Assessment of the Effect of Organic Matter on Rare Earth Elements and Yttrium Using the Zhijin Early Cambrian Phosphorite as an Example

: The geochemistry of rare earth elements and yttrium (REY) in phosphorite has been widely studied. However, the effect of organic matter on REY enrichment has not been well determined. We utilized paired inorganic ( δ 13 C carb ) and organic ( δ 13 C carb ) carbon isotopes, total organic carbon (TOC), and REY content ( ∑ REY) of the Zhijin Motianchong (MTC) phosphorite and compared them with those of Meishucun (MSC) phosphorite to reveal the effect of organic matter on REY. The δ 13 C carb of the MTC area ( ≈ 0‰) is heavier than that of the MSC area ( − 5.23‰ to − 1.13‰), whereas δ 13 C org is lighter ( − 33.85‰ to − 26.34‰) in MTC than in MSC ( − 32.95‰ to − 25.50‰). Decoupled δ 13 C carb and δ 13 C org in MTC indicate the contribution of chemoautotrophic organisms or methanotrophic bacteria. Compared to the MSC phosphorite, the MTC phosphorite has higher ∑ REY and TOC, and these parameters have a positive relationship. MTC phosphorite has REY patterns resembling those of contemporary organic matter. Furthermore, dolomite cement has a higher ∑ REY than dolomite in the phosphorus-bearing dolostone. Additionally, pyrites are located on the surface of ﬂuorapatite in the Zhijin phosphorites. It is reasonable to suggest that the REY was released into the pore water owing to the anaerobic oxidation of organic matter at the interface between seawater and sediment, resulting in the REY enrichment of Zhijin phosphorites.


Introduction
The carbon isotopes of carbonate (δ 13 C carb ) record the composition of dissolved inorganic carbon (DIC), in which 13 C is relatively enriched owing to the assimilation of 12 C during photosynthesis [1][2][3]. Therefore, heavier 13 C enrichment is often associated with increased organic carbon burial [4]. However, lighter 12 C enrichment is often associated with a largely dissolved organic carbon (DOC) reservoir, overturning of anoxic seawater, shutdown of primary productivity (e.g., Neoproterozoic snowball earth), oxidation of organic carbon reservoirs, and methane hydrate destabilization [5][6][7][8][9][10][11]. Carbon isotopes of organic matter (δ 13 C org ) can be influenced by clastic organic carbon, DOC reservoir, post-sedimentary diagenesis/metamorphism, hydrocarbon contamination, and carbon isotope fractionation during primary and secondary productivity [12]. Furthermore, the difference between δ 13 C carb and δ 13 C org (∆ 13 C = δ 13 C carb − δ 13 C org ) reflects the comprehensive influence of chemical, biological, and geological processes [2]. The global mean ∆ 13 C values of >32‰, 28-32‰, and <28‰ are related to chemoautotrophs, carbon isotope fractionation produced by photosynthesis, and decreased primary fractionation, respectively [4]. In general, the coupled δ 13 C carb and δ 13 C org caused by photosynthesis are preserved in sedimentary carbonates and organic matters. However, photosynthesis hardly explains their decoupling [13,14]. Two important hypotheses regarding decoupled δ 13 C carb and δ 13 C org have been propounded: (1) δ 13 C org records DOC signals owing to the buffer influence of a large DOC reservoir [5][6][7], and (2) photosynthetic organic matter mixes with exogenous organic matter, including detrital material [15], chemoautotrophic and methanotrophic organisms [8,13,16], and oil from hydrocarbon source rock [17]. In addition, there was an obviously vertical δ 13 C gradient from surface seawater to deep seawater during the Ediacaran-Cambrian transition [18,19]. Furthermore, there was obvious redox stratification in the ocean during this period, with oxic surface water and anoxic or euxinic deep water [20]. Ocean stratification is considered to be an important factor controlling the vertical δ 13 C gradient [8,13,21].
In general, the rare earth elements and yttrium (REY) can be divided into light REY (LREE, La-Nd), middle REY (MREE, Sm-Ho), and heavy REY (HREE, Er-Lu + Y). Previous studies have suggested that microbial breakdown of buried organic matter plays an important role in phosphorite formation [22,23]. This could be supported by the Post-Archean Australian Shale (PAAS)-normalized REY pattern showing HREE depletion relative to MREE in sedimentary apatite, which is similar to that of organic matter [24][25][26][27]. Although the sum of REY (∑REY) exhibits a positive correlation with TOC in phosphorites in northern Iran, the ∑REY only varies from 87.9 to 292 ppm [28]. In contrast, phosphorites were enriched with higher REY concentrations (>500 ppm) in the Enoch Valley Mine, United States, and they showed no correlation between TOC and ∑REY [29]. Overall, the relationship between organic matter and REY enrichment remains controversial.
Zhijin phosphorite began to form in the early Cambrian Period and it was enriched with a high quantity of REY. The whole rock exhibits a positive correlation between REY and P 2 O 5 [30], and the ∑REY in fluorapatite is approximately 2000 ppm [31]. Previous studies suggested that REY was incorporated into fluorapatite from pore water, and PAASnormalized samples exhibit REY patterns of HREE depletion relative to MREE [32][33][34][35][36]. Furthermore, degradation of organic matter occurred during the REY enrichment of Zhijin phosphorite [36]. However, there is no direct evidence to support this viewpoint in Zhijin phosphorite. The contemporary Meishucun (MSC) phosphorus deposit has a ∑REY of approximately 400 ppm [37,38], which is much lower than that of Zhijin phosphorite. The TOC in Zhijin is relatively higher (0.5-5.2 wt.%) than that of MSC (<0.1 wt.%) [18,39], indicating a more important role regarding organic matter in Zhijin than in MSC even though they formed in the same geological period and similar paleogeographic position. Therefore, whether organic matter influenced REY enrichment in Zhijin compared with MSC must be ascertained.
To explain the effect of organic matter on REY enrichment, we conducted a study focused on the drill hole ZK2407 through the phosphorus rock series in the Motianchong (MTC) ore block, Zhijin deposit, South China. The δ 13 C carb , δ 18 O, δ 13 C org , TOC, and geochemical compositions of the phosphorus rock series and dolomites were analyzed. These results will be helpful in improving our understanding of the mechanism of REY enrichment in phosphorite deposits.

