INTRODUCTION
The majority of terrestrial vascular plants can be categorized as C3 or C4 plants according to their photosynthetic pathways (Hatch and Slack, Reference Hatch and Slack1970), with the former being the most common and ancestral type of photosynthesis (Maeda and Fernie, Reference Maeda and Fernie2021). Atmospheric CO2 diffuses into the mesophyll cells of C3 plants in response to concentration gradients and directly occupies carbon fixation sites within the cells (Hatch and Slack, Reference Hatch and Slack1970). This process is susceptible to photorespiration and leads to the consumption of photosynthate, which is enhanced by increasing temperature and light intensity (Sage et al., Reference Sage, Sage and Kocacinar2012). In contrast, the C4 photosynthetic pathway begins with the temporary fixation of intracellular CO2 into a four-carbon intermediate compound. These intermediate compounds are gathered and pumped into bundle sheath cells, releasing CO2 in high concentrations at carbon fixation sites (Hatch and Slack, Reference Hatch and Slack1970; Ehleringer and Monson, Reference Ehleringer and Monson1993). By spatially separating the carbon fixation and the Calvin cycle, C4 plants suppress photorespiration and increase the efficiency of photosynthesis (Sage et al., Reference Sage, Sage and Kocacinar2012). Therefore, C4 plants have a unique ecological significance due to carbon assimilation advantages under high temperature and light-intensity conditions (Pearcy and Ehleringer, Reference Pearcy and Ehleringer1984). The origins and expansion of C4 plants reflect combinations of significant environmental and climatic events (Ehleringer et al., Reference Ehleringer, Sage, Flanagan and Pearcy1991; Edwards et al., Reference Edwards, Osborne, Strömberg and Smith2010). Distinct carbon isotope discrimination mechanisms across photosynthetic pathways result in significant δ13C disparity between C3 and C4 plants (O'Leary, Reference O'Leary1981; Farquhar et al., Reference Farquhar, O'Leary and Berry1982, Reference Farquhar, Ehleringer and Hubick1989). The δ13C of soil organic matter and pedogenic carbonate are thus powerful proxies in documenting the variation of C3/C4 plants and have been widely used in paleovegetation reconstruction (Cerling et al., Reference Cerling, Harris, MacFadden, Leakey, Quade, Eisenmann and Ehleringer1997; An et al., Reference An, Huang, Liu, Guo, Clemens and Li2005; Edwards et al., Reference Edwards, Osborne, Strömberg and Smith2010; Sun et al., Reference Sun, Lü, Zhang, Wang and Liu2012; Carrapa et al., Reference Carrapa, Clementz and Feng2019).
Pedogenic carbonates and soil organic matter are the most common archives in terrestrial sediments, inheriting the δ13C signal of in situ vegetation (Cerling, Reference Cerling1984). The δ13C of pedogenic carbonates (δ13CCarb) should match the δ13C of coexisting organic matter (δ13CTOC), which has been well constrained by the theoretical model of Cerling (Reference Cerling1984) and its subsequent modifications (Davidson, Reference Davidson1995; Quade et al., Reference Quade, Rech, Latorre, Betancourt, Gleeson and Kalin2007). Pedogenic carbonates are generally formed by the dissolution and re-precipitation of primary carbonates by soil water and soil-respired CO2. Thus, the carbon isotope difference (Δ13C) between soil organic matter and pedogenic carbonates mainly comes from isotopic fractionation during two processes: (1) the diffusion of soil-respired CO2 within soil pores, producing a positive fractionation of 4.2–4.4‰ (Cerling, Reference Cerling1984; Davidson, Reference Davidson1995); and (2) the precipitation of carbonates from the soil solution, accompanied by isotopic exchanges between the carbonate phase and dissolved CO2 (Cerling, Reference Cerling1984; Cerling and Quade, Reference Cerling, Quade, Swart, Lohmann, Mckenzie and Savin1993). At the isotopic equilibrium state, carbonates are enriched in 13C relative to soil CO2 by 9–12‰ within a temperature range of 0–25℃ (Romanek et al., Reference Romanek, Grossman and Morse1992). The total isotopic difference between pedogenic carbonates and soil-respired CO2 is expected to fall within a range of approximately 14–17‰ (Cerling and Quade, Reference Cerling, Quade, Swart, Lohmann, Mckenzie and Savin1993; Quade et al., Reference Quade, Rech, Latorre, Betancourt, Gleeson and Kalin2007; Zamanian et al., Reference Zamanian, Pustovoytov and Kuzyakov2016). Surveys of modern soils have also confirmed this prediction, with δ13CCarb values generally being 14–17‰ more positive than those of coexisting δ13CTOC (Cerling and Quade, Reference Cerling, Quade, Swart, Lohmann, Mckenzie and Savin1993; Amundson et al., Reference Amundson, Stern, Baisden and Wang1998). Hence, the δ13C offset (Δ13C) of 14–17‰ is frequently applied in assessing the reliability of paired δ13C records in paleosol sequences (Cerling et al., Reference Cerling, Quade, Wang and Bowman1989; Cerling and Quade, Reference Cerling, Quade, Swart, Lohmann, Mckenzie and Savin1993; Sheldon and Tabor, Reference Sheldon and Tabor2009). However, a growing body of research has observed Δ13C exceeding this theoretical range (Rao et al., Reference Rao, Zhu, Chen and Zhang2006; Sanyal et al., Reference Sanyal, Sarkar, Bhattacharya, Kumar, Ghosh and Agrawal2010; Agrawal et al., Reference Agrawal, Sanyal, Sarkar, Jaiswal and Dutta2012; Vögeli et al., Reference Vögeli, Najman, van der Beek, Huyghe, Wynn, Govin, van der Veen and Sachse2017; Ghosh et al., Reference Ghosh, Sanyal, Sangode and Nanda2018), leading to uncertainty in comparing quantitative reconstructions of C4 vegetation between different archives.
