Distribution Patterns of Humus and Mineral Composition in Dark-Brown, Meadow, and Paddy Soils in Northeast China

Abstract

This study aimed to investigate vertical variations in dissolved organic matter (DOM) properties, humus (HS) composition, humic acid (HA) characteristics, and clay mineral dynamics, with a particular focus on the vertical distribution of HS components and mineral composition across Dark-brown, Meadow, and Paddy soil profiles. Results indicated that: (1) DOM in all three soil types was predominantly endogenous, primarily derived from microbial metabolism with minimal contributions from plant residues. (2) Vertical trends in DOM carbon content (C_DOM) were specific to soil type: in Dark-brown soil, C_DOM slightly increased from the Ap to Bt layer, followed by a sharp increase in the C layer; Meadow soil exhibited a significant decrease in C_DOM in the AB layer but remained relatively stable in other layers; Paddy soil showed a consistent decline in C_DOM with increasing depth. (3) HS and its fractions exhibited vertical variability: Paddy soil showed higher HS content in surface layers; carbon contents of water-soluble substances, HA, and humic-extracted acid (C_WSS, C_HA, and C_HE) decreased with depth in Dark-brown and Paddy soils, whereas they remained relatively stable in deeper layers of Meadow soil. (4) HA characteristics, including C/N ratio, functional groups, and aromaticity, were influenced by both depth and soil type: the Ap2 layer of Paddy soil effectively restricted the downward movement of organic matter; Fe³⁺ complexation played a key role in HA stabilization in Dark-brown soil; Meadow soil exhibited transitional HS properties. (5) Clay mineral assemblages were dominated by 2:1 type minerals (illite, smectite, illite–smectite interstratifications), showing distinct vertical weathering patterns: illite content decreased with depth due to hydrolysis, while proton-driven dissolution promoted kaolinite formation in surface layers, particularly in Dark-brown soil 2:1 minerals enhancing organic–mineral complexation in Meadow soil. The findings of this study provided a scientific basis for optimizing soil carbon pool management and offer insights into organic–mineral interactions that can enhance organic matter sequestration in agricultural soils.

1. Introduction

The Songnen Plain is a major grain-crop production region in Northeast China. This zone is mainly a temperate monsoon climate, and grain crops are harvested once a year. In this area, apart from Black soil and Albic soil, the main types of agricultural soil include Dark-brown soil, Meadow soil, and Paddy soil, with a total area of approximately 1.0326 × 10⁸ hm². These soil types constitute vital agricultural resources, serving as a cornerstone for national food security. As the largest carbon (C) reservoir in terrestrial ecosystems, soil organic matter (SOM) plays a key role in mitigating global warming and enhancing soil functionality [1]. Humus (HS) constitutes the primary component of SOM and accounts for the largest proportion of SOM [2]. However, intensive farming practices combined with inadequate organic fertilizer input lead to severe C loss. Over time, this “input-output imbalance” diminishes soil fertility and reduces C storage capacity [3]. Clay minerals are integral to maintaining soil health and productivity. Through physical (e.g., structure and water retention), chemical (e.g., nutrient retention and buffering capacity), and biological (e.g., microbial habitat support) mechanisms, clay minerals sustain the equilibrium of soil ecosystems. Moreover, they play a critical role in forming stable organic–mineral complexes via processes such as surface adsorption and interlayer complexation with SOM [4]. This interaction effectively protects OM from microbial decomposition, thereby facilitating long-term C sequestration in soils [5].

Recent studies have highlighted the pivotal role of humification and clay mineral weathering in shaping SOM dynamics. Tarasova et al. [6] proposed that adding vermiculite could enhance the soil’s physical structure and foster stronger bonds between SOM and mineral particles. Tombácz et al. [7] observed that active surfaces formed through clay mineral weathering facilitated the recombination of HS and minerals. Hidayat et al. [8] revealed that in near-volcanic soils containing an illite–kaolinite combination, the chelation with Fe³⁺–HS enhanced the bioavailability of potassium (K), calcium (Ca), and magnesium (Mg), demonstrating the regulatory influence of specific minerals on nutrient availability. Drewnik et al. [9] showed that carbonate-mediated stabilization played an essential role, with the carbonate–clay–SOM complex driving HS accumulation in Chernozem. Hampl et al. [10] established a mechanical link between interfacial selectivity and iron dynamics, showing that biotic weathering induced tensile stresses through the oxidation of Fe²⁺ to Fe³⁺. This disrupted the mineral lattice, exposing hydroxylated edges that supported SOM binding—a critical step in forming Saprosol soils. Zhang et al. [11] demonstrated that long-term fertilization increased the ratio of the C content of HA to FA in Paddy soil by 130.0%, attributed to the co-precipitation of iron and SOM. However, this process also led to an exponential decline in SOM with depth, driven by anaerobic micro-environments dissolving protective Fe³⁺ bridges. Li et al. [12] studying the Loess Plateau, reported a consistent vertical pattern in dissolved organic matter (DOM), where C content of DOM decreased with depth while tryptophan content increased.

Most existing studies have primarily focused on the interactive effects between surface SOM and minerals, with a notable lack of longitudinal comparative analyses on HS and mineral composition across different soil types within similar agricultural ecosystems. In view of this, the study selected three typical dryland soil profiles from Dark-brown, Meadow, and Paddy soil in Jilin Province as the research subject and employed a comprehensive suite of analytical techniques. These included three-dimensional excitation-emission matrix (3D-EEM) fluorescence spectroscopy for DOM, Fourier Transform Infrared (FTIR) spectroscopy, and elemental composition analysis of HA, HS composition analysis, combined with X-ray diffraction (XRD) of mineral. Using these methods, key indicators such as the composition of HS, the molecular structure of HA, and clay mineral composition were systematically assessed to elucidate the development of targeted sustainable soil management strategies and enhance C sequestration potential.

2. Materials and Methods

2.1. Study Site and Soil Sampling Procedures

The study area was located in Longtan District and Panshi City, Jilin City, China. Based on the specific areas within the electronic fence of the Third National Soil Census system and thorough on-site investigations, three representative soil types were selected as sampling sites: Kaoshan Village in Jiangbei Township (Dark-brown soil), located at E 126.52957713752103, N 43.97054004668869; Hashi Village in Gangyao Town (meadow soil), located at E 126.69213517315029, N 44.199627152853594; and Zhenxing Village in Yantongshan Town, Panshi City (paddy soil), located at E 126.0070519959121, N 44.42577565055015. These sampling points were evenly distributed across the hilly terrain of the residual region of Changbai Mountain, with elevations ranging from 180 to 600 m and slopes varying between 2° and 6°. The area falls within a temperate monsoon climate zone, characterized by an average annual temperature of 4.5 °C and annual precipitation of 630–680 mm. Notably, dark-brown soil and meadow soil have been continuously used for corn cultivation for more than 10 years, while paddy soil has only recently been converted to corn farming within the past two years. The corn planting at the three locations employed a single application of basal fertilizer. The application rates of nitrogen (N), phosphorus pentoxide (P₂O₅), and potassium oxide (K₂O) were 220 kg/ha, 80 kg/ha, and 90 kg/ha, respectively. Detailed geological, geomorphological, and agricultural ecological parameters were presented in Table S1.

