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The role of dissolved organic matter in soil organic carbon stability under water erosion
Ecological Indicators 102 (2019) 724–733
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Original Articles
The role of dissolved organic matter in soil organic carbon stability under water erosion
T
Xuqin Zhanga,c, Zhongwu Lia,b,c, , Xiaodong Nied, , Mei Huanga,c, Danyang Wanga,c, Haibing Xiaob, Chun Liua,c, Hao Penga,c, Jieyu Jianga,c, Guangming Zenga,c ⁎
⁎
a
College of Environmental Science and Engineering, Hunan University, Changsha 410082, PR China State Key Laboratory of Soil Erosion and Dryland Farming on the Loess Plateau, Institute of Soil and Water Conservation CAS and MWR, Yangling, Shaanxi Province 712100, PR China c Key Laboratory of Environmental Biology and Pollution Control (Hunan University), Ministry of Education, Changsha 410082, PR China d Guangdong Key Laboratory of Integrated Agro-environmental Pollution Control and Management, Guangdong Institute of Eco-environmental Science & Technology, Guangzhou 510650, PR China b
ARTICLE INFO
ABSTRACT
Keywords: Soil erosion Deposition Dissolved organic matter Soil organic carbon Stability
Dissolved organic matter (DOM) can respond rapidly to external interference and is of great importance to soil carbon cycling. However, little information is available about the characteristics of DOM in response to water erosion and its role in soil organic carbon (SOC) stability. To contribute this area, a typical watershed in the Southern Red Soil Hilly Region was studied and soil samples were collected from two contrasting sites (eroding versus depositional sites). The DOM concentration and spectral characteristics, SOC stability and their relationships were determined. Results showed that DOM concentration in the deposited topsoil (0–5 cm, 0.69 g kg−1) was significantly higher than that in the eroded soil (0.27 g kg−1). Dissolved organic matter rich in aromatic, hydrophobic and high molecular weight moieties were concentrated in the eroded topsoil rather than in the deposited topsoil. Three fluorescent components in DOM were finally identified: tryptophan-like (C1), humic-like (C2) and protein-like (C3) components. The humic-like C2 in the eroded topsoil was significantly higher than that of the deposited topsoil (P < 0.05). The largest cumulative CO2-C (23.29 mg CO2-C kg−1 soil d−1) was found in the deposited soil of 0–5 cm and it was approximately three times as high as that of the eroded soil. Regression analyses showed that the cumulative CO2 was significantly and positively related to the DOM concentration in two sites, and remarkably negatively correlated with the abundance of aromatic, hydrophobic, high molecular weight and humic-like substances in DOM. Our results indicated that the DOM was strongly affected by water erosion and can be an effective indicator for predicting the SOC stability feedback to erosion.
1. Introduction The soil organic carbon (SOC) pool, estimated at 1550 Pg, acts as an essential component of the global C cycle (Lal, 2003; Batjes, 2014). With the increasing concentration of carbon dioxide (CO2) in the atmosphere, SOC sequestration has gradually been considered a feasible solution to mitigate the greenhouse effect (Lal, 2004). Gaseous emissions from the soil are aggravated by soil degradation, and soil erosion has been considered the most widespread form of the soil degradation (Lal, 2003). Soil erosion influences the SOC dynamics first by a selective process referring to the preferential removal of the light and fine fractions from upslopes to downslopes (Bajracharya et al., 2000), and second by altering the SOC mineralization rate, which was related to the destruction and redistribution of SOC within the watershed (Polyakov and Lal, ⁎
2004). There has been considerable discussions about the effects of these two processes on SOC dynamics. Some studies suggested that the preferential removal of labile organic matter (LOM) led to SOC enrichment in depositional sites, and the deep burial of such C-enriched sediments can improve soil aggregation and thus protect the OC from degradation (Wang et al., 2013). Simultaneously, with the exposure of subsurface soil to air at eroding sites, more photosynthetic products will be sorbed and fixed in the soil, due to the increased availability of noncarbon saturated reactive mineral surface (Berhe et al., 2007; Oost et al., 2007). However, there are also studies believed that the breakdown of aggregates and the exposure of formerly physically protected SOC at eroding sites can increase the mineralization of SOC (Wei et al., 2016). The enrichment of such sediments rich in LOC at depositional sites could lead to a high rate of mineralization (Bajracharya et al.,
Corresponding authors at: College of Environmental Science and Engineering, Hunan University, Changsha 410082, PR China (Z.W. Li) E-mail addresses: [email protected] (Z. Li), [email protected] (X. Nie).
https://doi.org/10.1016/j.ecolind.2019.03.038 Received 14 October 2018; Received in revised form 19 March 2019; Accepted 20 March 2019 Available online 28 March 2019 1470-160X/ © 2019 Elsevier Ltd. All rights reserved.
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Fig. 1. Study location and the cross-sectional diagram representing depth profile and general slope.
2000). Given the debates about the effects of soil erosion on SOC dynamics, there is a need to better understand the response of SOC stability to water erosion. The stability of SOC is often described the resistance against microbial degradation, mineralization incubation is thus generally considered a relatively more representative method to determine the stability of SOC. Information about the stability of SOC is estimated by the amount and rate of CO2-C released during laboratory incubations (Wang et al., 2014b). Several studies have suggested that the mineralization of SOC relies heavily on the size of the labile organic matter pool (Marschner and Kalbitz, 2003; Xiao et al., 2017b), while dissolved organic matter (DOM) is considered to be the most migratory and active fraction of SOM (Zsolnay, 2003; Bolan et al., 2011). Although the DOM accounts for only a small fraction of the total soil organic matter (Chantigny, 2003), it plays a very important role in various biogeochemical processes (biodegradation, sorption, etc). The DOM is subject to sorption by soils (or minerals) and functions as a banding agent for the formation of stable aggregates (Denef et al., 2002; Bolan et al., 2011). The sorption of DOM onto the soil minerals results in the formation of soil organo-mineral complexes, thereby protecting the soil organic carbon from degradation (Bolan et al., 2011). In contrast, the breakdown of soil aggregates or desorption of DOM from the soil matrix could promote the degradation of organic substances (Kuzyakov and Cheng, 2001). The DOM serving as a readily available carbon source for microorganisms (Qualls, 2005; Bolan et al., 2011), it is of great importance to the carbon cycle (i.e., CO2 and CH4). The soil erosion process involves four stages: detachment, breakdown, transport and the deposition of sediments (Lal, 2003). All these stages could not only lead to spatial differentiation of the dissolved organic matter (DOM) over sites and depths but also alter the properties of DOM. Because the quality of DOM seriously affects the biogeochemical processes involved in DOM cycle (Aiken et al., 2011; Jiang et al., 2017), changes in DOM induced by water erosion might be emphasized as an essential factor responsible for the variation of SOC stability at sites and depths. To obtain information on the sources, composition and even reactivity of the DOM in soils or the aquatic environment, optical tools represented by Ultra violet-visible (UV–visible) absorption and fluorescence spectroscopy have been widely applied (Gao et al., 2017; Liu et al., 2019). For example, Weishaar et al. (2003) examined the chemical composition and reactivity of DOM by specific UV absorbance (SUVA) measurements suggested that SUVA could be a useful parameter for characterizing the dissolved aromatic carbon content. Fang et al. (2014) investigated the effects of nitrogen dose and form on the composition and stability of the soil DOM by employing excitation emission matrix (EEM) fluorescence spectroscopy. Liu et al. (2019) identified and quantified the fluorescent components of the soil DOM derived from different types using EEM-parallel factor analysis (PARAFAC). Furthermore, some specific spectroscopic indicators, such as fluorescence index (FI), biological index (BIX) and humification index
(HIX), were utilized to quantify the contributions of various sources to DOM (Derrien et al., 2017; Jiang et al., 2017). There have been attempts to determine the role of the DOM in the terrestrial ecological process and to investigate the biogeochemical dynamics of DOM in soils and sediments. However, knowledge about the patterns in which DOM responds to soil erosion and their effect on SOC stability is rather limited. Thus, objectives of our study were to: (1) investigate the spectral characteristics of the DOM feedback to soil erosion; (2) explore the differences in the SOC stability between eroding and depositional sites; and (3) estimate the relationships between DOM properties and SOC stability. 2. Materials and methods 2.1. Study site Soil samples were collected from the Soil and Water Conservation Monitoring Station of Shaoyang City, Hunan Province, China (111°31′E, 27°18′N) (Fig. 1). The local climate is subtropical monsoon with an annual mean precipitation of 1218.5–1473.5 mm. The main soil types are Ultisols based on the USDA Soil Taxonomy (Soil Survey Staff, 2014), and it is developed from Quaternary parent material and has a sandy and clay textures. The soil is poor in nutrients and has been subjected to severe water erosion (average soil erosion modulus of 2650 t km−2 yr−1), due to frequent heavy rains in spring and summer. Topographically, the slopes in this region generally range from 10° to 15°. Masson pine was the main vegetation prior to the 1960s. With the expansion of agricultural production scale, large areas of pine forest were destroyed and converted to arable land. Since then, soil erosion in this watershed has intensified, and soil fertility has dropped off sharply. By 1993, some slope farmlands were planted with Masson pine, while some were abandoned for vegetation restoration through natural succession to reduce soil degradation. 2.2. Sampling and analysis The upslope and downslope of the watershed covered with nothing but few grasses were selected as eroding and depositional sites, respectively (Fig. 1). Three transects were constructed at each site. In each transect, six soil core samples were collected from the depth: 0–5, 5–10, 10–20, 20–30, 30–40, 40–60, 60–80, 80–100, 100–120 and 120–150 cm by a 5-cm diameter corner and then mixed together. Thus, a total of 60 soil samples (3 transects, 2 sites and 10 depths) were collected. Each sample was divided into three subsamples by visual inspection: one was air-dried at room temperature around 25 °C for the analysis of soil physical and chemical properties; one was sieved through a 2 mm mesh sieve and stored subsequently at 4 °C for the determination of DOM (Nie et al., 2018); and the third was passed through a 2 mm mesh sieve and stored at −70 °C for the incubation of carbon mineralization (Xiao et al., 2017a). 725
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Prior to the soil analysis, the plant residues and gravel in the soil samples were removed. Air-dried soil samples were ground and passed through a 0.25-mm mesh sieve for the analysis of SOC. No carbonates were found in the soil samples as tested repeatedly by adding 1 M HCl. The concentration of SOC was finally determined with a CN analyzer (Vario Max CN, Elementer, Hanau, Germany). Because the distribution of SOC has a certain continuity in the vertical direction, the soil layers with the maximum difference in C content can be selected to study the spatial variation characteristics of C along depth. According to the distribution of SOC over ten depths (S1), samples from the most representative soil layers, 0–5, 5–10, 20–30 and 120–150 cm, were finally selected for further analyses, and the soils at depths shallower than 20 cm were defined as topsoil while the soils deeper than 20 cm were recorded as the subsoil. The air-dried soil samples from the selected layers were crushed to pass through a 2-mm mesh sieve for the determination of soil texture and pH. Soil texture was measured using the pipette method (Gee and Bauder, 1986). Soil pH was measured by a digital pH meter (Hanna Instruments Inc., USA) at a soil and water ratio of 1: 2.5. Soil bulk density (BD) was determined by the cutting ring method. The basic physical and chemical properties of study sites are presented in Table 1.