Previous studies have mainly attributed the positive Δ13C anomaly to the contamination of pedogenic carbonates by external carbon sources, such as the mixing of detrital carbonates from dust sources or the infiltration of atmospheric CO2 into the soil profile (Cerling, Reference Cerling1984; Rao et al., Reference Rao, Zhu, Chen and Zhang2006; Sun et al., Reference Sun, Lü, Zhang, Wang and Liu2012; Da et al., Reference Da, Zhang, Li and Ji2020). However, this hypothesis cannot explain some observations in monsoon regions. In these cases, carbonates have a pedogenic origin, and the δ13CTOC value indicates a mixed C3/C4 vegetation (Wang and Follmer, Reference Wang and Follmer1998; Sinha et al., Reference Sinha, Tandon, Sanyal, Gibling, Stuben, Berner and Ghazanfari2006; Sanyal et al., Reference Sanyal, Sarkar, Bhattacharya, Kumar, Ghosh and Agrawal2010; Agrawal et al., Reference Agrawal, Sanyal, Sarkar, Jaiswal and Dutta2012; Shu et al., Reference Shu, Wang, Zhou, Ao, Niu, Wen and Li2021). The accumulation of soil organic matter continues with vegetation growth, and δ13CTOC represents the long-term average isotopic composition of the total plant biomass (Rao et al., Reference Rao, Guo, Cao, Shi, Jiang and Li2017). While the precipitation of pedogenic carbonates in monsoon regions is generally concentrated in warm seasons, δ13CCarb inherits the δ13C of soil-respired CO2 during the same period (Peters et al., Reference Peters, Huntington and Hoke2013; Kelson et al., Reference Kelson, Huntington, Breecker, Burgener, Gallagher, Hoke and Petersen2020). In a mixed C3/C4 ecosystem, C4 plants with improved photosynthetic efficiency under high temperatures make an enhanced contribution to rhizosphere respiration and cause 13C enrichment in respired CO2 (Breecker et al., Reference Breecker, Sharp and McFadden2009; Shimoda et al., Reference Shimoda, Murayama, Mo and Oikawa2009; Huth et al., Reference Huth, Cerling, Marchetti, Bowling, Ellwein and Passey2019). Consequently, the soil-respired CO2 during the pedogenic carbonate formation is biased toward respiration contributed by C4 plants, resulting in the C4 plant signal being overrepresented in the δ13CCarb record. This hypothesis requires further testing, because existing studies mainly focus on the theoretical calculation of fractionation processes and the cross-comparison of δ13C records in modern soil (Monger et al., Reference Monger, Cole, Buck and Gallegos2009; Montañez, Reference Montañez2013; Fischer-Femal and Bowen, Reference Fischer-Femal and Bowen2021; Sarangi et al., Reference Sarangi, Agrawal and Sanyal2021), with few considering paleosol sequences over long timescales.
In this study, we analyzed δ13CCarb and coexisting δ13CTOC records of the Shaozhai section in the central Chinese Loess Plateau (CLP) since the Pliocene. We made efforts to exclude the contribution of external carbon sources to carbonates by utilizing a range of morphological, mineralogical, and geochemical evidence. We also analyzed the isotopic composition of pedogenic carbonates to evaluate any seasonal bias in carbonate formation. Our aim was to test the hypothesis that variations in carbon sources are responsible for the positive Δ13C anomaly and to investigate the mechanism underlying this phenomenon. This study offers novel insights into comprehending the environmental implications of δ13C records associated with depositional biases and contributes to accurately utilizing the δ13C proxy in paleovegetation reconstruction.
MATERIALS AND METHODS
Study area and sampling
The modern CLP is dominated by mixed grassland ecosystems comprising C3 plants and C4 plants (Wang and Ma, Reference Wang and Ma2016). Most C4 plants are warm-climate grasses belonging to the Poaceae, such as Bothriochloa ischaemum and Setaria viridis, followed by species from the Chenopodiaceae, such as Salsola collina (Wang et al., Reference Wang, Feng, Han, Zhou, Tan and Su2008). On the other hand, C3 plants in this region include herbaceous plants, shrubs, and woody plants, among which herbaceous species from the Poaceae and Compositae have the widest distribution (Jiang and Ding, Reference Jiang and Ding2005; Liu et al., Reference Liu, Huang, An, Clemens, Li, Prell and Ning2005). The abundance of C3 shrubs gradually increases with aridity, and they prevail in the northwest of the CLP. Woody species of Pinus and Ulmus can be found in the southeastern region and deep gullies or low river terraces (Jiang et al., Reference Jiang, Cheng, Yang and Yang2013). The relative abundance of C4 plants in the modern CLP region increases from around 10% in the northwest to more than 60% in the southeast, as revealed by field investigations (An et al., Reference An, Huang, Liu, Guo, Clemens and Li2005; Rao et al., Reference Rao, Guo, Cao, Shi, Jiang and Li2017). This pattern corresponds with the increasing trend of modern annual temperature and precipitation from northwest to southeast.
The Shaozhai section (34°59.68′N, 107°48.61′E) is located in the Gansu Province of China, which lies in the central CLP and has a typical monsoon climate (Fig. 1). The climate of this region is characterized by warm, humid summers and cold, dry winters. July is the hottest month, and December is the coldest. During 1981–2010, the average summertime (June–August) temperature was 22.7°C, and the average wintertime (December–February) temperature was −0.7°C. The average annual temperature and precipitation were 10.4°C and 694 mm, respectively, with approximately 55% of the total annual precipitation occurring in summer (http://data.cma.cn).
Figure 1. Geographic map of the sections discussed in this study (left) and photographs of the Shaozhai section (right). The red star shows the location of Shaozhai section, and the orange circles indicate the other sections in the Chinese Loess Plateau with comparable published isotope records (An et al., Reference An, Huang, Liu, Guo, Clemens and Li2005; Da et al., Reference Da, Zhang, Wang, Balsam and Ji2015). Photographs of the Shaozhai section depict the interbedded loess–paleosol sequence in the upper part and the relatively uniform red clay formation in the lower part.