A deep pit over 120 cm was dug at the sampling point to expose the soil profile. The smooth and rough surfaces were artificially trimmed, as shown in Figure 1. Soil samples were collected from each designated soil layer using a leather hammer and cloth bags, according to the specific stratigraphic layers of each soil profile. At each sampling depth, three replicate samples were collected and combined to form a representative composite sample. In the field, visible root materials were manually removed. The samples were then air-dried at room temperature, ground, and sieved through a 2 mm nylon sieve to eliminate gravel and coarse organic matter. Subsequently, each sample was further ground using a wooden mortar to achieve a fine particle size of less than 0.25 mm for subsequent laboratory analysis. To prevent cross-contamination, all sampling tools were thoroughly cleaned with 95% ethanol between each sampling event.

2.2. Experimental Method

2.2.1. Three-Dimensional Fluorescence Spectrum (3D-EEM) of Dissolved Organic Matter (DOM)

Soil samples were dissolved in distilled water at a weight-to-volume ratio of 1:30 and agitated at 25 °C for 2 h. Subsequently, the mixture was centrifuged at 12,879× g for 15 min and filtered through a 0.45 μm membrane filter. The dissolved organic matter (DOM) C content (C_DOM) was quantified using a total organic C analyzer (vario TOC select, Elementar, Langenselbold, Germany). Fluorescence excitation–emission matrices (EEMs) of the DOM extracts were analyzed using a fluorescence spectrometer (F-7000, Hitachi, Tokyo, Japan), with excitation (E_x) and emission (E_m) wavelengths scanned in the ranges of 200–550 nm and 250–550 nm, respectively. The slit widths for both E_x and E_m were set to 5 nm, with a scanning speed of 2400 nm/min. To minimize Rayleigh and Raman scattering effects, ultrapure water was used as a blank control and subtracted from the EEM data. Fluorescence spectra were visualized using MATLAB 2020 (MathWorks, Natick, MA, USA). The fluorescence index (FI) is defined as the ratio of fluorescence intensity at an excitation wavelength of 370 nm to emissions at wavelengths of 470 nm to 520 nm. The autochthonous contribution (BIX) is calculated as the ratio of emission intensity at 380 nm to 430 nm when excited at 310 nm [13]. The humification index (HIX), defined as the ratio of the fluorescence peak area at 435–480 nm to that at 300–345 nm under 254 nm excitation.

2.2.2. Extraction of Humus Components and Determination of Their C Contents

The HS composition was analyzed using a modified Kumada method. A 5.0 g sample was placed in a 100 mL polyethylene centrifuge tube, and 30 mL of distilled water was added. The mixture was thoroughly stirred and extracted in a 70 °C constant temperature water bath oscillator for 1 h. After extraction, centrifugation was performed at 3500 r/min for 15 min, and the supernatant was transferred to a 50 mL volumetric flask. Then, 20 mL of distilled water was added to the residual sample in the centrifuge tube, followed by thorough stirring and centrifugation again under the same conditions. The supernatants from both steps were combined to obtain the water-soluble substance (WSS). For extracting humic-extracted acid (HE), the distilled water was replaced with a mixture of 0.1 mol/L NaOH and 0.1 mol/L Na₂P₂O₇, and the residue was extracted twice using the same procedure. The collected solution was designated as HE. The remaining residue was dried at 60 °C and sieved through a 0.15 mm mesh to obtain the humin (Hu) fraction.

A 30 mL aliquot of the HE solution was adjusted to pH 1.0–1.5 with 0.5 mol/L H₂SO₄, maintained in a 70 °C water bath for 1.5 h, and left to stand overnight. The next day, the solution containing the precipitate was filtered into a 50 mL volumetric flask and made up to volume, which constituted the fulvic acid (FA). The precipitate on the filter paper was washed with dilute acid, then dissolved in 0.05 mol/L warm NaOH in another 50 mL volumetric flask and diluted to volume with distilled water to prepare the humic acid (HA) base solution. The C contents of WSS, HE, HA, and Hu (C_WSS, C_HE, C_HA, and C_Hu) were determined via the external heating potassium dichromate oxidation method. The C content of FA (C_FA) was calculated by subtraction (C_HE–C_HA), and the C_HA/C_FA ratio was derived accordingly. All reagents mentioned above were of analytical grade and were procured from Shanghai Sinopharm Group Co., LTD., Shanghai, China. Fourier infrared spectroscopy (FTIR-850, Tianjin Gangdong Science and Technology Development Co., LTD., Tianjin, China) was used for qualitative analysis of HA functional groups, and its elemental composition was determined with a 2400 Series II CHNS/O Elemental Analyzer (PerkinElmer Inc., Waltham, MA, USA).

2.2.3. Clay Mineral Separation and Extraction

To determine the composition, the siphon method and sieving method were employed. A 30 g air-dried soil sample was weighed, and SOM was removed using 30% hydrogen peroxide. Decalcification was achieved with diluted hydrochloric acid. The sample underwent ultrasonic dispersion treatment at 150 W for 3 min. Coarse sand particles larger than 0.002 mm were separated using the sieving method. Clay particles smaller than 0.002 mm and silt particles ranging from 0.02 mm to 0.002 mm were extracted using the siphon method within a specified timeframe, with repeated extractions performed until no suspended clay particles smaller than 0.002 mm remained. The collected material was subjected to precipitation, centrifugation, and drying.

X-ray Diffraction: After the <0.002 mm clay fraction was deironed using the sodium disulfite–sodium citrate–sodium bicarbonate method (commonly referred to as the DCB method), the samples were saturated with 1 mol/L potassium acetate and 1 mol/L magnesium acetate solutions. The resulting clay samples were prepared into potassium-saturated oriented specimens (K-air) and magnesium-saturated oriented specimens (Mg-air). Following natural air drying, the oriented samples were analyzed using an X-ray diffractometer (Beijing Puxi General Instrument Co., Ltd., Beijing, China, XD3, CuKa radiation, Ni filter, 360 kV, 25 mA, with a step size of 0.04°). Subsequently, the potassium-saturated oriented samples were subjected to two heating treatments in a muffle furnace at 300 °C and 550 °C for 2 h each (referred to as K-300 and K-550, respectively). Meanwhile, the magnesium-saturated oriented samples underwent further testing after glycerol saturation treatment (Mg-gly) [14].

2.3. Statistical Analysis

Excel 2003, Origin 9.0, and SPSS 20.0 were utilized for statistical analysis and data visualization, while MDI Jade 9.0 was employed for pattern matching. One-way analysis of variance (ANOVA) in SPSS Statistics 20.0 was used to test for significant differences. Statistical significance was determined at a level of p < 0.05. The data were subjected to LSD post hoc test. The presented data represented the mean values of triplicate determinations, with the standard deviation displayed as the error band on the line graph. XRD diffraction spectra obtained under five distinct processing conditions were superimposed and compared to identify clay minerals. Data accuracy was refined using a combination of five-point smoothing, baseline correction, and Gaussian curve fitting techniques. The semi-quantitative analysis of Vermiculite (Ver), illite (I), montmorillonite-illite mixed layers (I/S), montmorillonite (S), chlorite (Chl), and kaolinite (Kaol) was performed using magnesium/potassium heat treatment diffraction spectra. The calculation formulas were provided as follows [15]:

$$w(\mathrm{Kaol}+\mathrm{Chl}+\mathrm{Ver})={\displaystyle \frac{{\mathrm{I}}_{0.7\mathrm{nm}\left(\mathrm{air}\right)}/1.5}{{\mathrm{I}}_{0.7\mathrm{nm}\left(\mathrm{air}\right)}/1.5{+\mathrm{I}}_{1.0\mathrm{nm}(550 °\mathrm{C})}}}\times 100$$