Fig. 2. The concentrations of dissolved organic carbon (DOC) at different depths and sites (eroding site and depositional site). The error bars represent the standard errors of the means. Different letters (a, b and c) indicate significant differences among various depths at P < 0.05 and different numbers (1 and 2) indicate significant differences between eroding site and depositional site for each soil depth at P < 0.05.
2.3. Carbon mineralization incubation
Milli-Q water to ensure that the UV absorbance at 254 nm was less than 0.3, thus minimizing the inner filtering effects in the later fluorescence scanning (Huang et al., 2018). The UV absorption spectra were recorded between 200 and 800 nm using a UV-2550 spectrophotometer (Shimadzu, Kyoto, Japan), with Millipore water in a 1 cm quartz cuvette as a blank. The DOM three-dimensional fluorescence spectra were measured by an F-7000 Fluorescence spectrophotometer (F-7000, Hitachi) with a 1 cm quartz four-pass-dish. The excitation wavelength ranged from 200 to 550 nm at 5-nm increments and the emission wavelength was recorded at 250 to 600 nm at 2-nm increments. The scanning speed was set at 2400 nm min−1. Milli-Q water was used as a blank. The calibration of EEM data was based on the Milli-Q water Raman peak (Ex = 350 nm) (Lawaetz and Stedmon, 2009), and finally, the fluorescence intensity was corrected to Raman Units (R.U).
Incubation experiments were carried out with soil samples at depths of 0–5, 5–10, 20–30 and 120–150 cm and these samples have adjusted its water contents to 60% water holding capacity (Wang et al., 2014a). Three replicate samples, each one 50 g of the soils from each site and depth were placed in 500 mL incubation bottles and then transferred to the incubator (BIC-400, Shanghai, China) with a preset temperature of 25 °C and kept in the dark for 7 days preincubation, aiming to stimulate microbial activity (Wang et al., 2014a). Three bottles without soil were set as blank controls. The incubation bottles were equipped with two hose-tube valves placed in the lids and sealed with rubber O-rings. Vaseline was applied to the edge of lids to prevent gas leakage (Shimamura et al., 2003). After processing, bottles were incubated in the dark for 82 days (25 °C). During the incubation period on days 2, 5, 12, 19, 26, 33, 40, 47, 54, 61, 68, 75 and 82, the headspace of the incubation bottles was sampled by a sampling units (SwageLok, Solon, USA) from which a 0.2 mL gas sample was taken with a 1 mL syringe for the determination of CO2. After each gas sampling, the incubation bottles were flushed with ambient air and the concentrations of CO2 were analyzed by gas chromatography (GC-2010, Shimadzu, Japan). The C mineralization rate was finally determined by dividing the cumulative CO2-C by 82 days.
2.5. The PARAFAC analyses The PARAFAC modeling was performed based on a total of 72 EEMs of the soil samples (4 depths, 3 transects, 2 sites and 3 replications) using the DOMFluor toolbox in 2015(b) (MathWorks, Natick, MA) (Stedmon and Bro, 2008). Two to seven component models with nonnegativity constraints were used in the PARAFAC analysis. Residual analysis, split half analysis, random initialization and visual inspection were applied to identify the number of fluorescence components (Huang et al., 2019). During the analysis, one sample was removed as an outlier. The relative abundance of each component was reflected by the value of the maximum fluorescence intensity (Fmax) per unit of DOC (Fmax/DOC), while the percentage of each PARAFAC component to the total fluorescence was calculated as dividing the Fmax of each component by the sum of the components’ Fmax (Wu et al., 2012).
2.4. DOM extraction and spectral determination The soil samples for the incubations were used to study the characteristics of DOM and this fraction was extracted with 0.5 M K2SO4 at a 1:5 w/v solution ratio. After shaking and centrifugation, the supernatant was collected through a 0.45 μm filter. The DOM concentration was expressed as dissolved organic carbon (DOC) and was detected by a Shimadzu TOCVCPH (Shimadzu Corp., Kyoto, Japan) (Nie et al., 2018). Subsequently, the DOC concentration of all samples was diluted to 10 mg L−1 or less with Table 1 Soil physicochemical prosperities in eroding site and depositional site. Sites
Soil depth (cm)
SOC (g kg−1)
Sand (%)
Eroding site
0–5 5–10 20–30 120–150
9.88 6.35 2.14 1.61
16.23 12.72 17.51 16.89
Depositional site
0–5 5–10 20–30 120–150
16.96 ± 1.56 9.85 ± 0.74 2.56 ± 0.74 1.34 ± 0.25
± ± ± ±
1.60 0.77 0.77 0.28
± ± ± ±
Silt (%)
Clay (%)
BD (g cm−3)
pH
1.25 5.22 3.88 7.17
60.91 62.23 59.64 61.83
± ± ± ±
1.21 6.24 1.56 6.33
22.86 25.05 22.85 21.29
± ± ± ±
0.91 1.52 2.52 1.08
3.72 3.66 4.04 4.17
± ± ± ±
0.07 0.02 0.16 0.07
1.36 1.43 1.51 1.70
± ± ± ±
0.01 0.03 0.01 0.09
6.49 ± 1.71 4.53 ± 1.82 16.57 ± 3.38 19.16 ± 3.60
52.26 51.82 58.31 60.32
± ± ± ±
1.25 0.91 0.97 3.59
41.25 43.65 28.37 20.52
± ± ± ±
1.59 1.17 5.44 0.99
3.91 4.02 4.13 4.29
± ± ± ±
0.04 0.01 0.07 0.12
1.02 1.30 1.51 1.62
± ± ± ±
0.03 0.10 0.11 0.10
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Fig. 3. The specific UV-absorbance (a) at 254 nm (SUVA254); (b) the ratio of UV absorbance at 250 nm and 365 nm (E2/E3) at each depth from eroding site and depositional site. The error bars represent the standard errors of the means.
2.6. Spectroscopic indices
The DOC in the deposited soils at depths of 0–5 and 5–10 cm were approximately twice as high as those in the eroded soils (P < 0.05), while no significant differences were found in the soils deeper than 20 cm (P > 0.05). A previous study has shown that the labile C pool in the soils is susceptible to release, due to changes in the soil microenvironment (Ahn et al., 2009). Soil erosion induced the detachment and breakdown of the soil aggregates at surface layers, and then the soil particles were susceptible to some causative agents (e.g., runoff and gravity) and finally deposited in low-lying areas. These processes not only affected the sizes of the DOM pools, but also altered their spatial redistribution in soils and over landscapes thus leading to the differences between the eroding and depositional sites. The formation of densely packed layers could mainly be responsible for the decrease in DOM with depth. As a result of the high clay content in red soil, a dense layer formed and restricted the leaching and migration of DOM from the surface to the deep layers (Nie et al., 2018). Therefore, the DOM contents in deep soils were at a relatively low and stable level, duo to the shortage of new carbon input. Absorption spectroscopic indices characterizing the DOM properties are displayed in Fig. 3. The SUVA254 values at both the eroding and depositional sites showed a similar downtrend with increasing depth, while a reverse trend was found for the E2/E3 values. These results indicated that the hydrophobic and aromatic components in the DOM assembled on the soil surface and the molecular weight of the DOM at both sites decreased with the soil depth. The SUVA254 in the soil at depths of 0–5 and 5–10 cm at the eroding site were significantly higher than those at the depositional site (P < 0.05). Conversely, E2/E3 values in the deposited soil at depths of 0–5 and 5–10 cm were much higher rather than in the eroded soils. Results implied that the aromaticity, hydrophobicity and molecular weight of the DOM in the eroded topsoil were significantly higher than that in the deposited topsoil. The selective migration of soil materials along hillsides induced by surface runoff in the soil erosion process might be responsible for these phenomenon (Neff and Asner, 2001). The DOM rich in aromatic, hydrophobic and high molecular weight moieties were preferentially sorbed on the mineral surfaces of soil matrix, while the DOM high in hydrophilic fractions seemed to be mobile and active (Guggenberger and Kaiser, 2003; Kaiser et al., 2004). During the erosion processes, the hydrophilic components characterized by low aromaticity and low molecular weight would thus be preferentially removed by the surface runoff and finally enriched in the depositional site, whereas the hydrophobic components lagged behind and concentrated in the eroding site. No significant differences in SUVA254 and E2/E3 were found between the soils deeper than 20 cm at the two sites, which confirmed that soil erosion severely affected the
Absorption coefficients, a(λ), were obtained by the following equation,
a( ) = 2.303A( )/L
(1)
where A(λ) is the absorbance at wavelength λ and L is the optical pathlength (L = 0.01 m). The ratio of the absorbance at 250 nm to the absorbance at 365 nm was defined as E2/E3, which was strongly and negatively correlated with the average molecular weight of DOM (Peuravuori and Pihlaja, 1997). The specific UV absorbance at 254 nm (SUVA 254) was calculated as a 254/DOC (L mg C−1 m−1), which was positively correlated with the aromaticity or hydrophobicity of DOM (Weishaar et al., 2003). The fluorescence index (FI) was defined as the ratio of emission (Em) intensities at 470 nm to those at 520 nm, with an excitation (Ex) wavelength of 370 nm (Cory et al., 2010), which has been widely used to differentiate between terrestrial and microbial DOM (∼1.9 for microbial sources and ∼1.4 for terrestrial sources) (McKnight et al., 2001). The biological index (BIX) was calculated as the ratio of emission intensity at 380 nm to that at 430 nm (excitation wavelength was kept at 310 nm), which corresponded to a predominantly autochthonous origin for BIX > 1 and a low autochthonous component for 0.6 < BIX < 0.7 (Huguet et al., 2009). The humification index (HIX) was calculated as the ratio of the integral area under emission wavelength 435–480 nm to that under 300–345 nm at excitation 254 nm, which tends to increase with a stronger humic character (Huguet et al., 2009). 2.7. Statistical analyses Statistical analyses were performed with SPSS version 20.0 (SPSS Inc., Chicago, IL, USA). Differences in DOC concentration, cumulative CO2-C and spectroscopic indices among sites and depths were assessed by one-way ANOVA with a least significant difference (LSD) test at 95% confidence level. The results with P < 0.05 were reported as significant, while results with P > 0.05 were considered as non significant. A linear (or nonlinear) regression analysis was performed in the Origin 9.1 (Origin Lab, Northampton, MA) to assess the relationships between the cumulative CO2C and DOM properties (i.e. DOC content, SUVA254, E2:E3 and Fmax). 3. Results and discussion 3.1. The quantity and quality characteristics of DOM responses to soil erosion The DOC concentration at the eroding and depositional sites decreased by 74% and 94% from 0–5 to 120–150 cm, respectively (Fig. 2). 727
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Fig. 4. Three fluorescence components in DOM derived from different sites and depths identified by EEM-PARAFAC analysis.