The Shaozhai section comprises a continuous aeolian deposit of 225 m from the base to the top, covering the middle Pliocene to the Quaternary (Cheng et al., Reference Cheng, Qiao, Liu, Li, Peng, Li, Qi and Wang2014; Lv et al., Reference Lv, Chunxia, Fu, Wu, Hao, Qiao and Guo2022). The Pliocene red clay is generally marked by stronger reddening, finer grain size, and lower porosity than the overlying Quaternary sequence, which exhibits a noticeable color difference between reddish paleosol layers and yellowish loess layers (Lv et al., Reference Lv, Chunxia, Fu, Wu, Hao, Qiao and Guo2022). However, loess and paleosol layers in the Pliocene red clay exhibit only a slight color difference, with paleosol being mainly reddish-brown or dark brownish and loess layers being yellowish-brown or brownish (Cheng et al., Reference Cheng, Qiao, Liu, Li, Peng, Li, Qi and Wang2014). The Pliocene red clay formation typically has a diffuse boundary between loess–paleosol layers, unlike the distinct boundary observed in the Quaternary consequences. These features indicate that the red clay formation has experienced stronger pedogenesis under a warm and stable climate, making weathered loess layers difficult to distinguish from soil units. Therefore, the red clay formation is divided into extremely thick soil units containing multiple soil B and carbonate nodule horizons (Supplementary Fig. 1).
Carbonate nodules in the upper part of the Quaternary sequence are generally embedded at the bottom of the paleosol layer or the top of the underlying loess layer. They have regular ellipsoidal or irregular tubular shapes with diameters ranging from 5 to 10 cm. Cracks and Fe-Mn films are observable within the nodules. The surfaces of some nodules are slightly porous and relatively light in color, while the inner parts are dense, hard, and dark in color. Nodules in the lower part of the Quaternary sequence have a slightly increased diameter, and their morphology is closer to a homogeneous massive structure. Carbonate nodules in the red clay formation are often scattered irregularly in the soil matrix or aggregated in a horizon at the top of weakly developed soil layers. They have uneven surfaces and diameters of 2–5 cm. Some nodules are composed of relatively pure micritic carbonate crystals, resulting in a white color (Supplementary Fig. 1).
To prevent potential contamination of the regolith, trenches (0.5–1 m deep) were dug into the section outcrop to expose fresh aeolian deposits. We applied different sampling strategies for collecting bulk samples and pedogenic carbonate samples. The soil matrix within each soil unit was selected for bulk samples to ensure representativeness and avoid interference from large nodules. Bulk samples were collected at intervals of 10 cm for magnetic susceptibility testing to complete stratigraphic correlation and establish the age model of this profile. When collecting pedogenic carbonate samples, we strove to cover each soil unit and choose ellipsoidal nodules with a complete shape, hard texture, and 2–7 cm diameter. We collected 137 pedogenic carbonate samples to analyze stable isotopes and carbonate mineralogical and elemental composition.
Magnetic susceptibility analysis and chronological framework
Bulk samples were dried and weighed into 10.0 g subsamples for low-frequency magnetic susceptibility testing using a Bartington MS2 magnetic susceptibility meter (Bartington Instruments, Oxfordshire, UK) at a frequency of 0.47 kHz. Each sample was independently measured three times, and a total of 2257 samples were tested. The magnetic susceptibility analyses were conducted at the Laboratory of Soil Structure and Mineralogy, Institute of Geology and Geophysics, Chinese Academy of Sciences.
A previous study of magnetostratigraphy has been carried out in this section, and the pedostratigraphic division of our sampling sequence is consistent with the published results (Qi et al., Reference Qi, Qiao, Liu, Wang and Peng2021). Therefore, the age control points were obtained by comparison of magnetic susceptibility records, and the continuous age–depth model was derived by the interpolation method of Kukla et al. (Reference Kukla, Heller, Liu, Tong and Liu1988). Our section spanned the interval from 0.1 to 4.5 Ma. The detailed comparison of magnetic susceptibility records is presented in Supplementary Figure 2.
Analyses of carbonate mineral and elemental composition
We randomly selected 16 carbonate nodules to analyze the carbonate mineral and elemental composition. These nodules were washed and dried, then mechanically crushed. The core fragments were ground in an agate mortar until the particle diameter was less than 45 μm. X-ray diffraction (XRD) analyses for mineral composition of powdered samples were conducted with a PANalytical diffractometer (Malvern Panalytical B.V., Almelo, Netherlands) with Ni-filtered Cu-Kα radiation (40 kV, 40 mA). For elemental analysis, an appropriate amount of powder was placed in a plastic centrifuge tube with 0.2 mol/L acetic acid overnight to dissolve the carbonate fraction. The Mn, Mg, and Ca in the dissolved product were measured by inductively coupled plasma mass spectrometry (PerkinElmer, Waltham, MA, USA) to obtain the elemental concentration. Analytical precision is less than 4% for Mn/Ca and Mg/Ca ratios. The XRD and elemental composition analyses were conducted at the Laboratory of Soil Structure and Mineralogy, Institute of Geology and Geophysics, Chinese Academy of Sciences.
Stable isotope analysis
A total of 137 pedogenic carbonate samples were washed repeatedly with pure water and scrubbed to remove soil attached to their surfaces. To avoid potential heterogeneity in micro-sampling, the nodule was cut, and the core part was broken into pieces, from which about 5 g of subsamples were ground to below 200 mesh and thoroughly mixed before further analysis. Samples were divided into two groups, one of which was used for the measurement of the carbon and oxygen isotope ratio (δ13CCarb and δ18OCarb) by a Thermo-Finnigan MAT-253 isotope ratio mass spectrometer (Thermo Scientific, Waltham, MA, USA) coupled with the Gasbench carbonate device (Thermo Scientific) reacted at 75°C with the oversaturated H3PO4. Following the conventional stable isotope analysis protocol (Thomas, Reference Thomas, Paul and Melillo1991), we performed one test on most samples and incorporated repeated tests on duplicate samples and different internal standards. Results are reported in per mil (‰) notation relative to the Vienna Pee Dee Belemnite (VPDB). Every eighth sample was treated as a replicate, and two laboratory internal standards with known isotopic values (GBW04416, δ13CVPDB = +1.61‰, δ18OVPDB = −11.59‰; GBW04417, δ13CVPDB = −6.06‰, δ18OVPDB = −24.12‰) were added for repeated testing. The standard deviation was thus obtained for the whole data set, and it was better than 0.15‰ for the δ13CCarb measurements and better than 0.20‰ for the δ18OCarb measurements.