(1)

$$w(\mathrm{Kaol})={\displaystyle \frac{{\mathrm{h}}_{0.358\mathrm{nm}\left(\mathrm{gly}\right)}}{{\mathrm{h}}_{0.358\mathrm{nm}\left(\mathrm{gly}\right)}{+\mathrm{h}}_{0.353\mathrm{nm}\left(\mathrm{gly}\right)}}}\times w(\mathrm{Kaol}+\mathrm{Chl})$$

(2)

$$w\left(\mathrm{I}\right)={\displaystyle \frac{\left({\mathrm{h}}_{0.7\mathrm{nm}\left(\mathrm{air}\right)}{/\mathrm{h}}_{0.7\mathrm{nm}\left(\mathrm{air}\right)}\right)\times {\mathrm{I}}_{1.0\mathrm{nm}\left(\mathrm{gly}\right)}}{{\mathrm{I}}_{1.0\mathrm{nm}\left(550 °\mathrm{C}\right)}}}\times \left[\begin{array}{c}100 - w(\mathrm{Kaol}+\mathrm{Chl})\end{array}\right]$$

(3)

$$w\left(\mathrm{S}\right)={\displaystyle \frac{{\mathrm{I}}_{1.7\mathrm{nm}\left(\mathrm{gly}\right)}/4}{{\mathrm{I}}_{1.0\mathrm{nm}(550 °\mathrm{C})}}}\times \left[\begin{array}{c}100 - w(\mathrm{Kaol}+\mathrm{Chl})\end{array}\right]$$

(4)

$$w(\mathrm{Chl})={\displaystyle \frac{{\mathrm{I}}_{1.4\mathrm{nm}(550 °\mathrm{C})}}{{\mathrm{I}}_{1.4\mathrm{nm}\left(\mathrm{gly}\right)}}}\times w(\mathrm{Chl}+\mathrm{Ver})$$

(5)

$$w(\begin{array}{c}\mathrm{Ver})\end{array}=w(\begin{array}{c}\mathrm{Chl}+\mathrm{Ver})\end{array}-w(\begin{array}{c}\mathrm{Chl})\end{array}$$

(6)

$$w(\mathrm{I}/\mathrm{S})=100-w(\mathrm{S})-w(\mathrm{I})-w(\mathrm{Kaol}+\mathrm{Chl})$$

(7)

Note: I_0.7nm(air): the diffraction peak intensity at 0.7 nm on the K-air spectrum. I_{0.7nm(550 °C)}: the diffraction peak intensity at 1.0 nm on the K-550 °C spectrum. h_0.358nm(gly): the peak height at 0.358 nm on the Mg-gly spectrum; h_0.353nm(gly): the peak height at 0.353 nm on the Mg-gly spectrum; h_1.76nm(gly): the diffraction peak intensity at 1.76 nm on the Mg-gly spectrum. h_0.7nm(gly): the diffraction peak intensity at 1.0 nm on the Mg-gly spectrum. h_0.7nm(gly): the peak height at 0.7 nm on the Mg-gly spectrum. h_{1.4nm(550 °C)}: the peak height at 1.4 nm on the K-550 °C spectrum. h_1.4nm(gly): The peak height at 1.4 nm on the Mg-gly spectrum.

3. Results

3.1. C Content and Optical Indices of Dissolved Organic Matter (DOM)

DOM is among the most dynamic and active organic C pools in terrestrial ecosystems [16]. Spectral characteristics and C content of DOM (C_DOM) serve as an effective measure to assess the state and dynamic changes in subterranean organic C [17]. This study examined the spectral characteristics of DOM and C_DOM across three soil profiles (Dark-brown soil, Meadow soil, and Paddy soil), with detailed data provided in Figure 2 and

Table 1. C_DOM and optical indices of DOM in samples of each layer of the three soil types.

Soil Types	Layers of Soil Profile	CDOM (mg/L)	Fluorescence Index (FI)	Autochthonous Contribution (BIX)	Humification Index (HIX)
Dark-brown soil	Ap	0.231 c	2.1 b	0.7 b	5.5 a
	AB	0.232 c	2.2 ab	0.7 b	4.7 ab
	Bt	0.243 b	2.2 ab	0.7 b	4.6 b
	C	2.212 a	2.3 a	0.8 a	0.9 c
Meadow soil	Ap	0.277 a	2.1 b	0.8 c	6.3 a
	AB	0.236 b	2.2 a	0.8 c	4.5 b
	Bg	0.277 a	2.1 b	0.9 b	3.1 c
	Br	0.274 a	2.2 a	1.0 a	2.4 d
Paddy soil	Ap1	0.351 s	2.0 d	0.7 c	5.8 a
	Ap2	0.340 b	2.3 a	0.9 a	2.6 d
	Br	0.231 c	2.1 c	0.8 b	3.1 c
	C	0.200 d	2.2 b	0.8 b	5.5 b

Note: Lowercase letters denote statistically significant differences among various soil layers within the single soil profile at the 0.05 significance level, based on one-way analysis of variance (ANOVA). Post-hoc multiple comparisons were conducted using both the LSD test and Tukey’s S-B(k) test. The sample size for each group was n = 3.

. In Dark-brown soil, C_DOM increased relatively steadily from 0.231 mg/L in the Ap layer to 0.243 mg/L in the Bt layer, followed by a sharp rise to 2.212 mg/L in the C layer. For Meadow soil, C_DOM exhibited a significant decrease in the AB layer but remained stable across other layers. In Paddy soil, C_DOM progressively declined with increasing depth, from 0.351 mg/L to 0.200 mg/L.

FI is used to quantify the proportion of microbial sources in SOM, and could be traced back to the humic origins of DOM. When FI ≤ 1.4, DOM was considered predominantly derived from plant materials such as roots, stems, and leaves, indicating a terrestrial source. Conversely, FI ≥ 1.9 denoted that DOM originated mainly from microbial metabolism and degradation products, indicative of an endogenous source. FI values between 1.4 and 1.9 suggested a mixture of microbial and plant-derived sources [13]. In this study, all soil profiles exhibited FI values exceeding 1.9, while the lowest FI value was observed in the Ap or Ap1 layer.

BIX reflects the proportion of newly produced, autogenic DOM within the overall DOM pool [18]. A higher BIX value signifies fresher DOM. Across all treatments in this study, BIX values ranged from 0.7 to 1.0. While BIX in Dark-brown and Meadow soils showed slight increases with soil depth, Paddy soil exhibited distinct phasic behavior, characterized by an initial increase followed by a subsequent decline.

HIX is used to evaluate the humification degree and aromaticity of DOM [19]. Higher HIX value indicated stronger humification and greater aromaticity, with values exceeding 4.0 suggesting a highly humified state [20]. Significant variances in HIX values were observed across the three soil profiles. In Dark-brown soil, HIX values decreased progressively from 5.5 in the Ap layer to 0.9 in the C layer. Meadow soil exhibited a similar trend, with values declining from 6.3 in the Ap layer to 2.4 in the Br layer. Meanwhile, Paddy soil demonstrated a nonlinear vertical distribution, with HIX values decreasing sharply from 5.8 in the Ap1 layer to 2.6 in the Ap2 layer, followed by an increase to 3.1 in the Br layer and 5.5 in the C layer, indicating clear stratification during the humification process.

3.2. C Content of Humus Components and C_HA/C_FA

As shown in Figure 3a, the C_WSS contents in the Dark-brown soil, Meadow soil, and Paddy soil exhibited distinct patterns of profile differentiation. In Dark-brown soil, the C_WSS dropped sharply from 1.79 g/kg in the Ap layer to 0.38 g/kg in the C layer, marking a significant reduction of 78.8%. Meadow soil demonstrated a more gradual decline, with C_WSS decreasing from 1.79 g/kg in the Ap layer to 1.58 g/kg in the Br layer—an overall reduction of just 11.7%. Paddy soil exhibited a unique C_WSS pattern, characterized by an initial increase followed by a decrease. Specifically, its C_WSS rose from 2.81 g/kg in the Ap1 layer to a peak of 3.26 g/kg in the Ap2 layer, before falling to 1.72 g/kg in the C layer—a reduction of 38.8% when compared to the Ap1 layer.