soils at the surface rather than the deep layers (Nie et al., 2018). The EEM maps were generated from the soil samples collected from the eroding and depositional sites (Fig. 4). Four main fluorescence peaks were observed, and three fluorescent components were identified, namely, C1, C2 and C3 (Table 2). The components C1 showed the fluorescence characteristics of the tryptophan-like material (Chai et al., 2012). The component C2 was related to humic-like material (Coble, 1996) and the component of C3 was bound up with the protein-like material (Wang et al., 2017). The identified components and their quantitative characteristics of
Table 2 Spectral characteristics of three fluorescent components identified by EEMPARAFAC. Components
Ex/Em max (nm)
Substance
References
C1 C2 C3
260/352 230/424 220/294(364)
tryptophan-like Humic-like Protein-like
(Chai et al., 2012) (Coble, 1996) (Wang et al., 2017)
Ex/Em max(nm): the excitation wavelength and the emission wavelength where the maximum values located.
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Fig. 5. Distribution of three PARAFAC-derived components in samples from different depth at eroding and depositional sites. The maximum fluorescence intensity (Fmax) per unit of DOC (Fmax/DOC) and the percentage of each PARAFAC component to the total fluorescence for the eroding site (a, c) and depositional site (b, d).
sites showed no significant difference (P > 0.05). A possible reason for the changes in the DOM components over sites and depths was that the soils and minerals were selective for DOM sorption (Kaiser et al., 2004). The compounds rich in carboxyl and aromatic C would be preferentially retained by soils during an erosion event, while the compounds high in sugars and amino sugars were susceptible to leaching and migration. The FI values of the eroded and depositional soil were both larger than 1.9 (Fig. 6) and showed little variation with increasing depth. No significant differences in the FI values were found between the eroded soil and deposited soils, which indicated that the DOM derived from both sites may have a similar microbial source. The BIX values of both sites were increased with depth, and this value in the soil under 20 cm was larger than 1. The BIX values in eroding site were not much different from that in depositional site. These results suggested that the subsoil DOM had a predominantly autochthonous origin, while the topsoil DOM seemed to be mainly allochthonous sources. The HIX values from the two sites showed a similar downtrend with increasing depth, implying that the topsoil OC hold a higher degree of humification than the subsoil. The HIX values in the eroded soil shallower than 20 cm were a bit larger than those in the deposited soil, which indicated that the eroded soil contained more condensed polyaromatic structures and less oxygen-containing functional groups than the deposited soil (Liu et al., 2019).
Fig. 6. Fluorescence index of DOM derived from different depths and sites.
samples from the eroding and depositional sites are presented in Fig. 5. The relative abundance of the tryptophan-like component C1 increased with depth at both sites. The humic-like component C2 decreased with depth, and the protein-like component C3 showed no obvious changes with increasing depth. For the soil DOM at depths shallower than 20 cm, tryptophan-like component C1 and humic-like component C2 accounted for almost three quarters of the total fluorescent components at the eroding site, while the humic-like C2 and the protein-like component C3 accounted for 83% of the total fluorescent components at the depositional site. For the soil DOM at depths deeper than 20 cm, the tryptophan-like components C1 acted as the primary fractions in both the eroding and depositional sites, and the contents of each component from the two
3.2. The stability of SOC at the eroding and depositional sites The C mineralization incubation is generally considered to be an effective method of assessing the resistance of C to decomposition (Gregorich et al., 2015). The cumulative CO2-Cs emitted over an 82 day incubation period from the eroding and the depositional sites are presented in Fig. 7. The cumulative CO2 emitted from the deposited soil at a depth of 0–5 cm (1909.60 ± 84.55 mg CO2-C kg−1 soil) was approximately 3 times that of from the eroded soil (624.80 ± 63.45 mg 729
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CO2-C kg−1 soil), and statistical analysis showed a significant difference between these two values. The C mineralization rates for the eroded and deposited soils were high during the first 20 days and then decreased (Fig. 7), which was similar to the results reported by Wei et al. (2016). In that report, the initial amount of labile OC was considered as an essential part of this phenomenon. The maximum mineralization rate (23.29 ± 1.03 mg CO2-C kg−1 soil d−1) was observed in the deposited soil at a depth of 0–5 cm, followed by the eroded soil at a depth of 0–5 cm (7.62 ± 0.27 mg CO2-C kg−1 soil d−1), and the minimum mineralization rate was found in the soil of the eroding site at a depth of 120–150 cm (3.00 ± 0.12 mg CO2-C kg−1 soil d−1). Statistical analysis illustrated that the mineralization rates in the topsoil (0–5 and 5–10 cm) at both sites were significantly (P < 0.05) higher than those in the subsoil (20–30 and 120–150 cm). No significant difference was observed between the subsoil of the two sites (P > 0.05). Two main reasons may be responsible for this observation: (i) the subsoil suffered small disturbances by erosion, and thus the OC in subsoil was maintained in a relatively stable state (Bernal et al., 2016); and (ii) the subsoil was short of microbial substrates (e.g., labile organic compounds, molecular oxygen, or available nutrients) (Yakov, 2010). The cumulative CO2, as well as the mineralization rate, of deposited topsoil was significantly greater than that from the eroded topsoil. It can be reasonably assumed that the difference in the bioavailable carbon content within the eroded and deposited soils contributed to difference in the CO2 emission flux from the two sites. During erosional events, the selective migration process has brought a large amount of labile organic matter from the eroding to the depositional sites and thus led to the increase of bioavailable carbon in depositional site. The variations in the cumulative CO2 and mineralization rate between the two sites implied that the OC of eroded topsoil was more stable than the deposited topsoil.
Fig. 7. Cumulative C mineralized over 82 days incubation period for soil samples from each depth at eroding site and depositional site. The vertical bars are standard deviations. E1: eroding site at 0–5 cm; E2: eroding site at 5–10 cm; E3: eroding site 20–30 cm; E4: eroding site 120–150 cm; D1: depositional site at 0–5 cm; D2: depositional site at 5–10 cm; D3: depositional site at 20–30 cm; D4: depositional site at 120–150 cm.
3.3. Relationships between DOM and the stability of SOC under erosion and deposition The DOM serves as the most bioavailable fraction to soil microorganisms (Marschner and Kalbitz, 2003). Therefore, the size of the DOM pool may play a decisive role in the mineralization of SOC. In our study, an enrichment of light and small particles was found in the topsoil at the depositional site, and the highest cumulative CO2-C was observed at the same site. Our experimental data showed that the variation of DOC content followed a trend similar to that of the cumulative CO2-C and mineralization rate. This finding was in agreement with several studies (Wang et al., 2013; Xiao et al., 2017b). Wang et al. (2013) observed that low mineralization rates were often accompanied by small contents of the water-extractable OC, based on the data collected in two contrasting sites in the Belgian Loess Belt. Xiao et al.
Fig. 8. Relationships between the concentrations of DOC and the cumulative CO2-C at eroding site and depositional site.
Fig. 9. Relationships between SUVA254 and the cumulative CO2-C in the topsoil of eroding site (a) and depositional site (b).
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Fig. 10. Relationships between E2:E3 and the cumulative CO2-C in the topsoil of eroding site (a) and depositional site (b).
Fig. 11. Relationships between the abundance of components C2 and the cumulative CO2-C in the topsoil of eroding site (a) and depositional site (b).