The other group of samples was put into a centrifuge tube and reacted with 6 N HCl for 48 h to remove carbonates. Then, deionized water was added to rinse samples until neutral and freeze-dried. The residue of each sample was wrapped in a tin cup, and the carbon isotope ratio of organic matter occluded within the carbonate nodule (δ13CTOC) was determined by a MAT-253 isotope ratio mass spectrometer (Thermo Scientific) coupled with a FLASH EA1112 elemental analyzer (Thermo Scientific). We followed the same procedure used in the carbonate isotope analysis, and two laboratory internal standards (GBW04407, δ13CVPDB = −22.43‰; urea, δ13CVPDB = −49.1‰) were used to calibrate carbon isotope results. The standard deviation for measurements of δ13CTOC was better than 0.2‰ for the whole data set.
Given that the carbon isotopic composition of atmospheric CO2 (δ13Catm) has varied from the Pliocene to the present day, all raw carbon isotope ratios were corrected for secular changes in δ13Catm (Tipple et al., Reference Tipple, Meyers and Pagani2010; Schmitt et al., Reference Schmitt, Schneider, Elsig, Leuenberger, Lourantou, Chappellaz and Köhler2012). We used the mid-Holocene δ13Catm of −6.3‰ as the basis for this correction.
The stable isotope analyses were conducted at the Laboratory for Stable Isotope Geochemistry, Institute of Geology and Geophysics, Chinese Academy of Sciences.
RESULTS
The magnetic susceptibility result of the Shaozhai section is shown in Figure 2 and Supplementary Table 1. The loess layers have low magnetic susceptibility values and have undergone weak pedogenesis, whereas paleosol layers have high values due to the pedogenic formation of magnetic minerals.
Figure 2. Magnetic susceptibility and carbon and oxygen isotope composition of pedogenic carbonates (δ13CCarb and δ18OCarb), carbon isotope composition of organic matter occluded within carbonate nodules (δ13CTOC), and carbon isotope offset between δ13CCarb and δ13CTOC (Δ13C) of the Shaozhai section. Yellow and orange areas in the lithologic column represent the Quaternary loess and paleosol layers, while pink and red areas represent loess–paleosol units of the red clay formation. VPDB, Vienna Pee Dee Belemnite.
The δ13CCarb varied from −2.6‰ to −9.9‰, with a mean value of −7.4‰; the δ18OCarb varied from −7.3‰ to −12.0‰, with a mean value of −9.4‰; the δ13CTOC ranged from −20.4‰ to −26.2‰, with a mean value of −23.6‰ (Fig. 2, Supplementary Table 2). The long-term trend of δ13CTOC is consistent with that of δ13CCarb. Both of them reached the highest values in the entire sequence during 3–2.5 Ma, which coincides with the widespread expansion of C4 plants in the CLP during the late Pliocene (Ding and Yang, Reference Ding and Yang2000; An et al., Reference An, Huang, Liu, Guo, Clemens and Li2005; Suarez et al., Reference Suarez, Passey and Kaakinen2011; Sun et al., Reference Sun, Lü, Zhang, Wang and Liu2012; Zhou et al., Reference Zhou, Shen, Sun, Bird, Ma, Taylor, Liu, Peterse, Yi and Zheng2014). However, during this period (3–2.5 Ma), the increase in δ13CCarb values (approximately 4‰) was greater than that of the less-variable δ13CTOC (approximately 2‰), causing the isotopic offset between δ13CCarb and δ13CTOC (Δ13C) to exceed the theoretical range of 14–17‰ (Fig. 2).
The XRD results of nodule samples showed that carbonate fractions in nodules comprise mainly pure calcite minerals (Fig. 3); their Mg/Ca ratios ranged from 9.23 to 26.52 mol/mol, and Mn/Ca ratios ranged from 0.10 to 0.63 mmol/mL (Supplementary Table 3).
Figure 3. The mineralogical and elemental composition of carbonate nodules from the Shaozhai section. (A and B) The X-ray diffractograms of nodule samples and the carbonate mineral were entirely composed of calcites. (C) The Mg/Ca and Mn/Ca ratios of our nodule samples (blue circles), bulk carbonates from modern desert topsoils (gray circles and yellow shaded area), L1–S1 layers across the Chinese Loess Plateau (CLP; orange shaded area), the Holocene microcodium and biogenic carbonates (green shaded area). These elemental composition data of different carbonates were obtained from Li et al. (Reference Li, Chen and Chen2013) and Li and Li (Reference Li and Li2014).
DISCUSSION
Interpretation of the δ13CCarb record
The source and origin of carbonates in aeolian deposits are usually complex, and only the pedogenic carbonate that forms via recrystallization during pedogenesis can faithfully reflect the local vegetation conditions. Abiotic factors, such as the mixing of detrital carbonate from dust sources and the infiltration of atmospheric CO2 into the soil profile, could also drive increases in δ13CCarb and Δ13C values. Therefore, we examine these two influences separately in the following sections.
Influence of detrital carbonates
The coexistence of detrital and pedogenic carbonate in aeolian sequences is sometimes unavoidable due to dust accumulation and incomplete dissolution of soil parent materials (Rao et al., Reference Rao, Zhu, Chen and Zhang2006; Sheng et al., Reference Sheng, Chen, Ji, Chen, Li and Teng2008). To avoid the potential contamination of detrital carbonates, we carefully selected carbonate nodules with typical pedofeatures, including distinct ellipsoidal outer boundaries and a firm internal structure similar to the soil matrix (Fig. 3A and B). Such morphological features indicate that the carbonate nodules had a purely secondary origin (Barta, Reference Barta2011; Zamanian et al., Reference Zamanian, Pustovoytov and Kuzyakov2016).