Figure 3b illustrates the profile distribution of C_HE across three representative soil types, highlighting significant regional variations. In Dark-brown soil, C_HE decreased from 4.97 g/kg in the Ap layer to 0.77 g/kg in the C layer, marking a reduction of 84.5%. In Meadow soil, C_HE declined from 3.06 g/kg in the Ap layer to 1.15 g/kg in the Br layer, reflecting a decrease of 62.4%. In Paddy soil, C_HE dropped sharply from 16.58 g/kg in the Ap1 layer to just 0.64 g/kg in the C layer, representing a dramatic reduction of 96.1%.

Figure 3c illustrates the vertical distribution of C_HA across three distinct soil types. In Dark-brown soil, C_HA decreased by 85.5%, from 2.76 g/kg in the Ap layer to 0.40 g/kg in the C layer. Meadow soil showed a more gradual decline, with C_HA dropping from 1.85 g/kg in the Ap layer to 0.57 g/kg in the Br layer, resulting in a cumulative loss of 69.2%. In contrast, Paddy soil underwent the most dramatic reduction, with a cumulative loss rate of 97.7%, as C_HA decreased from 9.75 g/kg in the upper layers to just 0.22 g/kg in the C layer.

Figure 3d illustrates that the C_HA/C_FA ratios across various soil layers exhibited a clear decreasing trend with increasing depth. Higher C_HA/C_FA ratios indicated greater conversion of FA to HA, suggesting that the HS structure was more mature and stable [21]. The elevated ratios observed in the Ap1 layer signified a more stable and mature form of HS, whereas the lower ratios found in deeper layers implied a less stable and more mobile fraction of HS [22]. In Dark-brown soil, the C_HA/C_FA ratios decreased gradually from the Ap layer (1.25) to the C layer (1.11), representing an overall reduction of just 11.2%. In contrast, Meadow soil and Paddy soil showed sharp declines in C_HA/C_FA ratios with increasing depth, by 36.4% and 64.3%, respectively.

Figure 3e illustrates that Hu was the most stable and resistant fraction of HS, with its C content serving as an indicator of HS’s long-term stability and C storage capacity [23]. In Dark-brown soil, C_Hu was notably high in the Ap layer, measuring 11.51 g/kg, indicating a significant accumulation of stable OM. However, deeper within the soil profile, C_Hu decreased markedly, dropping sharply to 3.05 g/kg in the AB layer and further declining to 1.50 g/kg in the Bt layer. Upon reaching the C layer, C_Hu stabilized at 0.96 g/kg. Among the three soil types studied, Meadow soil exhibited the lowest overall C_Hu, with minimal variation throughout its profile. The C_Hu ranged from 1.27 g/kg in the Ap1 layer to 0.78 g/kg in the C layer. By comparison, Paddy soil demonstrated a notably high content of C_Hu in the Ap1 layer, measuring 16.94 g/kg; The C_Hu rapidly decreased to 12.29 g/kg in the Ap2 layer, continued declining to 0.53 g/kg in the Br layer, and finally stabilized at 0.45 g/kg in the C layer.

3.3. Elemental Composition and FTIR Spectra of HA

FTIR was extensively employed to analyze the structural characteristics and functional group composition of HS [24]. To assess variations in FTIR across different soil types and profile depths, the changes in HA functional groups in Dark-brown soil, Meadow soil, and Paddy soil were compared, as shown in Figure 4.

As shown in

Table 2. The relative FTIR intensity of HA (% of total area) in different profile soil layers of three soil types.

Soil Types	Layers of Soil Profile	cm−1
		3428–3437	2922–2925a	2852–2855b	1773–1779d	1602–1604e	1443–1449c	1029–1047	882	(a + b + c)/d	(a + b + c)/e
Dark-brown soil	Ap	37.7	7.25	2.29	3.35	13.4	19.8	19.5	2.37	8.76	2.19
	AB	42.9	9.29	2.68	2.46	7.60	21.7	19.4	2.47	13.7	4.43
	Bt	39.7	11.1	2.80	2.35	6.20	22.1	17.8	3.67	15.3	5.81
	C	36.7	10.6	3.10	1.72	4.35	18.7	15.5	5.07	18.8	7.45
Soil Type	Layers of Soil Profile	cm−1
		3423–3442	2921–2923a	2852–2856b	1773–1779d	1600–1604e	1430–1440c	1027–1047	882	(a + b+ c)/d	(a + b + c)/e
Meadow soil	Ap	39.3	8.63	2.17	2.69	11.4	19.6	22.0	2.17	11.3	2.67
	AB	42.5	8.19	2.04	2.35	11.1	20.9	20.7	2.32	13.2	2.80
	Bg	38.4	7.67	2.11	1.99	6.21	18.0	19.8	2.39	14.0	4.47
	Br	36.6	7.86	2.00	1.45	5.87	18.2	19.6	2.47	19.4	4.78
Soil Type	Layers of Soil Profile	cm−1
		3424–3443	2920–2924a	2852–2856b	1773–1779d	1601–1604e	1437–1443c	1028–1047	882	(a + b+ c)/d	(a+ b + c)/e
Paddy soil	Ap1	34.8	6.53	2.09	2.16	13.3	19.8	19.5	1.85	13.2	2.14
	Ap2	38.3	7.34	1.81	2.01	12.2	18.6	19.1	2.20	13.8	2.27
	Br	36.0	7.76	2.05	1.84	11.7	19.1	18.7	2.20	15.7	2.47
	C	34.6	7.68	2.38	1.61	11.2	16.3	17.8	2.14	16.4	2.35

Note: The data presented in Table 2 represented the best result obtained from three repeated trials.

, the broadband range of 3428–3437 cm⁻¹ was typically associated with O–H stretching vibrations [25]. For all three soil types, the peak intensity observed within this range initially increased before subsequently decreasing. The band in the range of 1021–1050 cm⁻¹ corresponded to the asymmetric stretching or vibration of C–O in polysaccharides [26]. As soil depth increased, the polysaccharide content in Dark-brown soil, Meadow soil, and Paddy soil decreased by 20.5%, 10.9%, and 8.7%, respectively. The band at 882 cm⁻¹ was attributed to the stretching vibration of Fe–O [27]. Meadow soil and Paddy soil showed increases of 13.8% and 15.7%, respectively, while Dark-brown soil demonstrated the most significant increase, at 113.9%. The vibration intensities of the aliphatic −CH₃ and −CH₂ absorption peaks were observed at 2921–2925 cm⁻¹, 2852–2856 cm⁻¹, and 1443–1449 cm⁻¹, respectively [28]. These peaks were designated as a, b, and c. The stretching vibration of −COOH occurred in the range of 1773–1779 cm⁻¹ and was denoted by d [29]. The (a + b + c)/d ratio served as an indicator of the relative proportion of aliphatic C to carboxyl C groups. A lower aliphatic C/carboxyl C ratio suggested that the HA had undergone more extensive decomposition and oxidation processes, leading to a relative enrichment of −COOH [30]. The (a + b + c)/d ratio of HA for Dark-brown soil, Meadow soil, and Paddy soil increased with depth, showing respective increases of 114.6%, 71.7%, and 24.2%. The molecular stretching vibration peak of the aromatic C=C double bond was observed in the range of 1602–1604 cm⁻¹ and was denoted by e [17]. The (a + b + c)/e ratio represented the relative proportion of aliphatic C to aromatic C and served as an indicator for assessing the degree of aromatization of HA [31]. The (a + b + c)/e ratio for Dark-brown soil, Meadow soil, and Paddy soil increased by 240.2%, 79.0%, and 9.8%, respectively.