(2017b) suggested that labile organic matter was the primary ratelimiting factor for microbial respiration in the hilly area of the Loess Plateau. The above mentioned observations suggested that the DOC content may actually have a positive relationship with the stability of SOC. To test this hypothesis, a regression analysis between DOC contents and the cumulative CO2-C was performed (Fig. 8). The DOC contents at the eroding and depositional sites were significantly positively correlated with the amount of cumulative CO2-C (R2 = 0.92 and 0.99, respectively). This finding not only provided a further evidence for the statement by Ahn et al. (2009) that DOM can be used as the most efficient soil variable to quantify the potential for soil C mineralization, but also implied that the DOM contents exerted an important influence on the stability of SOC. Several studies on the chemical structure of DOM have showed that the enrichment of aromatic and hydrophilic components is often accompanied by a low degradation rate, due to its strong resistance (Kalbitz et al., 2003). Vitousek et al. (2003) suggested that the components with small molecular weights, such as carbohydrates and amino acids are more easily used by microorganisms than hydrophobic fractions. In our study, the DOM at depths of 0–5 and 5–10 cm in the eroded soil had a higher aromatic (hydrophobic) component content and a larger molecular weight than that in the deposited soil, which was contrary to the trend for the cumulative CO2-C at the two sites. This finding implied that the enrichment of aromatic (hydrophobic or high molecular weight) components may be a rate-limiting factor for topsoil C mineralization. The high aromaticity (hydrophobicity and molecular weight) of DOM may have a positive effect on the stability of SOC. To test this hypothesis, regression analyses between the cumulative CO2-C and absorption spectroscopic indices were performed. The results showed
that the values of SUVA254 in the topsoil were markedly negatively related to the cumulative CO2-C (Fig. 9), while the E2/E3 values of the soil at depths of 0–5 and 5–10 cm were significantly positively correlated to the cumulative CO2-C (Fig. 10), which implied that the aromaticity, as well as the hydrophobicity and molecular weight of the DOM was essential to the SOC biostability at depth shallower than 20 cm. However, it was notable that such correlations between the cumulative CO2-C and absorption spectroscopic indices were not found in the subsoil at either the eroding or depositional sites, which indicated that the stability mechanisms of the topsoil organic carbon may differ from those of deep subsoil organic carbon (Bernal et al., 2016). Referring to the chemical components of DOM, tryptophan-like and protein-like components were often characterized by simple structures (Marhuenda-Egea et al., 2007), while humic-like components were characterized by large molecular, conjugated aromatic moieties and complex structures (Guo et al., 2012). In our study, we found that the tryptophan-like component C1 played a dominant role in the soil below 20 cm, while the humic-like component C2 had a large proportion in the surface 0–20 cm layer. The relative abundance of component C2 in the eroded topsoil was significantly higher than that in the deposited topsoil (P < 0.05), while the abundance of relatively simple components (tryptophan-like C1 and protein-like C3 combined) in the top soils of two sites did not differ significantly (P > 0.05). Our results supported the finding by Kaiser et al. (2004), who proposed that the topsoil-derived DOM had more aromatic substances whereas the subsoilderived DOM contained more simple-structured substances. Our results showed that the DOM components in the topsoil of the eroding site were more complex than those of the depositional site, and the humiclike component C2 might be the main reason for these differences. Past studies on DOM have shown that the more complex the DOM structure is, the more difficult it is to microbial degradation (Wang and 731
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Zhang, 2010). Although the components of the DOM in the soil below 20 cm seemed more available to microorganisms than that in the topsoil 0–20 cm. The cumulative CO2-C and mineralization rate of the subsoil OC was much lower than that of the topsoil. These finding implied that the components of DOM were not an essential limiting factor for the biostability of subsoil OC. When it came to the importance of the DOM components to topsoil OC, results were quite different. Regression analyses between the topsoil DOM components and the cumulative CO2-C have showed that the relative abundance of the component C2 was significantly negatively correlated with the cumulative CO2-C (Fig. 11). On the basis of these findings we may conclude that the humic-like component C2 of the DOM is essential to C mineralization. The fraction of the humic-like component in the DOM exerts considerable influence on the biostability of topsoil OC under erosion and deposition.
induced terrestrial carbon sink. Bioscience 57, 337–346. Bernal, B., Mckinley, D.C., Hungate, B.A., White, P.M., Mozdzer, T.J., Megonigal, J.P., 2016. Limits to soil carbon stability; deep, ancient soil carbon decomposition stimulated by new labile organic inputs. Soil Biol. Biochem. 98, 85–94. Bolan, N.S., Adriano, D.C., Kunhikrishnan, A., James, T., McDowell, R., Senesi, N., 2011. Dissolved organic matter: biogeochemistry, dynamics, and environmental significance in soils. Adv. Agron. Elsevier 1–75. Chai, X., Liu, G., Zhao, X., Hao, Y., Zhao, Y., 2012. Complexion between mercury and humic substances from different landfill stabilization processes and its implication for the environment. J. Hazard. Mater. 209–210, 59. Chantigny, M.H., 2003. Dissolved and water-extractable organic matter in soils: a review on the influence of land use and management practices. Geoderma 113, 357–380. Coble, P.G., 1996. Characterization of marine and terrestrial DOM in seawater using excitation-emission matrix spectroscopy. Mar. Chem. 51, 325–346. Cory, R.M., Miller, M.P., McKnight, D.M., Guerard, J.J., Miller, P.L., 2010. Effect of instrument-specific response on the analysis of fulvic acid fluorescence spectra. Limnol. Oceanogr.-Meth. 8, 67–78. Denef, K., Six, J., Merckx, R., Paustian, K., 2002. Short-term effects of biological and physical forces on aggregate formation in soils with different clay mineralogy. Plant Soil 246, 185–200. Derrien, M., Yang, L., Hur, J., 2017. Lipid biomarkers and spectroscopic indices for identifying organic matter sources in aquatic environments: a review. Water Res. 112, 58–71. Fang, H., Cheng, S., Yu, G., Xu, M., Wang, Y., Li, L., Dang, X., Wang, L., Li, Y., 2014. Experimental nitrogen deposition alters the quantity and quality of soil dissolved organic carbon in an alpine meadow on the Qinghai-Tibetan Plateau. Appl. Soil Ecol. 81, 1–11. 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Huang, M., Li, Z., Huang, B., Luo, N., Zhang, Q., Zhai, X., Zeng, G., 2018. Investigating binding characteristics of cadmium and copper to DOM derived from compost and rice straw using EEM-PARAFAC combined with two-dimensional FTIR correlation analyses. J. Hazard. Mater. 344, 539–548. Huang, M., Li, Z., Luo, N., Yang, R., Wen, J., Huang, B., Zeng, G., 2019. Application potential of biochar in environment: insight from degradation of biochar-derived DOM and complexation of DOM with heavy metals. Sci. Total Environ. 646, 220–228. Huguet, A., Vacher, L., Relexans, S., Saubusse, S., Froidefond, J.M., Parlanti, E., 2009. Properties of fluorescent dissolved organic matter in the Gironde Estuary. Org. Geochem. 40, 706–719. Jiang, T., Kaal, J., Liang, J., Zhang, Y., Wei, S., Wang, D., Green, N.W., 2017. Composition of dissolved organic matter (DOM) from periodically submerged soils in the Three Gorges Reservoir areas as determined by elemental and optical analysis, infrared spectroscopy, pyrolysis-GC–MS and thermally assisted hydrolysis and methylation. Sci. Total Environ. 603–604, 461–471. Kaiser, K., Guggenberger, G., Haumaier, L., 2004. Changes in dissolved lignin-derived phenols, neutral sugars, uronic acids, and amino sugars with depth in forested Haplic Arenosols and Rendzic Leptosols. Biogeochemistry 70, 135–151. Kalbitz, K., Schmerwitz, J., Schwesig, D., Matzner, E., 2003. Biodegradation of soil-derived dissolved organic matter as related to its properties. Geoderma 113, 273–291. Kuzyakov, Y., Cheng, W., 2001. Photosynthesis controls of rhizosphere respiration and organic matter decomposition. Soil Biol. Biochem. 33, 1915–1925. Lal, R., 2003. Soil erosion and the global carbon budget. Environ. Int. 29, 437. Lal, R., 2004. Soil carbon sequestration to mitigate climate change. Geoderma 123, 1–22. Lawaetz, A.J., Stedmon, C.A., 2009. fluorescence intensity calibration using the Raman scatter peak of water. Appl. Spectrosc. 63, 936–940. Liu, C., Li, Z., Berhe, A.A., Xiao, H., Liu, L., Wang, D., Peng, H., Zeng, G., 2019. Characterizing dissolved organic matter in eroded sediments from a loess hilly catchment using fluorescence EEM-PARAFAC and UV–Visible absorption: insights from source identification and carbon cycling. Geoderma 334, 37–48. Marhuenda-Egea, F.C., Martínez-Sabater, E., Jordá, J., Moral, R., Bustamante, M.A., Paredes, C., Pérez-Murcia, M.D., 2007. Dissolved organic matter fractions formed during composting of winery and distillery residues: evaluation of the process by fluorescence excitation–emission matrix. Chemosphere 68, 301–309. Marschner, B., Kalbitz, K., 2003. Controls of bioavailability and biodegradability of dissolved organic matter in soils. Geoderma 113, 211–235. McKnight, D.M., Boyer, E.W., Westerhoff, P.K., Doran, P.T., Kulbe, T., Andersen, D.T., 2001. Spectrofluorometric characterization of dissolved organic matter for indication of precursor organic material and aromaticity. Limnol. Oceanogr. 46, 38–48. Neff, J.C., Asner, G.P., 2001. Dissolved organic carbon in terrestrial ecosystems: synthesis and a model. Ecosystems 4, 29–48. Nie, X., Li, Z., Huang, J., Liu, L., Xiao, H., Liu, C., Zeng, G., 2018. Thermal stability of organic carbon in soil aggregates as affected by soil erosion and deposition. Soil Till. Res. 175, 82–90. Oost, K.V., Quine, T.A., Govers, G., Gryze, S.D., Six, J., Harden, J.W., Ritchie, J.C.,