Furthermore, striking mineralogical and geochemical differences between detrital and pedogenic carbonates can be used to identify potential contamination. In the study area, detrital carbonates carried by dust are mainly from marine carbonate strata in the arid regions of central Asia (Yang and Ding, Reference Yang and Ding2008; Chen and Li, Reference Chen and Li2011; Zhang et al., Reference Zhang, Lu, He, Xie, Wang, Zhang and Breecker2022). These carbonates inherit the chemical composition of seawater in geologic history and have experienced long-term burial metamorphism. The detrital carbonate usually includes a large proportion of dolomite minerals, with an elemental composition characterized by high Mg/Ca or Mn/Ca ratios (Li et al., Reference Li, Chen and Chen2013; Li and Li, Reference Li and Li2014; Meng et al., Reference Meng, Liu, Zhao, He, Chen and Ji2019). However, our XRD results indicate that the carbonate component of our samples is solely composed of calcite with no presence of dolomite observed (Fig. 3A and B). Because the dolomite dissolves more slowly than calcite during the soil weathering process, the disappearance of dolomite indicates the total dissolution of preexisting detrital carbonates (Tribble et al., Reference Tribble, Arvidson, Lane and Mackenzie1995; Meng et al., Reference Meng, Liu, Balsam, Li, He, Chen and Ji2015). Furthermore, our carbonate samples have particularly low Mg/Ca and Mn/Ca ratios (Fig. 3C), which are significantly lower than those of bulk carbonates from weakly weathered loess layers across the CLP and potential dust source regions, such as modern desert topsoil. In addition, their Mg/Ca and Mn/Ca ratios fall within the same range as those of other carbonates with unquestionable secondary origins, such as rhizoliths, pseudomycelia, and snail shells (Li et al., Reference Li, Chen and Chen2013).
The δ13C and δ18O of detrital carbonates are close to 0‰ (VPDB), significantly positive compared with those of pedogenic carbonates (Wang et al., Reference Wang, Zhang, Arimoto, Cao and Shen2005; Cao et al., Reference Cao, Zhu, Chow, Liu, Han and Watson2008; Sun et al., Reference Sun, Kutzbach, An, Clemens, Liu, Liu and Liu2015; Horton et al., Reference Horton, Defliese, Tripati and Oze2016). In the CLP region, the δ13C and δ18O of pedogenic carbonates are generally below −3‰ (VPDB) and −5‰ (VPDB), respectively (Ding and Yang, Reference Ding and Yang2000; Li et al., Reference Li, Sheng, Chen, Yang and Chen2007; Yang et al., Reference Yang, Ding, Wang, Tang and Gu2012; Luo et al., Reference Luo, Wang, An, Zhang and Liu2020). The present study observed that all δ13CCarb and δ18OCarb of our samples ranged from −2.6‰ to −9.9‰ and from −7.3‰ to −12.0‰, respectively (Fig. 4A and B), differing from the δ13C and δ18O values of detrital carbonates.
Figure 4. The δ13CCarb, δ18OCarb, and magnetic susceptibility in the Pliocene red clay (A and B) and Quaternary loess–paleosol sequence (C and D) in the Shaozhai section. VPDB, Vienna Pee Dee Belemnite.
Furthermore, the incomplete dissolution of detrital carbonates will be most evident in the Quaternary loess layers because of weak pedogenesis and rapid dust accretion (Guo et al., Reference Guo, Peng, Hao, Biscaye, An and Liu2004; Sun et al., Reference Sun, An, Clemens, Bloemendal and Vandenberghe2010). Consequently, loess layers will have a higher proportion of detrital carbonates and systematically higher δ13CCarb and δ18OCarb values than paleosol layers (Rao et al., Reference Rao, Zhu, Chen and Zhang2006; Liu et al., Reference Liu, Yang, Sun and Wang2011). According to the same principle, the δ13CCarb and δ18OCarb values of the Quaternary loess will be higher than those of the Pliocene red clay. However, this study showed that the Quaternary loess has lower δ13CCarb values and identical δ18OCarb values compared with the red clay formation. In addition, neither δ13CCarb nor δ18OCarb had differences in the range between loess and paleosol (Fig. 4).
Magnetic susceptibility has been suggested as a good indicator of the intensity of pedogenesis in the Quaternary loess–paleosol sequence, and it increases with increasing pedogenesis (Zhou et al., Reference Zhou, Oldfield, Wintle, Robinson and Wang1990). If detrital carbonates are present in significant amounts, δ13CCarb or δ18OCarb values and magnetic susceptibility will exhibit negative correlations, because weak pedogenesis accompanied by higher detrital carbonate contribution leads to an increase in the δ13CCarb or δ18OCarb (Rao et al., Reference Rao, Zhu, Chen and Zhang2006; Sun et al., Reference Sun, Yin, Crucifix, Clemens, Araya-Melo, Liu and Qiang2019). However, our data did not show this negative correlation in the Quaternary loess–paleosol sequence, while positive correlations between δ18OCarb and magnetic susceptibility were observed (Fig. 4C and D).
An increase in the content of fine-grained ferrimagnetic minerals created by pedogenesis is the primary reason for enhanced magnetic susceptibility in the Pliocene red clay (Hao et al., Reference Hao, Oldfield, Bloemendal and Guo2008, Reference Hao, Oldfield, Bloemendal, Torrent and Guo2009). Hence, the higher magnetic susceptibility still indicates relatively stronger pedogenesis (Nie et al., Reference Nie, King and Fang2007). If detrital carbonates are present in the Pliocene red clay, negative correlations between δ13CCarb or δ18OCarb and magnetic susceptibility will also be observed. However, similar to results from the Quaternary loess–paleosol sequence, neither δ13CCarb nor δ18OCarb is negatively correlated with magnetic susceptibility, whereas δ18OCarb is positively correlated with magnetic susceptibility in the Pliocene red clay (Fig. 4A and B).
In summary, mineralogical and geochemical evidence strongly supports the purely pedogenic origin of our carbonate samples and excludes detrital carbonate contamination. Previous studies were conducted on some sections in the central and southern parts of the CLP, which are at the same latitude as the Shaozhai section and far from the northern sandy lands (Yang et al., Reference Yang, Ding, Wang, Tang and Gu2012; Zhang and Liu, Reference Zhang and Liu2013). These studies also suggested that the contribution of detrital carbonates is negligible in carbonate nodules, which is consistent with our conclusion.