Paleogeography
The Yangtze Block developed a platform facies, a deep-water basin facies, and a transition zone between the two facies during the early Cambrian Period [18]. The platform facies are dominated by shallow-water carbonate deposits (e.g., Yanjiahe Formation), the basin facies by black shales and chert deposits (e.g., Niutitang formation and Liuchapo Formation), and the transition zone between the two facies by carbonate and black shale interbedding [18,26,40]. The early Cambrian Period was one of the most important phosphorusforming periods worldwide, wherein thickly bedded phosphorites were mainly formed at the platform facies [41] (Figure 1). The Zhijin phosphorite deposit (red pentagram, Figure 1) is located on the southwest margin of the Yangtze Platform.

Paleogeography
The Yangtze Block developed a platform facies, a deep-water basin facies, an sition zone between the two facies during the early Cambrian Period [18]. The p facies are dominated by shallow-water carbonate deposits (e.g., Yanjiahe Format basin facies by black shales and chert deposits (e.g., Niutitang formation and L Formation), and the transition zone between the two facies by carbonate and bla interbedding [18,26,40]. The early Cambrian Period was one of the most importa phorus-forming periods worldwide, wherein thickly bedded phosphorites were formed at the platform facies [41] (Figure 1). The Zhijin phosphorite deposit (red gram, Figure 1) is located on the southwest margin of the Yangtze Platform.  [18]. The red pentagram represents the Zhijin Area.

Deposit Geology
The Zhijin phosphorus rock series is located in Zhijin County, Bijie City, G Province. The outcropping strata of the study area span from the Ediacaran, Ca Carboniferous, and Permian systems to the Triassic system. The Zhijin phospho series spreads toward the northeast ( Figure 2a) and a Zhijin phosphorus deposit o thickness occurs in the Gezhongwu Formation, underlain by the Ediacaran Dengy mation (dolostone) and overlain by the Lower Cambrian Niutitang Formatio shale) [30] (Figure 2b). In this study, the boundary between the Dengying Forma the Gezhongwu Formation can be interpreted as the Ediacaran-Cambrian bounda thermore, the Ni-Mo sulfide layer, which formed immediately above the Zhijin p rus deposit, has a Re-Os isochron age of 521 ± 5 Ma [42], which means that th phosphorite deposit was formed in the early Cambrian Period.

Deposit Geology
The Zhijin phosphorus rock series is located in Zhijin County, Bijie City, Guizhou Province. The outcropping strata of the study area span from the Ediacaran, Cambrian, Carboniferous, and Permian systems to the Triassic system. The Zhijin phosphorus rock series spreads toward the northeast ( Figure 2a) and a Zhijin phosphorus deposit of~20 m thickness occurs in the Gezhongwu Formation, underlain by the Ediacaran Dengying Formation (dolostone) and overlain by the Lower Cambrian Niutitang Formation (black shale) [30] (Figure 2b). In this study, the boundary between the Dengying Formation and the Gezhongwu Formation can be interpreted as the Ediacaran-Cambrian boundary. Furthermore, the Ni-Mo sulfide layer, which formed immediately above the Zhijin phosphorus deposit, has a Re-Os isochron age of 521 ± 5 Ma [42], which means that the Zhijin phosphorite deposit was formed in the early Cambrian Period.