4. Conclusions The results of our study showed that the differences in the DOM contents, spectral characteristics and SOC stability between eroding and depositional sites were limited to the depth of 0–10 cm. The variation of the DOM contents in the two sites indicated that water erosion resulted in a preferential removal of the light and fine fractions from the upslope to the downslope sites. The DOM extracted from the eroded soils had more aromatics, hydrophobic and larger molecular size fractions than that of deposited soils, implying that soil erosion caused the enrichment of the labile matters in the deposited topsoil and the accumulation of relatively resistant substances in the eroded topsoil. Compared with the sediments at the depositional area, the soil at the eroded site showed a greater abundance of the humic-like component C2 and had a higher degree of humification, which suggested that the eroded soil contained more condensed polyaromatic structures and less oxygen-containing functional groups than in the deposited soil. Incubation results suggested that the eroded SOC was much more stable than the deposited SOC. The DOC contents and DOM spectral characteristics at the two sites were strongly related to the cumulative CO2-C, which revealed that the quantity and quality responses of the DOM to soil erosion had an important impact on the stability of SOC. Combined results suggested that the DOM was severely affected by water erosion and it was of great importance in assessing the stability of SOC. Acknowledgments Financial support for this study was provided by the National Natural Science Foundation of China (41271294, 51521006), the ‘Hundred-talent Project’ of the Chinese Academy of Sciences and the GDAS' Special Project of Science and Technology Development (2018GDASCX-1002). Appendix A. Supplementary data Supplementary data to this article can be found online at https:// doi.org/10.1016/j.ecolind.2019.03.038. References Ahn, M.Y., Zimmerman, A.R., Comerford, N.B., Sickman, J.O., Grunwald, S., 2009. Carbon mineralization and labile organic carbon pools in the sandy soils of a North Florida Watershed. Ecosystems 12, 672–685. Aiken, G.R., Hsu-Kim, H., Ryan, J.N., 2011. Influence of dissolved organic matter on the environmental fate of metals, nanoparticles, and colloids. Environ. Sci. Technol. 45, 3196–3201. Bajracharya, R.M., Lal, R., Kimble, J.M., 2000. Erosion effects on carbon dioxide concentration and carbon flux from an ohio alfisol. Soil Sci. Soc. Am. J. 64, 694–700. Batjes, N.H., 2014. Total carbon and nitrogen in the soils of the world. Eur. J. Soil Sci. 65, 10–21. Berhe, A.A., Harte, J., Harden, J.W., Torn, M.S., 2007. The significance of the erosion-
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Soil Sci. 64, 219–228. Wang, X., Zhang, F., Kung, H.T., Ghulam, A., Trumbo, A.L., Yang, J., Ren, Y., Jing, Y., 2017. Evaluation and estimation of surface water quality in an arid region based on EEM-PARAFAC and 3D fluorescence spectral index: a case study of the Ebinur Lake Watershed, China. Catena 155, 62–74. Wang, Z.P., Zhang, T., 2010. Characterization of soluble microbial products (SMP) under stressful conditions. Water Res. 44, 5499–5509. Wei, S., Zhang, X., McLaughlin, N.B., Yang, X., Liang, A., Jia, S., Chen, X., 2016. Effect of breakdown and dispersion of soil aggregates by erosion on soil CO2 emission. Geoderma 264, 238–243. Weishaar, J.L., Aiken, G.R., Bergamaschi, B.A., Fram, M.S., Fujii, R., Mopper, K., 2003. Evaluation of specific ultraviolet absorbance as an indicator of the chemical composition and reactivity of dissolved organic carbon. Environ. Sci. Technol. 37, 4702–4708. Wu, J., Zhang, H., Shao, L.M., He, P.J., 2012. Fluorescent characteristics and metal binding properties of individual molecular weight fractions in municipal solid waste leachate. Environ. Pollut. 162, 63–71. Xiao, H., Li, Z., Chang, X., Huang, J., Nie, X., Liu, C., Liu, L., Wang, D., Dong, Y., Jiang, J., 2017a. Soil erosion-related dynamics of soil bacterial communities and microbial respiration. Appl. Soil Ecol. 119, 205–213. Xiao, H., Li, Z., Dong, Y., Chang, X., Deng, L., Huang, J., Nie, X., Liu, C., Liu, L., Wang, D., 2017b. Changes in microbial communities and respiration following the revegetation of eroded soil. Agric. Ecosyst. Environ. 246, 30–37. Yakov, K., 2010. Priming effects: interactions between living and dead organic matter. Soil Biol. Biochem. 42, 1363–1371. Zsolnay, Á., 2003. Dissolved organic matter: artefacts, definitions, and functions. Geoderma 113, 187–209.
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Original Articles
The role of dissolved organic matter in soil organic carbon stability under water erosion
T
Xuqin Zhanga,c, Zhongwu Lia,b,c, , Xiaodong Nied, , Mei Huanga,c, Danyang Wanga,c, Haibing Xiaob, Chun Liua,c, Hao Penga,c, Jieyu Jianga,c, Guangming Zenga,c ⁎
⁎
a
College of Environmental Science and Engineering, Hunan University, Changsha 410082, PR China State Key Laboratory of Soil Erosion and Dryland Farming on the Loess Plateau, Institute of Soil and Water Conservation CAS and MWR, Yangling, Shaanxi Province 712100, PR China c Key Laboratory of Environmental Biology and Pollution Control (Hunan University), Ministry of Education, Changsha 410082, PR China d Guangdong Key Laboratory of Integrated Agro-environmental Pollution Control and Management, Guangdong Institute of Eco-environmental Science & Technology, Guangzhou 510650, PR China b
ARTICLE INFO
ABSTRACT
Keywords: Soil erosion Deposition Dissolved organic matter Soil organic carbon Stability
Dissolved organic matter (DOM) can respond rapidly to external interference and is of great importance to soil carbon cycling. However, little information is available about the characteristics of DOM in response to water erosion and its role in soil organic carbon (SOC) stability. To contribute this area, a typical watershed in the Southern Red Soil Hilly Region was studied and soil samples were collected from two contrasting sites (eroding versus depositional sites). The DOM concentration and spectral characteristics, SOC stability and their relationships were determined. Results showed that DOM concentration in the deposited topsoil (0–5 cm, 0.69 g kg−1) was significantly higher than that in the eroded soil (0.27 g kg−1). Dissolved organic matter rich in aromatic, hydrophobic and high molecular weight moieties were concentrated in the eroded topsoil rather than in the deposited topsoil. Three fluorescent components in DOM were finally identified: tryptophan-like (C1), humic-like (C2) and protein-like (C3) components. The humic-like C2 in the eroded topsoil was significantly higher than that of the deposited topsoil (P < 0.05). The largest cumulative CO2-C (23.29 mg CO2-C kg−1 soil d−1) was found in the deposited soil of 0–5 cm and it was approximately three times as high as that of the eroded soil. Regression analyses showed that the cumulative CO2 was significantly and positively related to the DOM concentration in two sites, and remarkably negatively correlated with the abundance of aromatic, hydrophobic, high molecular weight and humic-like substances in DOM. Our results indicated that the DOM was strongly affected by water erosion and can be an effective indicator for predicting the SOC stability feedback to erosion.
1. Introduction The soil organic carbon (SOC) pool, estimated at 1550 Pg, acts as an essential component of the global C cycle (Lal, 2003; Batjes, 2014). With the increasing concentration of carbon dioxide (CO2) in the atmosphere, SOC sequestration has gradually been considered a feasible solution to mitigate the greenhouse effect (Lal, 2004). Gaseous emissions from the soil are aggravated by soil degradation, and soil erosion has been considered the most widespread form of the soil degradation (Lal, 2003). Soil erosion influences the SOC dynamics first by a selective process referring to the preferential removal of the light and fine fractions from upslopes to downslopes (Bajracharya et al., 2000), and second by altering the SOC mineralization rate, which was related to the destruction and redistribution of SOC within the watershed (Polyakov and Lal, ⁎
2004). There has been considerable discussions about the effects of these two processes on SOC dynamics. Some studies suggested that the preferential removal of labile organic matter (LOM) led to SOC enrichment in depositional sites, and the deep burial of such C-enriched sediments can improve soil aggregation and thus protect the OC from degradation (Wang et al., 2013). Simultaneously, with the exposure of subsurface soil to air at eroding sites, more photosynthetic products will be sorbed and fixed in the soil, due to the increased availability of noncarbon saturated reactive mineral surface (Berhe et al., 2007; Oost et al., 2007). However, there are also studies believed that the breakdown of aggregates and the exposure of formerly physically protected SOC at eroding sites can increase the mineralization of SOC (Wei et al., 2016). The enrichment of such sediments rich in LOC at depositional sites could lead to a high rate of mineralization (Bajracharya et al.,
Corresponding authors at: College of Environmental Science and Engineering, Hunan University, Changsha 410082, PR China (Z.W. Li) E-mail addresses: [email protected] (Z. Li), [email protected] (X. Nie).
https://doi.org/10.1016/j.ecolind.2019.03.038 Received 14 October 2018; Received in revised form 19 March 2019; Accepted 20 March 2019 Available online 28 March 2019 1470-160X/ © 2019 Elsevier Ltd. All rights reserved.
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Fig. 1. Study location and the cross-sectional diagram representing depth profile and general slope.