As shown in

Table 3. Elemental composition and atomic ratios of HA in different profile soil layers of three soil types.

Soil Type	Layers of Soil Profile	C (%)	N (%)	H (%)	O (%)	H/C Ratio	C/N Ratio	O/C Ratio
Dark-brown soil	Ap	19.65 a	1.02 a	0.01 d	27.9 a	0.006 d	22.5 d	1.07 a
	AB	14.38 b	0.20 b	0.03 c	20.5 b	0.025 c	83.9 c	1.07 a
	Bt	14.17 c	0.14 c	0.04 b	18.0 c	0.034 b	118.1 b	0.95 b
	C	13.74 d	0.02 d	0.05 a	13.2 d	0.044 a	801.5 a	0.72 c
Meadow soil	Ap	15.96 a	0.71 a	0.02 b	42.9 a	0.015 c	26.2 d	2.02 a
	AB	14.79 b	0.35 b	0.02 b	25.6 b	0.016 c	49.3 c	1.30 a
	Bg	11.24 c	0.14 c	0.02 b	17.6 c	0.021 b	93.7 b	1.17 b
	Br	11.05 c	0.02 d	0.04 a	16.0 c	0.043 a	644.6 a	1.09 c
Paddy soil	Ap1	26.19 a	1.79 a	0.06 a	35.3 a	0.027 c	17.1 c	1.01 a
	Ap2	21.31 b	1.38 b	0.06 a	28.1 b	0.034 bc	18.0 c	0.99 a
	Br	14.65 c	0.11 c	0.05 b	17.8 c	0.041 b	155.4 b	0.91 b
	C	9.69 d	0.04 d	0.06 a	11.4 d	0.074 a	282.6 a	0.88 b

Note: Lowercase letters denote statistically significant differences among various soil layers within the single soil profile at the 0.05 significance level, based on one-way analysis of variance (ANOVA). Post hoc multiple comparisons were conducted using both the LSD test and Tukey’s S-B(k) test. The sample size for each group was n = 3.

, the contents of C, oxygen (O), nitrogen (N), and hydrogen (H) in HA showed a marked decline with increasing soil depth. Specifically, in Paddy soil, the relative C content of HA decreased from 26.19% at the Ap layer to 9.69% at the C layer, reflecting a substantial reduction of 63.0%. In Meadow soil, this value diminished from 15.96% to 11.05%, resulting in a decrease of 30.8%, while Dark-brown soil experienced a reduction from 19.65% to 13.74%, corresponding to a decline of 30.1%. With increasing soil depth, the N content of HA decreased by 98.0% in Dark-brown soil, 97.2% in Meadow soil, and 97.8% in Paddy soil. The H content of HA exhibited enrichment in Meadow and Dark-brown soils. In Meadow soil, H content demonstrated a 100% increase, while Dark-brown soil showed a notable increase of 400%. In contrast, the H content of HA from Paddy soil remained relatively stable across different soil layers. The O content of HA, which acted as an indicator of the presence of O−containing functional groups, displayed a downward trend with increasing soil depth. Specifically, the O contents of HA decreased by 62.7% in Meadow soil, 67.7% in Paddy soil, and 52.7% in Dark-brown soil.

The C/N ratio was a critical parameter for evaluating both maturity and stability within HA [32]. The C/N ratios of HA ranged from 22.5 to 801.5 in Dark-brown soil, 26.2 to 644.6 in Meadow soil, and 17.1 to 282.6 in Paddy soil. Extreme C/N ratios were observed in the parent materials, exceeding 600 in Meadow and Dark-brown soils. The H/C atomic ratio provided insights into the aromatic and aliphatic characteristics of HA [33]. Generally, the H/C ratio was positively correlated with the aliphatic groups of HA, yet negatively correlated with the degree of aromatization [34]. As soil depth increased, the H/C ratio of HA demonstrated significant variations across different soil types: a 633.3% increase in Dark-brown soil, a 186.7% increase in Meadow soil, and a 174.1% increase in Paddy soil. The O/C ratio reflected the presence of O–containing functional groups in HA, including –COOH, –OH, and C–O–C. As soil depth increased, the O/C ratio showed varying degrees of reduction: specifically, Dark-brown soil exhibited a 32.7% decline, Meadow soil showed a 46.0% decline, and Paddy soil reported a 12.9% decline.

3.4. Identification of Clay Mineral Composition and Determination

The results of the XRD analysis (Figure 5) revealed a high degree of consistency in the composition of clay minerals across various layers of the Dark-brown soil profile within the study area. Characteristic diffraction peaks were identified at 1.76 (1.80), 1.42, 1.01, 0.72, 0.51, 0.425, 0.357, 0.334, and 0.32 nm. Under Mg-air conditions, a distinct diffraction peak was observed at 1.42 nm, which expanded to 1.76 nm upon glycerol saturation. This expansion corresponded to a lattice increase caused by the insertion of water molecules between S layers [35]. The surface diffraction peak attributed to S was not pronounced, implying that S likely formed regular or disordered mixed-layer minerals in combination with I and Chl, such as I/S [36]. The location and morphology of these mixed-layer structures might differ from those seen in pure S. Consequently, this variation could lead to broadening or splitting of the 1.76 nm peak, complicating its identification [37]. Under Mg-air conditions, a reflection near 1.42 nm was observed, but this collapsed to approximately 1.01 nm after K-saturation, indicating the presence of Ver [13,14]. A stable and sharp diffraction peak at 0.425 nm confirmed the presence of quartz [38], while the peak at 0.32 nm corresponded to potassium feldspar. Following heat treatment at 550 °C, the peaks at 0.72 nm and 0.357 nm disappeared, likely due to the loss of interlayer −OH groups, which resulted in the recombination of the Kaol crystal structure [14]. According to crystallographic theory, Chl was expected to exhibit characteristic peaks at 1.42, 0.72, 0.47, and 0.357 nm [39]. However, no independent peak was detected at 0.47 nm, and interference from mixed signals of Ver and Kaol was observed at 1.42, 0.72, and 0.357 nm. Using semi-quantitative analysis, the Chl content was determined to range between 3.1% and 13.4%. The XRD analysis identified a homologous clay mineral composition in Meadow soil and Paddy soil, predominantly consisting of S, Ver, I, Chl, Kaol, and I/S. Nevertheless, notable differences were observed in the relative abundance of these minerals between the three soil types (p < 0.05).

Clay minerals played a crucial role in soil, influencing the transformation, stability, and functionality of SOM through mechanisms such as physical protection, chemical stabilization, and biological regulation [40]. Semi-quantitative X-ray diffraction (XRD) analysis (Table S2) revealed notable spatial variations in the clay mineral composition of Dark-brown soil, Meadow soil, and Paddy soil. These soils were predominantly composed of 2:1 silicate minerals (I, S, Ver, and I/S), which accounted for 78.2% to 86.4% of the total clay minerals, along with 1:1 minerals (Kaol and Chl). Across the different profiles of Dark-brown soil, Meadow soil, and Paddy soil, systematic changes were observed in both clay mineral composition and ICI.

