonstrated great reliability in converting atmospheric carbon into secondary minerals, and it is recognized as a negative emissions technology so long as the manufacturing and transport of the comminuted rock supply has a low carbon footprint. There are still many unanswered questions about this technique, so it might gain more popularity as more large-scale studies are conducted in the field and a better understanding of the co-benefits and potential negative impacts emerge.

Verifying the atmospheric origin of neo-formed soil carbonate is one of the main challenges in evaluating the efficiency and feasibility of engineered carbon dioxide removal technologies. Although tracking isotopic signatures over time and space provides an opportunity to enhance comprehension of the mechanisms governing PC formation and transport over a certain period, the other concern is to ensure this neo-formed carbonate is stable for a long time after accumulation. Developing predictive tools for investigating the long-term fate of carbonate in soil and subsoil systems will be critical in this context.