2000). Given the debates about the effects of soil erosion on SOC dynamics, there is a need to better understand the response of SOC stability to water erosion. The stability of SOC is often described the resistance against microbial degradation, mineralization incubation is thus generally considered a relatively more representative method to determine the stability of SOC. Information about the stability of SOC is estimated by the amount and rate of CO2-C released during laboratory incubations (Wang et al., 2014b). Several studies have suggested that the mineralization of SOC relies heavily on the size of the labile organic matter pool (Marschner and Kalbitz, 2003; Xiao et al., 2017b), while dissolved organic matter (DOM) is considered to be the most migratory and active fraction of SOM (Zsolnay, 2003; Bolan et al., 2011). Although the DOM accounts for only a small fraction of the total soil organic matter (Chantigny, 2003), it plays a very important role in various biogeochemical processes (biodegradation, sorption, etc). The DOM is subject to sorption by soils (or minerals) and functions as a banding agent for the formation of stable aggregates (Denef et al., 2002; Bolan et al., 2011). The sorption of DOM onto the soil minerals results in the formation of soil organo-mineral complexes, thereby protecting the soil organic carbon from degradation (Bolan et al., 2011). In contrast, the breakdown of soil aggregates or desorption of DOM from the soil matrix could promote the degradation of organic substances (Kuzyakov and Cheng, 2001). The DOM serving as a readily available carbon source for microorganisms (Qualls, 2005; Bolan et al., 2011), it is of great importance to the carbon cycle (i.e., CO2 and CH4). The soil erosion process involves four stages: detachment, breakdown, transport and the deposition of sediments (Lal, 2003). All these stages could not only lead to spatial differentiation of the dissolved organic matter (DOM) over sites and depths but also alter the properties of DOM. Because the quality of DOM seriously affects the biogeochemical processes involved in DOM cycle (Aiken et al., 2011; Jiang et al., 2017), changes in DOM induced by water erosion might be emphasized as an essential factor responsible for the variation of SOC stability at sites and depths. To obtain information on the sources, composition and even reactivity of the DOM in soils or the aquatic environment, optical tools represented by Ultra violet-visible (UV–visible) absorption and fluorescence spectroscopy have been widely applied (Gao et al., 2017; Liu et al., 2019). For example, Weishaar et al. (2003) examined the chemical composition and reactivity of DOM by specific UV absorbance (SUVA) measurements suggested that SUVA could be a useful parameter for characterizing the dissolved aromatic carbon content. Fang et al. (2014) investigated the effects of nitrogen dose and form on the composition and stability of the soil DOM by employing excitation emission matrix (EEM) fluorescence spectroscopy. Liu et al. (2019) identified and quantified the fluorescent components of the soil DOM derived from different types using EEM-parallel factor analysis (PARAFAC). Furthermore, some specific spectroscopic indicators, such as fluorescence index (FI), biological index (BIX) and humification index
(HIX), were utilized to quantify the contributions of various sources to DOM (Derrien et al., 2017; Jiang et al., 2017). There have been attempts to determine the role of the DOM in the terrestrial ecological process and to investigate the biogeochemical dynamics of DOM in soils and sediments. However, knowledge about the patterns in which DOM responds to soil erosion and their effect on SOC stability is rather limited. Thus, objectives of our study were to: (1) investigate the spectral characteristics of the DOM feedback to soil erosion; (2) explore the differences in the SOC stability between eroding and depositional sites; and (3) estimate the relationships between DOM properties and SOC stability. 2. Materials and methods 2.1. Study site Soil samples were collected from the Soil and Water Conservation Monitoring Station of Shaoyang City, Hunan Province, China (111°31′E, 27°18′N) (Fig. 1). The local climate is subtropical monsoon with an annual mean precipitation of 1218.5–1473.5 mm. The main soil types are Ultisols based on the USDA Soil Taxonomy (Soil Survey Staff, 2014), and it is developed from Quaternary parent material and has a sandy and clay textures. The soil is poor in nutrients and has been subjected to severe water erosion (average soil erosion modulus of 2650 t km−2 yr−1), due to frequent heavy rains in spring and summer. Topographically, the slopes in this region generally range from 10° to 15°. Masson pine was the main vegetation prior to the 1960s. With the expansion of agricultural production scale, large areas of pine forest were destroyed and converted to arable land. Since then, soil erosion in this watershed has intensified, and soil fertility has dropped off sharply. By 1993, some slope farmlands were planted with Masson pine, while some were abandoned for vegetation restoration through natural succession to reduce soil degradation. 2.2. Sampling and analysis The upslope and downslope of the watershed covered with nothing but few grasses were selected as eroding and depositional sites, respectively (Fig. 1). Three transects were constructed at each site. In each transect, six soil core samples were collected from the depth: 0–5, 5–10, 10–20, 20–30, 30–40, 40–60, 60–80, 80–100, 100–120 and 120–150 cm by a 5-cm diameter corner and then mixed together. Thus, a total of 60 soil samples (3 transects, 2 sites and 10 depths) were collected. Each sample was divided into three subsamples by visual inspection: one was air-dried at room temperature around 25 °C for the analysis of soil physical and chemical properties; one was sieved through a 2 mm mesh sieve and stored subsequently at 4 °C for the determination of DOM (Nie et al., 2018); and the third was passed through a 2 mm mesh sieve and stored at −70 °C for the incubation of carbon mineralization (Xiao et al., 2017a). 725
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Prior to the soil analysis, the plant residues and gravel in the soil samples were removed. Air-dried soil samples were ground and passed through a 0.25-mm mesh sieve for the analysis of SOC. No carbonates were found in the soil samples as tested repeatedly by adding 1 M HCl. The concentration of SOC was finally determined with a CN analyzer (Vario Max CN, Elementer, Hanau, Germany). Because the distribution of SOC has a certain continuity in the vertical direction, the soil layers with the maximum difference in C content can be selected to study the spatial variation characteristics of C along depth. According to the distribution of SOC over ten depths (S1), samples from the most representative soil layers, 0–5, 5–10, 20–30 and 120–150 cm, were finally selected for further analyses, and the soils at depths shallower than 20 cm were defined as topsoil while the soils deeper than 20 cm were recorded as the subsoil. The air-dried soil samples from the selected layers were crushed to pass through a 2-mm mesh sieve for the determination of soil texture and pH. Soil texture was measured using the pipette method (Gee and Bauder, 1986). Soil pH was measured by a digital pH meter (Hanna Instruments Inc., USA) at a soil and water ratio of 1: 2.5. Soil bulk density (BD) was determined by the cutting ring method. The basic physical and chemical properties of study sites are presented in Table 1.
Fig. 2. The concentrations of dissolved organic carbon (DOC) at different depths and sites (eroding site and depositional site). The error bars represent the standard errors of the means. Different letters (a, b and c) indicate significant differences among various depths at P < 0.05 and different numbers (1 and 2) indicate significant differences between eroding site and depositional site for each soil depth at P < 0.05.
2.3. Carbon mineralization incubation
Milli-Q water to ensure that the UV absorbance at 254 nm was less than 0.3, thus minimizing the inner filtering effects in the later fluorescence scanning (Huang et al., 2018). The UV absorption spectra were recorded between 200 and 800 nm using a UV-2550 spectrophotometer (Shimadzu, Kyoto, Japan), with Millipore water in a 1 cm quartz cuvette as a blank. The DOM three-dimensional fluorescence spectra were measured by an F-7000 Fluorescence spectrophotometer (F-7000, Hitachi) with a 1 cm quartz four-pass-dish. The excitation wavelength ranged from 200 to 550 nm at 5-nm increments and the emission wavelength was recorded at 250 to 600 nm at 2-nm increments. The scanning speed was set at 2400 nm min−1. Milli-Q water was used as a blank. The calibration of EEM data was based on the Milli-Q water Raman peak (Ex = 350 nm) (Lawaetz and Stedmon, 2009), and finally, the fluorescence intensity was corrected to Raman Units (R.U).
Incubation experiments were carried out with soil samples at depths of 0–5, 5–10, 20–30 and 120–150 cm and these samples have adjusted its water contents to 60% water holding capacity (Wang et al., 2014a). Three replicate samples, each one 50 g of the soils from each site and depth were placed in 500 mL incubation bottles and then transferred to the incubator (BIC-400, Shanghai, China) with a preset temperature of 25 °C and kept in the dark for 7 days preincubation, aiming to stimulate microbial activity (Wang et al., 2014a). Three bottles without soil were set as blank controls. The incubation bottles were equipped with two hose-tube valves placed in the lids and sealed with rubber O-rings. Vaseline was applied to the edge of lids to prevent gas leakage (Shimamura et al., 2003). After processing, bottles were incubated in the dark for 82 days (25 °C). During the incubation period on days 2, 5, 12, 19, 26, 33, 40, 47, 54, 61, 68, 75 and 82, the headspace of the incubation bottles was sampled by a sampling units (SwageLok, Solon, USA) from which a 0.2 mL gas sample was taken with a 1 mL syringe for the determination of CO2. After each gas sampling, the incubation bottles were flushed with ambient air and the concentrations of CO2 were analyzed by gas chromatography (GC-2010, Shimadzu, Japan). The C mineralization rate was finally determined by dividing the cumulative CO2-C by 82 days.
2.5. The PARAFAC analyses The PARAFAC modeling was performed based on a total of 72 EEMs of the soil samples (4 depths, 3 transects, 2 sites and 3 replications) using the DOMFluor toolbox in 2015(b) (MathWorks, Natick, MA) (Stedmon and Bro, 2008). Two to seven component models with nonnegativity constraints were used in the PARAFAC analysis. Residual analysis, split half analysis, random initialization and visual inspection were applied to identify the number of fluorescence components (Huang et al., 2019). During the analysis, one sample was removed as an outlier. The relative abundance of each component was reflected by the value of the maximum fluorescence intensity (Fmax) per unit of DOC (Fmax/DOC), while the percentage of each PARAFAC component to the total fluorescence was calculated as dividing the Fmax of each component by the sum of the components’ Fmax (Wu et al., 2012).
2.4. DOM extraction and spectral determination The soil samples for the incubations were used to study the characteristics of DOM and this fraction was extracted with 0.5 M K2SO4 at a 1:5 w/v solution ratio. After shaking and centrifugation, the supernatant was collected through a 0.45 μm filter. The DOM concentration was expressed as dissolved organic carbon (DOC) and was detected by a Shimadzu TOCVCPH (Shimadzu Corp., Kyoto, Japan) (Nie et al., 2018). Subsequently, the DOC concentration of all samples was diluted to 10 mg L−1 or less with Table 1 Soil physicochemical prosperities in eroding site and depositional site. Sites
Soil depth (cm)
SOC (g kg−1)
Sand (%)
Eroding site
0–5 5–10 20–30 120–150
9.88 6.35 2.14 1.61
16.23 12.72 17.51 16.89
Depositional site
0–5 5–10 20–30 120–150
16.96 ± 1.56 9.85 ± 0.74 2.56 ± 0.74 1.34 ± 0.25
± ± ± ±
1.60 0.77 0.77 0.28
± ± ± ±
Silt (%)
Clay (%)
BD (g cm−3)
pH
1.25 5.22 3.88 7.17
60.91 62.23 59.64 61.83
± ± ± ±
1.21 6.24 1.56 6.33
22.86 25.05 22.85 21.29
± ± ± ±
0.91 1.52 2.52 1.08
3.72 3.66 4.04 4.17
± ± ± ±
0.07 0.02 0.16 0.07
1.36 1.43 1.51 1.70
± ± ± ±
0.01 0.03 0.01 0.09
6.49 ± 1.71 4.53 ± 1.82 16.57 ± 3.38 19.16 ± 3.60
52.26 51.82 58.31 60.32
± ± ± ±
1.25 0.91 0.97 3.59
41.25 43.65 28.37 20.52
± ± ± ±
1.59 1.17 5.44 0.99
3.91 4.02 4.13 4.29
± ± ± ±
0.04 0.01 0.07 0.12
1.02 1.30 1.51 1.62
± ± ± ±
0.03 0.10 0.11 0.10
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Fig. 3. The specific UV-absorbance (a) at 254 nm (SUVA254); (b) the ratio of UV absorbance at 250 nm and 365 nm (E2/E3) at each depth from eroding site and depositional site. The error bars represent the standard errors of the means.