Influence of atmospheric CO2
The contribution of 13C-enriched atmospheric CO2 in the soil atmosphere can be identified by enhanced carbon isotope fractionation between δ13CCarb and δ13CTOC (Quade et al., Reference Quade, Rech, Latorre, Betancourt, Gleeson and Kalin2007). Modern surveys of pedogenic carbonates and soil organic matter have observed that Δ13C anomalies occurred at different depths in the soil profile (Retallack, Reference Retallack2009; Montañez, Reference Montañez2013; Myers et al., Reference Myers, Tabor, Jacobs and Bussert2016). The infiltration of atmospheric CO2 into the soil profile is primarily limited to the surface layer and does not reach soil layers below 50 cm (Cerling, Reference Cerling1984; Quade et al., Reference Quade, Rech, Latorre, Betancourt, Gleeson and Kalin2007). Therefore, Δ13C anomalies in deeper soils are expected to be unrelated to the infiltration of atmospheric CO2 (Fig. 5). Typical soil types in the modern CLP, such as alfisols, aridisols, and mollisols, generally have a depth to the carbonate nodular horizon (Bk) greater than the reach of atmospheric CO2 (Feng and Wang, Reference Feng and Wang2005; Yang et al., Reference Yang, Ding, Wang, Tang and Gu2012). Moreover, the depth of pedogenic carbonate accumulation is positively correlated with precipitation in semiarid areas (Retallack, Reference Retallack2005; Zamanian et al., Reference Zamanian, Pustovoytov and Kuzyakov2016), thus further reducing the potential contribution of atmospheric CO2.
Figure 5. Carbon isotope fractionation with depth in different types of modern soil profiles. The gray shaded area indicates the theoretical range (14–17‰) of Δ13C. Carbon isotope data were compiled from Retallack (Reference Retallack2009), Montañez (Reference Montañez2013), and Myers et al. (Reference Myers, Tabor, Jacobs and Bussert2016).
If atmospheric CO2 infiltrates the interior of the soil profile, then periods with higher atmospheric CO2 concentrations should be more prone to such infiltration than periods with lower CO2 concentrations. However, the highest Δ13C values in this study did not occur during the Pliocene warm period (~3.3 Ma), which had the highest CO2 concentrations. On the contrary, they occurred during the late Pliocene to the Early Quaternary period, when atmospheric CO2 dropped sharply (Fig. 6A). Other studies (Fig. 6) have also found that the highest anomalies did not occur during geologic periods when atmospheric CO2 concentrations were highest (An et al., Reference An, Huang, Liu, Guo, Clemens and Li2005; Bereiter et al., Reference Bereiter, Eggleston, Schmitt, Nehrbass-Ahles, Stocker, Fischer, Kipfstuhl and Chappellaz2015; Da et al., Reference Da, Zhang, Wang, Balsam and Ji2015; Rae et al., Reference Rae, Zhang, Liu, Foster, Stoll and Whiteford2021).
Figure 6. Comparison of atmospheric CO2 concentration and δ13CCarb and δ13CTOC records from different sections. (A) Δ13C from the Shaozhai section and the atmospheric CO2 concentration from the Pliocene; the black dotted line represents the theoretical upper limit of Δ13C (17‰). (B) The δ13CCarb and δ13CTOC records of the Shaozhai section. The individual data points with LOESS smoothing in thick solid lines and confidence intervals indicated by shaded areas. The gray shaded bars represent the loess layers L1, L2, L9, and L15. (C) Δ13C from the Baoji, Lingtai, and Lantian sections. (D and E) The δ13CCarb and δ13CTOC records of each section shown in C. Baoji, Lingtai, and Lantian data were collected from An et al. (Reference An, Huang, Liu, Guo, Clemens and Li2005) and Da et al. (Reference Da, Zhang, Wang, Balsam and Ji2015). The atmospheric CO2 concentration reconstruction results are compiled from Bereiter et al. (Reference Bereiter, Eggleston, Schmitt, Nehrbass-Ahles, Stocker, Fischer, Kipfstuhl and Chappellaz2015) and Rae et al. (Reference Rae, Zhang, Liu, Foster, Stoll and Whiteford2021).
In addition, loose and porous soil layers are typically more prone to existing atmospheric CO2 infiltration than dense soil layers. The Quaternary loess layers have coarser particles and higher porosity than the paleosol layers and Pliocene red clay, making them more vulnerable to atmospheric CO2 infiltration and more likely to produce pronounced Δ13C anomalies. However, the typical loess layers such as L1, L2, L9, and L15, with the highest dust accumulation rate, higher soil matrix porosity, and lower vegetation productivity, did not show the expected Δ13C anomalies (Fig. 6A and C).
If such infiltration did occur, δ13CCarb and δ13CTOC would unlikely show a consistent trend or strong correlation, because atmospheric CO2 has a much more positive δ13C composition than CO2 generated by plant and soil respiration (Quade et al., Reference Quade, Rech, Latorre, Betancourt, Gleeson and Kalin2007). The contribution of atmospheric CO2 tends to homogenize δ13CCarb to a fixed value independent of δ13CTOC. However, this and previous studies have shown a positive correlation between the δ13CCarb and δ13CTOC for each section (Fig. 6B, D, and E).
Therefore, the evidence from these four aspects allows us to conclude that Δ13C anomalies are unrelated to atmospheric CO2 infiltration.
Seasonality of C4 plant growth and carbonate precipitation controlled δ13CCarb anomaly
The low biomass of C4 plants in the Shaozhai section before 3 Ma is suggested by the δ13CCarb and δ13CTOC with average values of −7.3‰ and −23.5‰, respectively, which are in agreement with published pollen and δ13C records from the CLP (Ding and Yang, Reference Ding and Yang2000; Jiang et al., Reference Jiang, Han and Liu2001; An et al., Reference An, Huang, Liu, Guo, Clemens and Li2005; Ma et al., Reference Ma, Wu, Fang, Li, An and Wang2005; Li et al., Reference Li, Fang, Wu and Miao2011; Rao et al., Reference Rao, Zhang, Xue, Xu and Liu2012a). During this period, our δ13CCarb and δ13CTOC records showed a gradual positive trend with small fluctuations (Fig. 2). They maintained the congruent trend and Δ13C values within the expected range (14–17‰) in glacial cycles and on longer timescales. This demonstrates the consistency in the composition and variations of C3/C4 vegetation between the carbonate precipitation period and soil organic matter accumulation period.