From the perspective of material sources, the mineralogical composition of the parent material emerged as a primary determinant in the formation of clay minerals, directly influencing the types of weathering products in the soil and subsequently shaping the genesis of the clay minerals. The lower layers of the soil profile reflected certain intrinsic characteristics of the parent material. Specifically, the I content decreased by 47.5%, 47.6%, and 53.5% from the bottom to the top layers in Dark-brown soil, Meadow soil, and Paddy soil, respectively. Similarly, the S content demonstrated a downward trend from the deeper layers to the Ap1 layer, with reductions of 74.3%, 73.5%, and 75.9% in the respective soil types.

In contrast, a progressive increase in I/S was observed from the bottom to the top layers, with rises of 272.6%, 292.1%, and 254.5% in Dark-brown soil, Meadow soil, and Paddy soil, respectively. Additionally, the Ver content exhibited significant increases of 102.9%, 178.9%, and 95.5% in Dark-brown soil, Meadow soil, and Paddy soil, respectively. The Kaol content gradually increased from the bottom to the surface, with growth rates of 144.1%, 58.9%, and 74.4% for Dark-brown soil, Meadow soil, and Paddy soil, respectively. Conversely, the Chl content showed a decreasing trend from the bottom to the surface layer, with reductions of 61.2%, 58.5%, and 60.8% in Dark-brown soil, Meadow soil, and Paddy soil, respectively.

The Kübler index, defined as the half-peak width at 1.0 nm, served as a reliable and consistent indicator that exhibited a negative correlation with illite crystallinity [41]. Its values showed distinct trends with increasing soil depth. In Dark-brown soil, the index decreased significantly by 44.9%, dropping from 0.69 to 0.38. Meadow soil demonstrated a reduction from 0.60 in the Ap layer to 0.50 in the Bg layer, followed by a slight rebound to 0.61 in the C layer. Similarly, in Paddy soil, the index declined from 0.68 to 0.41, marking a decrease of 39.7%. The ICI, defined as the ratio of peak area at 0.5 nm to peak area at 1.0 nm, represented the relative abundance of Fe/Mg-illite (biotite and mica) when the ratio was less than 0.5, indicating physical weathering. Conversely, when the ratio exceeded 0.5, it reflected aluminum-rich illite (muscovite), suggesting strong chemical hydrolysis [42]. Similar to the Kübler index, the ICI also diminished with increasing soil depth: it dropped from 1.56 to 0.33 in Dark-brown soil, from 0.93 to 0.78 in Meadow soil, and from 0.88 to 0.41 in Paddy soil.

4. Discussion

4.1. Variations in the Fluorescence Index of DOM and Its C Content (C_DOM) with Depth in Dark-Brown Soil, Meadow Soil, and Paddy Soil

C_DOM showed a significant correlation between organic C stability and vertically structured microbial communities [17]. In the Dark-brown soil profile, C_DOM slightly increased from the Ap layer to the Bt layer (0.231–0.243 mg·L⁻¹), followed by a sharp rise to 2.212 mg·L⁻¹ in the C layer. This was mainly due to the low-oxygen conditions and nutrient deficiency in the C layer, which not only limited the survival of most microorganisms but also reduced biological interference [43]. As a result, the decomposition rate of DOM decreased, leading to the preservation of C_DOM [43]. At the same time, DOM could be either adsorbed onto mineral surfaces or co-precipitated with minerals, contributing to the stabilization of organic C [44]. The C_DOM in the AB layer of Meadow soil significantly decreased, while it remained relatively stable in other layers. The AB layer, as a transitional zone near the surface, had favorable aeration and received considerable inputs of fine root exudates and undecomposed plant residues [45]. It could be inferred that the microorganisms in this layer exhibited rapid metabolic rates, thereby accelerating the degradation of C_DOM [45]. The decrease in C_DOM in Paddy soil with increasing depth was closely related to the greater accumulation of plant residues in the topsoil and the slower decomposition of OM under waterlogged conditions. In contrast, the lower soil layers received fewer external organic inputs and maintained a more stable redox environment, which led to a relatively lower concentration of C_DOM in the deeper soil layers [46].

The FI values across all soil profiles were above 1.9, indicating that DOM mainly originated from microbial metabolic activities and degradation products, showing predominantly endogenous characteristics. This was closely linked to the actual situation of long-term straw burning and ineffective returning to the field for the three types of soil under corn cultivation conditions. The lowest FI values were found in the Ap and Ap1 layers, suggesting that although microbial sources still dominated in the surface soil, there might have been a minor influence from limited fresh plant residue inputs. As the main zone for plant residue input, the surface layer contained a small amount of partially degraded plant-derived OM, which could slightly reduce the relative contribution of microbial sources. This reflected minor exogenous influences occurring under the dominance of endogenous DOM sources. The BIX value range of 0.7–1.0 indicated that DOM in all three soil types contained a notable proportion of recently formed autotrophic OM. The BIX of Dark-brown soil and meadow soil moderately increased with depth, showing that although the total OM content in subsoil might be lower, the proportion of newly formed microbial-derived DOM was relatively higher. This was attributed to reduced external inputs in the subsoil environment, where microorganisms were more actively involved in decomposing and transforming native OM, leading to the continuous accumulation of freshly produced endogenous DOM [23,47]. The BIX of Paddy soil displayed a distinct phased pattern, characterized by an initial increase followed by a decrease. This variation was primarily caused by the conversion of paddy fields to dry fields and the influence of alternating dry–wet hydrological conditions across different soil layers [48]. In the Ap1 layer, frequent tillage enhanced the regular input of fresh OM. In contrast, in the Ap2 layer, alternating dry and wet conditions contributed to the temporary accumulation of native DOM [48]. As soil depth increased, processes such as re-oxidation and adsorption became more prominent in deeper layers, including the Br and C layers, ultimately reducing the freshness of DOM [49]. The vertical variation of HIX directly reflected differences in the degree of humification and aromaticity of DOM. The HIX values in Dark-brown soil and Meadow soil decreased with increasing depth, indicating that surface soils underwent a more complete humification process [50]. This was attributed to greater plant residue inputs at the surface and prolonged microbial activity promoting the accumulation of aromatic compounds [50]. In contrast, lower OM inputs and shorter humification periods in subsurface layers resulted in reduced HIX values. The non-linear distribution of HIX in paddy soil was closely associated with its unique tillage practices. Specifically, the Ap1 layer showed a high degree of humification due to frequent tillage and alternating wetting-drying cycles (HIX = 5.8). In the Ap2 layer, dryland farming in paddy fields altered the conditions for OM decomposition, leading to a sharp decline in humification. In deeper layers (Br and C), long-term anaerobic conditions further enhanced the humification process, resulting in distinct stratified characteristics [51]. These characteristics highlighted the significant impact of converting paddy fields to dry fields on the vertical material transformation within the soil profile.

4.2. Variations in Humus Composition with Depth in Dark-Brown Soil, Meadow Soil, and Paddy Soil

The vertical distribution of HS fractions, including C_WSS, C_HA, C_HE, and C_Hu across Dark-brown soil, Meadow soil, and Paddy soil profiles, revealed distinct humification pathways influenced by pedogenic conditions, redox dynamics, and land-use practices. Surface layers of Paddy soil exhibited higher concentrations of HS compared to those in Dark-brown and Meadow soils. This phenomenon was attributed to enhanced humification processes that occurred over mineralization under alternating aerobic and anaerobic conditions induced by the conversion of paddy fields to dry fields [52]. In contrast, the well-aerated surface layers of Dark-brown and Meadow soils favored rapid OM mineralization, thereby limiting the accumulation of HS [53]. Depth-dependent declines in C_WSS, C_HA, and C_HE were particularly pronounced in Dark-brown and Paddy soil due to reduced OM inputs and microbial activity [54]. On the contrary, the C contents of HS fractions in Meadow soil at greater depths remained relatively stable, which reflected the enrichment of OM in the marsh subsoil environment due to periodic flooding [55].