2.6. Spectroscopic indices
The DOC in the deposited soils at depths of 0–5 and 5–10 cm were approximately twice as high as those in the eroded soils (P < 0.05), while no significant differences were found in the soils deeper than 20 cm (P > 0.05). A previous study has shown that the labile C pool in the soils is susceptible to release, due to changes in the soil microenvironment (Ahn et al., 2009). Soil erosion induced the detachment and breakdown of the soil aggregates at surface layers, and then the soil particles were susceptible to some causative agents (e.g., runoff and gravity) and finally deposited in low-lying areas. These processes not only affected the sizes of the DOM pools, but also altered their spatial redistribution in soils and over landscapes thus leading to the differences between the eroding and depositional sites. The formation of densely packed layers could mainly be responsible for the decrease in DOM with depth. As a result of the high clay content in red soil, a dense layer formed and restricted the leaching and migration of DOM from the surface to the deep layers (Nie et al., 2018). Therefore, the DOM contents in deep soils were at a relatively low and stable level, duo to the shortage of new carbon input. Absorption spectroscopic indices characterizing the DOM properties are displayed in Fig. 3. The SUVA254 values at both the eroding and depositional sites showed a similar downtrend with increasing depth, while a reverse trend was found for the E2/E3 values. These results indicated that the hydrophobic and aromatic components in the DOM assembled on the soil surface and the molecular weight of the DOM at both sites decreased with the soil depth. The SUVA254 in the soil at depths of 0–5 and 5–10 cm at the eroding site were significantly higher than those at the depositional site (P < 0.05). Conversely, E2/E3 values in the deposited soil at depths of 0–5 and 5–10 cm were much higher rather than in the eroded soils. Results implied that the aromaticity, hydrophobicity and molecular weight of the DOM in the eroded topsoil were significantly higher than that in the deposited topsoil. The selective migration of soil materials along hillsides induced by surface runoff in the soil erosion process might be responsible for these phenomenon (Neff and Asner, 2001). The DOM rich in aromatic, hydrophobic and high molecular weight moieties were preferentially sorbed on the mineral surfaces of soil matrix, while the DOM high in hydrophilic fractions seemed to be mobile and active (Guggenberger and Kaiser, 2003; Kaiser et al., 2004). During the erosion processes, the hydrophilic components characterized by low aromaticity and low molecular weight would thus be preferentially removed by the surface runoff and finally enriched in the depositional site, whereas the hydrophobic components lagged behind and concentrated in the eroding site. No significant differences in SUVA254 and E2/E3 were found between the soils deeper than 20 cm at the two sites, which confirmed that soil erosion severely affected the
Absorption coefficients, a(λ), were obtained by the following equation,
a( ) = 2.303A( )/L
(1)
where A(λ) is the absorbance at wavelength λ and L is the optical pathlength (L = 0.01 m). The ratio of the absorbance at 250 nm to the absorbance at 365 nm was defined as E2/E3, which was strongly and negatively correlated with the average molecular weight of DOM (Peuravuori and Pihlaja, 1997). The specific UV absorbance at 254 nm (SUVA 254) was calculated as a 254/DOC (L mg C−1 m−1), which was positively correlated with the aromaticity or hydrophobicity of DOM (Weishaar et al., 2003). The fluorescence index (FI) was defined as the ratio of emission (Em) intensities at 470 nm to those at 520 nm, with an excitation (Ex) wavelength of 370 nm (Cory et al., 2010), which has been widely used to differentiate between terrestrial and microbial DOM (∼1.9 for microbial sources and ∼1.4 for terrestrial sources) (McKnight et al., 2001). The biological index (BIX) was calculated as the ratio of emission intensity at 380 nm to that at 430 nm (excitation wavelength was kept at 310 nm), which corresponded to a predominantly autochthonous origin for BIX > 1 and a low autochthonous component for 0.6 < BIX < 0.7 (Huguet et al., 2009). The humification index (HIX) was calculated as the ratio of the integral area under emission wavelength 435–480 nm to that under 300–345 nm at excitation 254 nm, which tends to increase with a stronger humic character (Huguet et al., 2009). 2.7. Statistical analyses Statistical analyses were performed with SPSS version 20.0 (SPSS Inc., Chicago, IL, USA). Differences in DOC concentration, cumulative CO2-C and spectroscopic indices among sites and depths were assessed by one-way ANOVA with a least significant difference (LSD) test at 95% confidence level. The results with P < 0.05 were reported as significant, while results with P > 0.05 were considered as non significant. A linear (or nonlinear) regression analysis was performed in the Origin 9.1 (Origin Lab, Northampton, MA) to assess the relationships between the cumulative CO2C and DOM properties (i.e. DOC content, SUVA254, E2:E3 and Fmax). 3. Results and discussion 3.1. The quantity and quality characteristics of DOM responses to soil erosion The DOC concentration at the eroding and depositional sites decreased by 74% and 94% from 0–5 to 120–150 cm, respectively (Fig. 2). 727
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Fig. 4. Three fluorescence components in DOM derived from different sites and depths identified by EEM-PARAFAC analysis.
soils at the surface rather than the deep layers (Nie et al., 2018). The EEM maps were generated from the soil samples collected from the eroding and depositional sites (Fig. 4). Four main fluorescence peaks were observed, and three fluorescent components were identified, namely, C1, C2 and C3 (Table 2). The components C1 showed the fluorescence characteristics of the tryptophan-like material (Chai et al., 2012). The component C2 was related to humic-like material (Coble, 1996) and the component of C3 was bound up with the protein-like material (Wang et al., 2017). The identified components and their quantitative characteristics of
Table 2 Spectral characteristics of three fluorescent components identified by EEMPARAFAC. Components
Ex/Em max (nm)
Substance
References
C1 C2 C3
260/352 230/424 220/294(364)
tryptophan-like Humic-like Protein-like
(Chai et al., 2012) (Coble, 1996) (Wang et al., 2017)
Ex/Em max(nm): the excitation wavelength and the emission wavelength where the maximum values located.
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Fig. 5. Distribution of three PARAFAC-derived components in samples from different depth at eroding and depositional sites. The maximum fluorescence intensity (Fmax) per unit of DOC (Fmax/DOC) and the percentage of each PARAFAC component to the total fluorescence for the eroding site (a, c) and depositional site (b, d).
sites showed no significant difference (P > 0.05). A possible reason for the changes in the DOM components over sites and depths was that the soils and minerals were selective for DOM sorption (Kaiser et al., 2004). The compounds rich in carboxyl and aromatic C would be preferentially retained by soils during an erosion event, while the compounds high in sugars and amino sugars were susceptible to leaching and migration. The FI values of the eroded and depositional soil were both larger than 1.9 (Fig. 6) and showed little variation with increasing depth. No significant differences in the FI values were found between the eroded soil and deposited soils, which indicated that the DOM derived from both sites may have a similar microbial source. The BIX values of both sites were increased with depth, and this value in the soil under 20 cm was larger than 1. The BIX values in eroding site were not much different from that in depositional site. These results suggested that the subsoil DOM had a predominantly autochthonous origin, while the topsoil DOM seemed to be mainly allochthonous sources. The HIX values from the two sites showed a similar downtrend with increasing depth, implying that the topsoil OC hold a higher degree of humification than the subsoil. The HIX values in the eroded soil shallower than 20 cm were a bit larger than those in the deposited soil, which indicated that the eroded soil contained more condensed polyaromatic structures and less oxygen-containing functional groups than the deposited soil (Liu et al., 2019).