At the Pliocene–Quaternary transition, δ13CCarb and δ13CTOC experienced a pronounced positive shift and reached their highest values at approximately 2.8 Ma, with the rise in δ13CCarb being more abrupt than that of δ13CTOC. Both indicate a considerable expansion of C4 plants, which is also observed in contemporaneous sections throughout the CLP (Ding and Yang, Reference Ding and Yang2000; Jiang et al., Reference Jiang, Han and Liu2001; An et al., Reference An, Huang, Liu, Guo, Clemens and Li2005; Suarez et al., Reference Suarez, Passey and Kaakinen2011; Rao et al., Reference Rao, Zhang, Xue, Xu and Liu2012a). The magnitude of δ13CCarb positive shift (ca. 4‰) and peak values (ca. −3‰) at the Shaozhai section are similar to what has been reported in previous studies (Ding and Yang, Reference Ding and Yang2000; An et al., Reference An, Kutzbach, Prell and Porter2001; Jiang et al., Reference Jiang, Han and Liu2001; Suarez et al., Reference Suarez, Passey and Kaakinen2011). However, this study and other δ13CTOC records exhibit a relatively minor positive excursion of 2.0‰ to 2.5‰ (An et al., Reference An, Kutzbach, Prell and Porter2001; Rao et al., Reference Rao, Zhang, Xue, Xu and Liu2012a).
The negative correlation between δ18OCarb and δ13CCarb became apparent after 2.8 Ma (Fig. 7B and C). Because the isotopic composition of pedogenic carbonates is inherited from meteoric waters and terrestrial vegetation, the relationship between δ18OCarb and δ13CCarb reflects the vegetation response to climate changes (Cerling and Quade, Reference Cerling, Quade, Swart, Lohmann, Mckenzie and Savin1993; Yang et al., Reference Yang, Ding, Wang, Tang and Gu2012; Kovda et al., Reference Kovda, Morgun and Gongalsky2014; Bayat et al., Reference Bayat, Karimzadeh, Eghbal, Karimi and Amundson2018). In the modern CLP, the monsoon climate results in a concentration of precipitation during warm seasons. Therefore, the oxygen isotope composition of meteoric waters during warm seasons shows a notable negative shift due to the rainfall effect (Vuille et al., Reference Vuille, Werner, Bradley and Keimig2005), producing a corresponding negative shift in the δ18OCarb (Zhang et al., Reference Zhang, Li, Yan and An2018). Although C4 plants maintain high photosynthetic efficiency under high temperatures and water-stressed conditions (Pearcy and Ehleringer, Reference Pearcy and Ehleringer1984; Sage et al., Reference Sage, Sage and Kocacinar2012), severe droughts in growing seasons are unfavorable for their growth, as indicated by the scarcity of C4 plants in Mediterranean climates (Rao et al., Reference Rao, Chen, Zhang, Xu, Xue and Zhang2012b). High temperatures and increased precipitation during growing seasons, as a result of the intensification of seasonality, benefit C4 plants by allowing them to fully utilize their adaptive advantages and causing a positive shift in δ13CCarb. Previous surveys of modern vegetation and the synchronous migration of the C4 plant distribution boundary and the monsoon rainfall belt have validated this hypothesis (An et al., Reference An, Huang, Liu, Guo, Clemens and Li2005; Von Fischer et al., Reference Von Fischer, Tieszen and Schimel2008; Yang et al., Reference Yang, Ding, Li, Wang, Jiang and Huang2015; Munroe et al., Reference Munroe, McInerney, Guerin, Andrae, Welti, Caddy-Retalic, Atkins and Sparrow2022). Therefore, pedogenic carbonates generally have a distinct negative correlation between δ13CCarb and δ18OCarb in the CLP (Yang et al., Reference Yang, Ding, Wang, Tang and Gu2012). This specific correlation suggests seasonal biases in both the metabolism of C4 plants and pedogenic carbonate formation, and it precludes the potential contamination of detrital carbonates and atmospheric CO2.
Figure 7. Modern meteorological data in the Chinese Loess Plateau and δ13CCarb, δ18OCarb, and Δ13C records of the Shaozhai section during the late Pliocene C4 expansion event. (A) The monthly average temperature (red), precipitation (blue), and monthly weighted precipitation δ18Op (black) of GNIP site Xi'an (IAEA/WMO, 2006). (B) The δ13CCarb–δ18OCarb correlation of pedogenic carbonates (P = 0.01, r = −0.41). Red circles indicate samples exhibiting a negative δ13CCarb–δ18OCarb correlation with high Δ13C values in the red dashed box of C. (C) The δ13CCarb, δ18OCarb, and Δ13C variation of the Shaozhai section; the black dotted line represents the theoretical upper limit of Δ13C (17‰).
Simultaneous with the rapid rise in our δ13CCarb record at 2.8 Ma, the Δ13C value of the Shaozhai section starts to show abnormally high values exceeding 17‰ because of larger variations in δ13CCarb compared with δ13CTOC (Fig. 7C). Given that we have excluded external contamination, soil CO2 produced by plant root respiration and decomposition of soil organic matter is considered the primary carbon source that contributes to pedogenic carbonates (Cerling, Reference Cerling1984; Zamanian et al., Reference Zamanian, Pustovoytov and Kuzyakov2016). Consequently, the increased range of δ13CCarb variation implies that greater vegetation changes occurred during periods of carbonate precipitation than those of organic matter accumulation. During warm seasons, root respiration dominates the soil CO2 due to vigorous plant metabolism (Hanson et al., Reference Hanson, Edwards, Garten and Andrews2000; Bond-Lamberty and Thomson, Reference Bond-Lamberty and Thomson2010). The respiration flux contributed by C3 and C4 plants at certain times probably does not match their respective proportions in annual plant biomass, because the competitive advantage of C3 and C4 plants differs seasonally (Lai et al., Reference Lai, Riley, Owensby, Ham, Schauer and Ehleringer2006; Shimoda et al., Reference Shimoda, Murayama, Mo and Oikawa2009). In mixed C3/C4 grassland ecosystems, the contribution of C4 plants to root respiration is greater in warm seasons than in other seasons due to their higher temperature tolerance and water-use efficiency (Pearcy and Ehleringer, Reference Pearcy and Ehleringer1984). High-resolution field observations have also confirmed apparent seasonal cycles in the carbon isotope composition of soil-respired CO2 (Breecker et al., Reference Breecker, Sharp and McFadden2009, Reference Breecker, McFadden, Sharp, Martinez and Litvak2012; Huth et al., Reference Huth, Cerling, Marchetti, Bowling, Ellwein and Passey2019). Therefore, the δ13C of soil CO2 in warm seasons overrepresents the abundance of C4 plants as a result of their increased metabolism, which is documented by pedogenic carbonates. We speculate that this mechanism could explain high Δ13C values observed in the present study after 2.8 Ma.