The C_WSS profiles indicated microbial accessibility and substrate recalcitrance [56]. In Dark-brown soil, microbial activities, dominated by the mineralization of unstable matrices, facilitated the decomposition and utilization of WSS, compounded by limited fresh OM content in deeper layers [57]. In Meadow soil profiles, C_WSS enrichment in the Ap layer was primarily driven by vegetation litter inputs and the direct supply of root exudates [58]. Compared to other zonal soils, C_WSS in Meadow soil exhibited the least reduction, which was attributed to its significantly higher total content of 2:1 phyllosilicates (S, I, and I/S) relative to Dark-brown and Paddy soil. The high specific surface area and cation exchange capacity of these phyllosilicates created ideal conditions for organic–mineral complexation [59,60]. Paddy soil’s unique C_WSS pattern displayed a nonlinear trend, first increasing and then sharply decreasing. This bimodal model underscored the “soil memory effect” in Paddy soil, where the geochemical legacy of the REDOX cycle persisted despite shifts in management practices [61].

From the perspective of the vertical attenuation law, the concentrations of C_HE and C_HA in all three soil types exhibited a significant decreasing trend from the top to the bottom layers. The attenuation amplitches of C_HE and C_HA in Paddy soil (C_HE decreased by 96.1%, C_HA decreased by 97.7%) were significantly higher than those in Dark-brown soil (C_HE 84.5%, C_HA 85.5%) and Meadow soil (C_HE 62.4%, C_HA 69.2%), and the Ap2 layer showed an obvious stratification effect. Beneath the Ap2 layer, both C_HA and C_HE demonstrated a pronounced decline, indicating that the Ap2 layer significantly impeded the vertical transport of OM below the Ap2 layer; the input of plant residues, root secretions was low, resulting in low contents of C_HE and C_HA [62]. The C_HA and C_HE attenuation amplitudes of Dark-brown soil were relatively low. The deep OM input in Dark-brown soil was predominantly derived from root residues [63]. Due to the periodic alternation between dry and wet conditions resulting from fluctuations in groundwater levels, a portion of the HA and HE in the Meadow soil remains accumulated within the Br layer [50].

The minimal decrease in C_HA/C_FA ratios observed in Dark-brown soil was likely due to the covalent complex formation between Fe³⁺ and the phenolic hydroxyl groups in HA, which contributed to their sustained stabilization [64]. Conversely, the sharp decline in C_HA/C_FA ratios in Paddy soil was attributed to anaerobic conditions fostering FA accumulation by inhibiting phenolic polymerization—a crucial pathway for HA formation—under reducing environments, thus enhancing the formation of low-molecular-weight organic components [65]. For Meadow soil, however, the decrease in C_HA/C_FA ratios reflected a transitional HS structure shaped by periodic root exudates combined with moderate microbial decomposition processes [66].

Hu, as a recalcitrant HS component, exhibited a higher degree of polymerization compared to HA and FA, reflecting divergent stabilization mechanisms [23]. In Dark-brown soil, lignin and cellulose, the primary inputs of litter, contributed an aromatic structure that formed stable polymers via free radical-mediated oxidative coupling reactions under aerobic conditions. These polymers established the core skeleton of C_Hu [67]. With increasing soil depth, biological activity diminished, inhibiting the formation of Hu [68]. In Meadow soil, the consistently low C_Hu suggested that organic C sources in the Ap layer originated predominantly from the root exudates of shallow-rooted herbs and maize. These exudates primarily comprised carboxyl-rich, low-molecular-weight organic acids (e.g., oxalic acid, malic acid) and neutral sugars (e.g., glucose, xylose) [69]. The short-chain polyphenols generated during decomposition exhibited limited capacity to form stable macromolecular polymers [70]. Additionally, the vertical distribution of roots was restricted, with root biomass concentrated within the Ap layer [71]. This concentration led to limited C_Hu accumulation in deeper layers due to substrate constraints. In contrast, unusually high C_Hu (16.94 g/kg) observed in the Paddy soil Ap1 layer might be attributed to the repeated incorporation of crop residues and the anaerobic preservation of aromatic compounds [72,73]. However, the C_Hu in C layer Paddy soil dropped sharply to 0.45 g/kg, which was attributed to the intense leaching caused by the periodic alternation between dry and wet conditions, and the anaerobic environment in the deep layer hindered the stable preservation of Hu [74,75].

4.3. Variations in Functional Groups and Elemental Composition of HA with Depth in Dark-Brown Soil, Meadow Soil, and Paddy Soil

As a key component of soil, HA structural complexity was significantly influenced by soil type and vertical stratification [76,77]. In dryland ecosystems, photosynthetic by-products, including lignin and cellulose, accumulate on the surface as precursors for HA synthesis [78]. With increasing soil depth, the pronounced reduction in plant-derived C and N inputs restricted HA synthesis, resulting in a decline in C, O, and N contents of HA molecules [79]. This decline was most obvious in Paddy soil, where the plough layer formed by converting paddy fields to dry fields impeded the downward movement of OM [80]. When converted to dryland farming, the Ap2 layer in Paddy soil remained unbroken, leading to notably higher C and N content of HA in the Ap1 and Ap2 layers compared to Dark-brown soil and Meadow soil. Nonetheless, OM content decreased rapidly below the Ap2 layer, causing a significant drop in C and N contents within HA [81].

In Dark-brown and Meadow soils, a substantial increase in H content indicated an upward trend in functional groups like –CH₃ and –CH₂–, while a decline in aromatic C=C bonds suggested reduced microbial activity and a slower humification process in deeper soil layers [82]. With soil depth, the C/N ratio of HA rose sharply, illustrating that N-rich compounds in surface soil were preferentially utilized by microorganisms, leaving behind more resistant aliphatic and aromatic residues [83]. During the early stages of soil formation, the parent material layer exhibited an extreme C/N ratio (>600), as the interaction between OM and minerals was not well-established, leading to a lowered efficiency of biologically driven humification [55]. Although the conversion from paddy fields to dry fields had occurred, the long-term utilization of paddy fields under continuous N application had resulted in significant N accumulation in the soil. A portion of this accumulated N was present in the form of organic N within HA, leading to a C/N ratio that was lower than those observed in Dark-brown and Meadow soils [84]. Furthermore, increasing soil depth triggered de-aromatization and deoxygenation of HA [85]. In the Ap (Ap1) layer, active plant residues and abundant microorganisms facilitated OM decomposition and humification, enriching functional groups like –OH and –COOH [86]. These –OH enrichments in the AB and Ap2 layers were linked to the O-rich Ap (Ap1) layer, where microbial activity promoted the formation of stable O–containing functional groups. Conversely, the decrease in the O level in the layers below the AB or Ap2 layers leads to partial depolymerization or oxidative cleavage of HA [87]. However, as soil depth increased, the sources of animal and plant residues diminished, microbial population sharply declined, and O₂ content in soil decreased, collectively reducing OM decomposition rates and reducing the abundance of –OH and –COOH groups [88].