Fig. 6. Fluorescence index of DOM derived from different depths and sites.
samples from the eroding and depositional sites are presented in Fig. 5. The relative abundance of the tryptophan-like component C1 increased with depth at both sites. The humic-like component C2 decreased with depth, and the protein-like component C3 showed no obvious changes with increasing depth. For the soil DOM at depths shallower than 20 cm, tryptophan-like component C1 and humic-like component C2 accounted for almost three quarters of the total fluorescent components at the eroding site, while the humic-like C2 and the protein-like component C3 accounted for 83% of the total fluorescent components at the depositional site. For the soil DOM at depths deeper than 20 cm, the tryptophan-like components C1 acted as the primary fractions in both the eroding and depositional sites, and the contents of each component from the two
3.2. The stability of SOC at the eroding and depositional sites The C mineralization incubation is generally considered to be an effective method of assessing the resistance of C to decomposition (Gregorich et al., 2015). The cumulative CO2-Cs emitted over an 82 day incubation period from the eroding and the depositional sites are presented in Fig. 7. The cumulative CO2 emitted from the deposited soil at a depth of 0–5 cm (1909.60 ± 84.55 mg CO2-C kg−1 soil) was approximately 3 times that of from the eroded soil (624.80 ± 63.45 mg 729
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CO2-C kg−1 soil), and statistical analysis showed a significant difference between these two values. The C mineralization rates for the eroded and deposited soils were high during the first 20 days and then decreased (Fig. 7), which was similar to the results reported by Wei et al. (2016). In that report, the initial amount of labile OC was considered as an essential part of this phenomenon. The maximum mineralization rate (23.29 ± 1.03 mg CO2-C kg−1 soil d−1) was observed in the deposited soil at a depth of 0–5 cm, followed by the eroded soil at a depth of 0–5 cm (7.62 ± 0.27 mg CO2-C kg−1 soil d−1), and the minimum mineralization rate was found in the soil of the eroding site at a depth of 120–150 cm (3.00 ± 0.12 mg CO2-C kg−1 soil d−1). Statistical analysis illustrated that the mineralization rates in the topsoil (0–5 and 5–10 cm) at both sites were significantly (P < 0.05) higher than those in the subsoil (20–30 and 120–150 cm). No significant difference was observed between the subsoil of the two sites (P > 0.05). Two main reasons may be responsible for this observation: (i) the subsoil suffered small disturbances by erosion, and thus the OC in subsoil was maintained in a relatively stable state (Bernal et al., 2016); and (ii) the subsoil was short of microbial substrates (e.g., labile organic compounds, molecular oxygen, or available nutrients) (Yakov, 2010). The cumulative CO2, as well as the mineralization rate, of deposited topsoil was significantly greater than that from the eroded topsoil. It can be reasonably assumed that the difference in the bioavailable carbon content within the eroded and deposited soils contributed to difference in the CO2 emission flux from the two sites. During erosional events, the selective migration process has brought a large amount of labile organic matter from the eroding to the depositional sites and thus led to the increase of bioavailable carbon in depositional site. The variations in the cumulative CO2 and mineralization rate between the two sites implied that the OC of eroded topsoil was more stable than the deposited topsoil.
Fig. 7. Cumulative C mineralized over 82 days incubation period for soil samples from each depth at eroding site and depositional site. The vertical bars are standard deviations. E1: eroding site at 0–5 cm; E2: eroding site at 5–10 cm; E3: eroding site 20–30 cm; E4: eroding site 120–150 cm; D1: depositional site at 0–5 cm; D2: depositional site at 5–10 cm; D3: depositional site at 20–30 cm; D4: depositional site at 120–150 cm.
3.3. Relationships between DOM and the stability of SOC under erosion and deposition The DOM serves as the most bioavailable fraction to soil microorganisms (Marschner and Kalbitz, 2003). Therefore, the size of the DOM pool may play a decisive role in the mineralization of SOC. In our study, an enrichment of light and small particles was found in the topsoil at the depositional site, and the highest cumulative CO2-C was observed at the same site. Our experimental data showed that the variation of DOC content followed a trend similar to that of the cumulative CO2-C and mineralization rate. This finding was in agreement with several studies (Wang et al., 2013; Xiao et al., 2017b). Wang et al. (2013) observed that low mineralization rates were often accompanied by small contents of the water-extractable OC, based on the data collected in two contrasting sites in the Belgian Loess Belt. Xiao et al.
Fig. 8. Relationships between the concentrations of DOC and the cumulative CO2-C at eroding site and depositional site.
Fig. 9. Relationships between SUVA254 and the cumulative CO2-C in the topsoil of eroding site (a) and depositional site (b).
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Fig. 10. Relationships between E2:E3 and the cumulative CO2-C in the topsoil of eroding site (a) and depositional site (b).
Fig. 11. Relationships between the abundance of components C2 and the cumulative CO2-C in the topsoil of eroding site (a) and depositional site (b).
(2017b) suggested that labile organic matter was the primary ratelimiting factor for microbial respiration in the hilly area of the Loess Plateau. The above mentioned observations suggested that the DOC content may actually have a positive relationship with the stability of SOC. To test this hypothesis, a regression analysis between DOC contents and the cumulative CO2-C was performed (Fig. 8). The DOC contents at the eroding and depositional sites were significantly positively correlated with the amount of cumulative CO2-C (R2 = 0.92 and 0.99, respectively). This finding not only provided a further evidence for the statement by Ahn et al. (2009) that DOM can be used as the most efficient soil variable to quantify the potential for soil C mineralization, but also implied that the DOM contents exerted an important influence on the stability of SOC. Several studies on the chemical structure of DOM have showed that the enrichment of aromatic and hydrophilic components is often accompanied by a low degradation rate, due to its strong resistance (Kalbitz et al., 2003). Vitousek et al. (2003) suggested that the components with small molecular weights, such as carbohydrates and amino acids are more easily used by microorganisms than hydrophobic fractions. In our study, the DOM at depths of 0–5 and 5–10 cm in the eroded soil had a higher aromatic (hydrophobic) component content and a larger molecular weight than that in the deposited soil, which was contrary to the trend for the cumulative CO2-C at the two sites. This finding implied that the enrichment of aromatic (hydrophobic or high molecular weight) components may be a rate-limiting factor for topsoil C mineralization. The high aromaticity (hydrophobicity and molecular weight) of DOM may have a positive effect on the stability of SOC. To test this hypothesis, regression analyses between the cumulative CO2-C and absorption spectroscopic indices were performed. The results showed
that the values of SUVA254 in the topsoil were markedly negatively related to the cumulative CO2-C (Fig. 9), while the E2/E3 values of the soil at depths of 0–5 and 5–10 cm were significantly positively correlated to the cumulative CO2-C (Fig. 10), which implied that the aromaticity, as well as the hydrophobicity and molecular weight of the DOM was essential to the SOC biostability at depth shallower than 20 cm. However, it was notable that such correlations between the cumulative CO2-C and absorption spectroscopic indices were not found in the subsoil at either the eroding or depositional sites, which indicated that the stability mechanisms of the topsoil organic carbon may differ from those of deep subsoil organic carbon (Bernal et al., 2016). Referring to the chemical components of DOM, tryptophan-like and protein-like components were often characterized by simple structures (Marhuenda-Egea et al., 2007), while humic-like components were characterized by large molecular, conjugated aromatic moieties and complex structures (Guo et al., 2012). In our study, we found that the tryptophan-like component C1 played a dominant role in the soil below 20 cm, while the humic-like component C2 had a large proportion in the surface 0–20 cm layer. The relative abundance of component C2 in the eroded topsoil was significantly higher than that in the deposited topsoil (P < 0.05), while the abundance of relatively simple components (tryptophan-like C1 and protein-like C3 combined) in the top soils of two sites did not differ significantly (P > 0.05). Our results supported the finding by Kaiser et al. (2004), who proposed that the topsoil-derived DOM had more aromatic substances whereas the subsoilderived DOM contained more simple-structured substances. Our results showed that the DOM components in the topsoil of the eroding site were more complex than those of the depositional site, and the humiclike component C2 might be the main reason for these differences. Past studies on DOM have shown that the more complex the DOM structure is, the more difficult it is to microbial degradation (Wang and 731
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Zhang, 2010). Although the components of the DOM in the soil below 20 cm seemed more available to microorganisms than that in the topsoil 0–20 cm. The cumulative CO2-C and mineralization rate of the subsoil OC was much lower than that of the topsoil. These finding implied that the components of DOM were not an essential limiting factor for the biostability of subsoil OC. When it came to the importance of the DOM components to topsoil OC, results were quite different. Regression analyses between the topsoil DOM components and the cumulative CO2-C have showed that the relative abundance of the component C2 was significantly negatively correlated with the cumulative CO2-C (Fig. 11). On the basis of these findings we may conclude that the humic-like component C2 of the DOM is essential to C mineralization. The fraction of the humic-like component in the DOM exerts considerable influence on the biostability of topsoil OC under erosion and deposition.
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4. Conclusions The results of our study showed that the differences in the DOM contents, spectral characteristics and SOC stability between eroding and depositional sites were limited to the depth of 0–10 cm. The variation of the DOM contents in the two sites indicated that water erosion resulted in a preferential removal of the light and fine fractions from the upslope to the downslope sites. The DOM extracted from the eroded soils had more aromatics, hydrophobic and larger molecular size fractions than that of deposited soils, implying that soil erosion caused the enrichment of the labile matters in the deposited topsoil and the accumulation of relatively resistant substances in the eroded topsoil. Compared with the sediments at the depositional area, the soil at the eroded site showed a greater abundance of the humic-like component C2 and had a higher degree of humification, which suggested that the eroded soil contained more condensed polyaromatic structures and less oxygen-containing functional groups than in the deposited soil. Incubation results suggested that the eroded SOC was much more stable than the deposited SOC. The DOC contents and DOM spectral characteristics at the two sites were strongly related to the cumulative CO2-C, which revealed that the quantity and quality responses of the DOM to soil erosion had an important impact on the stability of SOC. Combined results suggested that the DOM was severely affected by water erosion and it was of great importance in assessing the stability of SOC. Acknowledgments Financial support for this study was provided by the National Natural Science Foundation of China (41271294, 51521006), the ‘Hundred-talent Project’ of the Chinese Academy of Sciences and the GDAS' Special Project of Science and Technology Development (2018GDASCX-1002). Appendix A. Supplementary data Supplementary data to this article can be found online at https:// doi.org/10.1016/j.ecolind.2019.03.038. References Ahn, M.Y., Zimmerman, A.R., Comerford, N.B., Sickman, J.O., Grunwald, S., 2009. Carbon mineralization and labile organic carbon pools in the sandy soils of a North Florida Watershed. Ecosystems 12, 672–685. Aiken, G.R., Hsu-Kim, H., Ryan, J.N., 2011. Influence of dissolved organic matter on the environmental fate of metals, nanoparticles, and colloids. Environ. Sci. Technol. 45, 3196–3201. Bajracharya, R.M., Lal, R., Kimble, J.M., 2000. Erosion effects on carbon dioxide concentration and carbon flux from an ohio alfisol. Soil Sci. Soc. Am. J. 64, 694–700. Batjes, N.H., 2014. Total carbon and nitrogen in the soils of the world. Eur. J. Soil Sci. 65, 10–21. Berhe, A.A., Harte, J., Harden, J.W., Torn, M.S., 2007. The significance of the erosion-
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