From 2.8 to 1.5 Ma, the δ13CCarb and δ13CTOC in the Shaozhai section started to decrease and reached a minimum, indicating a decline in C4 plants. This decline also occurs in the central and western parts of the CLP and is believed to be primarily caused by the long-term cooling trend associated with enhanced Northern Hemisphere glaciation (Ding and Yang, Reference Ding and Yang2000; An et al., Reference An, Huang, Liu, Guo, Clemens and Li2005; Suarez et al., Reference Suarez, Passey and Kaakinen2011; Sun et al., Reference Sun, Lü, Zhang, Wang and Liu2012). According to surveys of δ13C end-member values of modern C3/C4 plants and correction for secular changes in atmospheric δ13CCO2 (+1.7‰), we select −25‰ and −8‰ as the most conservative thresholds for δ13CCarb and δ13CTOC indicating pure C3 plants (Wang et al., Reference Wang, Feng, Han, Zhou, Tan and Su2008; Tipple et al., Reference Tipple, Meyers and Pagani2010; Rao et al., Reference Rao, Guo, Cao, Shi, Jiang and Li2017; Jiang et al., Reference Jiang, Wu, Li, Lin and Yu2019), both of which are widely accepted and applied in paleovegetation reconstruction (Cerling et al., Reference Cerling, Harris, MacFadden, Leakey, Quade, Eisenmann and Ehleringer1997; Tipple and Pagani, Reference Tipple and Pagani2007; Edwards et al., Reference Edwards, Osborne, Strömberg and Smith2010). Therefore, our δ13CCarb and δ13CTOC values are close to the C3 plant end-member from 1.9 to 1.5 Ma, suggesting insufficient evidence to support the presence of C4 plants. Given the positive δ13C excursion in our record and the distinct δ13CCarb–δ18OCarb correlation (Fig. 7), anomalous Δ13C values from 2.8 to 1.9 Ma can be confidently correlated to the emergence and enhanced metabolism of C4 plants. Pedogenic carbonates inherit the δ13C of soil-respired CO2 with a positive bias during the growing season for C4 plants, leading to abnormally high Δ13C values that exceed the theoretical range.
After 1.5 Ma, our δ13CCarb and δ13CTOC records maintain a good consistency and experience a rebound with slow positive shifts continuing into the Late Quaternary. Despite δ13C records in the Shaozhai section being interrupted due to the complete leaching of carbonate in some well-developed paleosol layers such as S5–S8, they still attain relatively high values after 0.5 Ma. It shows that the abundance of C4 plants resumed a slow increase with the strengthening of the summer monsoon and finally reached a significant proportion in this region again. Although some high values of Δ13C are also present during this period (Fig. 6), we have refrained from employing the same mechanism to explain these anomalies out of caution. The credibility of several high Δ13C values in our record at 900 and 300 ka is somewhat suspect due to the scarcity of data points. In addition, the lack of a δ18OCarb record hinders further evaluation of the more obvious Δ13C anomaly in the Lantian section at 400 ka (An et al., Reference An, Huang, Liu, Guo, Clemens and Li2005). Therefore, we provisionally conclude that those high Δ13C values cannot robustly suggest the seasonal thriving of C4 plants, despite our comprehensive examination of all pedogenic carbonate samples and another evident C4 plant expansion event in the CLP around 0.5 Ma (Sun et al., Reference Sun, Lü, Zhang, Wang and Liu2012; Zhang et al., Reference Zhang, Liu and Qiang2013a; Zhou et al., Reference Zhou, Shen, Sun, Bird, Ma, Taylor, Liu, Peterse, Yi and Zheng2014).
CONCLUSION
This study investigated the discrepancy between δ13C records of pedogenic carbonate and organic matter occluded within carbonate nodules in the loess–paleosol sequence. It shows that a positive anomaly in δ13CCarb records mainly occurred in the C3/C4 mixed ecosystems and coincided with the expansion of C4 plants. Contamination of pedogenic carbonates by external carbon sources was precluded through morphological, mineralogical, and geochemical evidence. Consequently, we attribute this discrepancy to the fact that δ13CCarb mainly records summer vegetation composition, while δ13CTOC reflects year-round vegetation composition. Our study suggests that the seasonality of C4 plant growth and carbonate precipitation caused positive carbon isotope anomalies in pedogenic carbonates. Therefore, the δ13CCarb could represent the maximum relative abundance of C4 plants that can be achieved during the C4 plant expansion events. The reconstruction based on pedogenic carbonate can be regarded as an upper limit for the relative abundance of C4 plants.
Our results support the idea that the Δ13C has the potential to characterize climate seasonality (Wang et al., Reference Wang, Ambrose and Fouke2004; Sinha et al., Reference Sinha, Tandon, Sanyal, Gibling, Stuben, Berner and Ghazanfari2006; Sanyal et al., Reference Sanyal, Sarkar, Bhattacharya, Kumar, Ghosh and Agrawal2010). However, its application necessitates a thorough evaluation of the following factors. First, the source of carbonates must be carefully examined to ensure their pedogenic origin and to rule out the influence of atmospheric CO2 contamination during carbonate precipitation. Second, although plants of the same photosynthetic type also exhibit δ13C adjustments due to seasonal variations in temperature and precipitation (Stevenson et al., Reference Stevenson, Kelly, McDonald and Busacca2005; Breecker et al., Reference Breecker, Sharp and McFadden2009), Δ13C is better suited for use in the mixed C3/C4 ecosystem instead of a pure C3 or C4 ecosystem. Third, the period of carbonate precipitation is essential for interpreting the documented vegetation and climate changes. Pedogenic carbonate records the vegetation composition in warm seasons in midlatitude monsoon climate zones, while it could represent vegetation composition for entirely different seasons in low-latitude or polar regions (Zamanian et al., Reference Zamanian, Pustovoytov and Kuzyakov2016; Kelson et al., Reference Kelson, Huntington, Breecker, Burgener, Gallagher, Hoke and Petersen2020).
Supplementary Material
The supplementary material for this article can be found at https://doi.org/10.1017/qua.2023.66
Acknowledgments
This research was supported by the National Natural Science Foundation of China (grant nos. 41888101 and 41430531). Special thanks are extended to Nicholas Lancaster for insightful comments and valuable suggestions regarding the article.