The (a + b + c)/e ratio and H/C ratio rose with soil depth due to the presence of more aliphatic structures in HA at deeper soil layers [89]. Surface soil, rich in O and metabolic activity, favored the microbial consumption of aliphatic C and facilitated humification into more aromatic structures through peroxidation and polymerization reactions [90]. The aliphatic C/aromatic C ratio and H/C ratio in Dark-brown soil exhibited a more pronounced increase with depth compared to those in Meadow and Paddy soils. This trend suggested the incorporation of fine root fragments derived from deep-rooted vegetation, such as trees or perennial shrubs, during the formation of Dark-brown soil. These root materials were enriched in resilient lipids, including keratin and phospholipids [63]. Additionally, during HA downward migration, these low-molecular-weight aliphatic hydrocarbons, such as short-chain fatty acids and terpenoids, entered the C layer. The smaller lipid molecules, being more mobile than aromatic macromolecules, were prone to accumulating in deep soil through pore filtration [91]. In Meadow soil’s Ap layer, strong oxidation conditions supported the dissolution and polymerization of O-containing compounds, enabling condensation reactions with lignin-based aromatic structures [90]. With depth, reduced microbial activity slowed the decomposition of aliphatic compounds, thereby increasing the aliphatic C/aromatic C ratio [92]. In Paddy soil, the Ap2 layer provided physical barriers to aliphatic compounds, and long-term anaerobic or intermittently weak oxidation conditions minimized decomposition differences between surface and deep layers, reducing lipid preservation advantages [93]. Dark-brown soil, with unique root-derived lipid inputs, faced reduced oxidation efficiency of aliphatic compounds at deeper layers due to declining O content [63]. In contrast, Meadow soil under waterlogged conditions sustained a stable redox potential, mitigating oxidation gradients [94]. However, the frequent alternation of wet and dry conditions could accelerate HA decomposition, diminishing –COOH concentration [95]. Polysaccharides played a crucial role in HS formation, serving as both substrates for microbial degradation and active intermediates in polymerization reactions [96]. In Dark-brown soil, humification processes were relatively intense, and surface polysaccharides might serve as energy sources for microbial activity [81]. In contrast, Meadow soils, which were characterized by high groundwater levels and prolonged waterlogging, tended to promote anaerobic microbial activity, thereby slowing the decomposition of organic matter and maintaining polysaccharide content in HA [55]. Paddy soils, subject to alternating dry and wet conditions, exhibited relatively stable polysaccharide levels. During dry periods, aerobic microbial activity promotes polysaccharide decomposition, whereas subsequent flooding inhibits this process [97]. The relative abundance of Fe³⁺ increased in deeper soil layers, favoring the formation of stable Fe–O bonds within HA complexes [98], particularly in Dark-brown soils derived from basalt. When Fe³⁺ accumulated in deeper layers due to leaching, converting paddy fields to upland systems could elevate soil O levels, thereby enhancing iron oxidation and reducing its capacity to complex with HA [99].

4.4. Variations in Clay Mineral with Depth in Dark-Brown Soil, Meadow Soil, and Paddy Soil

This study systematically compared the vertical differentiation in structural characteristics of clay minerals in the profiles of Dark-brown soil, Meadow soil, and Paddy soil. XRD analysis (Table 2) showed that the proportion of 2:1 clay minerals (I, S, Ver, and I/S) was significantly higher than that of 1:1 clay minerals (Kaol and Chl), suggesting that these soils were in the early stages of weathering within a moderately leached environment [100]. Notably, the vertical distribution of I content exhibited a systematic decline, attributed to mineral phase transitions induced by successive hydrolysis reactions. Specifically, when interlayer K⁺ was replaced by hydrated cations (H⁺ or Ca²⁺) during leaching, this substitution disrupted the charge balance within the I lattice and initiated structural recombination following the sequence “I→I/S→S” [101]. The octahedral structure of Chl exhibited a notable susceptibility to the leaching effects of acidic substances, such as HA generated from the decomposition of OM in surface soils and organic acids secreted by plant roots [102]. This vulnerability resulted in the loss of Mg²⁺, leading to the collapse of the mineral structure and a subsequent reduction in surface content [102].

The dynamics of interlayer cations strongly influenced transformation pathways: when potassium loss between layers reached 40–60%, cation exchange caused significant lattice expansion accompanied by water molecule insertion, ultimately forming characteristic expandable 2:1 Ver structures [103]. Persistent K⁺ loss completely compromised the integrity of the interlayer support structure, leading to its reconstitution into 1:1 Kaol frameworks through desilication processes [104]. This process was most pronounced in topsoils, where successive H⁺ attacks accelerated the decomposition of 2:1 clay layers, promoting the formation of new Kaol phases and indicating advanced weathering stages [105]. Dark-brown soil exemplified this phenomenon due to its high OM content and specific oxidation conditions within the Ap layer [106]. The continuous release of protons from –COOH and –OH groups in HA facilitated the dissolution of I, with released Al³⁺ and silicic acid serving as key precursors for Kaol formation [107]. Additionally, the Kaol content exhibited a relationship with ICI. The ICI of Dark-brown soil and the C layer of Paddy soil was less than 0.5, indicating that Mg/Fe-illite was the dominant component. In contrast, the Bt layer and upper layers converted to aluminum-rich illite, confirming that Mg/Fe ions were either soluble or replaced by Al, corresponding with an increased Kaol content [100]. The vertical distribution pattern of low illite crystallinity (indicated by a high Kübler index) in the Ap or Ap1 layer, coupled with high crystallinity in deeper layers, suggested reduced intensity of weathering with increasing depth [108]. This reduction could be attributed to decreased soil permeability caused by a higher clay content or diminished mineral weathering rates due to reduced microbial attachment to mineral surfaces [109,110]. Such factors inhibited the migration of elements like K⁺ and Si⁴⁺, slowing down dealkalization and desilication processes while maintaining a certain degree of illite integrity [96]. In Meadow soil profiles, the synergistic effects of periodic flooding and freeze-thaw cycles enhanced the specific surface area of minerals, expedited the hydrolysis of edge micro-cracks induced by pore water phase transitions, and resulted in consistent weathering states of ICI and illite crystallinity in the bottom layers [111].

5. Conclusions

This study examined vertical variations in DOM properties, HS composition, HA characteristics, and clay mineral dynamics across Dark-brown, Meadow, and Paddy soil profiles. DOM in all soils was predominantly endogenous (FI > 1.9), derived mainly from microbial metabolism with minor plant residue inputs. Vertical trends in C_DOM concentrations varied by soil type: In the Dark-brown soil profile, C_DOM showed a slight increase from the Ap layer to the Bt layer and then a sharp surge in the C layer. For the Meadow soil, C_DOM decreased significantly in the AB layer while remaining relatively stable in other layers. In the Paddy soil, C_DOM exhibited a decreasing trend with increasing soil depth. Paddy soil surface layers had higher HS; C_WSS, C_HA, and C_HE declined with depth in Dark-brown/Paddy soils but stayed stable in deep Meadow soil. HA traits (C/N, functional groups, aromaticity) varied with depth and soil type: Paddy’s Ap2 layer blocked OM downward movement, Dark-brown stabilized HA via Fe³⁺ complexation, and Meadow had transitional HS. Clay mineral assemblages were dominated by 2:1 types (illite, smectite, illite–smectite interstratifications), with vertical weathering trends: illite decreased with depth via hydrolysis, while kaolinite formation was promoted by proton-driven dissolution in surface layers, especially in Dark-brown soil. These mineral transformations were closely linked to OM dynamics, with 2:1 minerals enhancing organic–mineral complexation in Meadow soil. This study’s conclusions were constrained by the focus on three soil types in specific regions, limiting broad generalization across diverse pedoclimatic zones. Additionally, short-term sampling might not capture long-term dynamics under changing land use.