Sampling
Representative samples of the phosphorus rock series were selected from the MTC profile, with a sampling spacing of 100 cm. The sampling site is shown as a red triangle in Figure 2a and a geologic column in Figure 2b. After sampling, the samples were cut to

Sampling
Representative samples of the phosphorus rock series were selected from the MTC profile, with a sampling spacing of 100 cm. The sampling site is shown as a red triangle in Figure 2a and a geologic column in Figure 2b. After sampling, the samples were cut to remove vertical veins, then washed, dried, and milled to 200 mesh for analysis. Furthermore, samples of phosphorites and phosphorus-bearing dolostones from the MTC profile were cut and polished into thin sections for in situ analysis.

Compositions of Major Elements, Trace Elements, and REY
For major element analysis, the sample powders were first dried at 105 • C. Then 0.66 g of sample powders were fused with a mixed flux (Li 2 B 4 O 7 -LiBO 2 -LiNO 3 , guaranteed reagent GR) at 1050 • C. After the melts cooled, they were determined by X-ray fluorescence (XRF) at ALS Minerals Co., Ltd. (Guangzhou, China). GBW07211 (phosphate ore), GBW07237 (zinc ore), and GBW07241 (tungsten ore) were used as standard materials to estimate the reliability of data. The error of the XRF analysis was less than 7.5% for major element oxides, and the detection limit of major elements was 0.01%.
For trace elements and REY analysis, 0.1 g of sample powder was melted at above 1025 • C for 30 min using a mixed flux (Li 2 B 4 O 7 -LiBO 2 , guaranteed reagent GR). Then the volume was fixed with nitric acid, hydrochloric acid, and hydrofluoric acid after cooling the molten liquid. Trace elements and REY were analyzed on an Agilent 7900 inductively coupled plasma mass spectrometer (ICP-MS) at ALS Minerals Co., Ltd. (Guangzhou, China). For the calibration of ICP-MS, a standard curve was generated using internal standard to determine the stability of the instrument, and test results were obtained after deducting background interference and interference between elements. OREAS-100a (granite and hematite breccia), OREAS-120 (greywacke, siltstone, and mudstone), and STSD-1 (stream sediment) were used as standard materials to estimate the reliability of data. The analysis error of trace elements and REE was less than 10%. The elements with a detection limit of 0.01 ppm, 0.03 ppm, 0.05 ppm, 0.1 ppm, and 2 ppm include Th, Tb, Ho, Tm, and Lu, Pr, Sm, Eu, Er, and Yb, Gd and Dy, La, Ce, Nd, Sc, and Y, and Zr, respectively. The most trace element and REY contents of blank samples were below the detection limit.

Scanning Electron Microscopy (SEM)
The images were obtained at the State Key Laboratory of Ore Deposit Geochemistry at the Institute of Geochemistry, Chinese Academy of Sciences (CAS), Guiyang. The samples were polished into thin sections and coated with Au before the experiment. Then, a highenergy electric beam focused into a microbeam of 1 µm was used to react with the minerals' surface in a JSM-7800F field emission SEM (Jeol Ltd., Tokyo, Japan), obtaining backscattered electron (BSE) and secondary electron (SE) images.

Analyses of Isotopes and Total Organic Carbon
C and O isotope values were determined at the Institute of Geochemistry, CAS, Guiyang. For δ 13 C carb and δ 18 O carb analyses,~200 µg of powder samples were collected on weighing paper and transferred to a standard 12-mL headspace sample vial. The samples were then flushed with helium using a Gasbench II device (Thermo Fisher Scientific, Breman, Germany) for 8 min. Phosphoric acid was added to dissolve the sample and produce CO 2 at 72 • C for at least 4 h. Finally, CO 2 gases were measured for δ 13 C and δ 18 O using a Gasbench II device attached to a MAT 253 gas source isotope ratio mass spectrometer (Thermo Fisher Scientific). The analytical precision was greater than 0.1‰ and 0.2‰ for δ 13 C carb and δ 18 O carb , respectively. Two international standards (IAEA-603 and IAEA-CO-8) and two Chinese standard samples (GBW04405 and GBW04416) were tested for each set of 16 samples.
For δ 13 C org analysis,~1.5 g of powder samples were reacted with 6N hydrochloric acid to remove all carbonate and phosphate minerals. The residues were washed with deionized water until they were neutralized, then dried at 60 • C. The dried samples were combusted at 950 • C to release CO 2 in a Thermo Finnigan Flash EA 2000, which determined the organic carbon isotope compositions using a Thermo Finnigan MAT 253 isotope ratio mass spectrometer. The analytical precision was greater than 0.2‰ for δ 13 C org . Three international standards (IAEA-CH-6, IAEA-CH-7, and IAEA-600) were tested for each set Minerals 2022, 12, 876 7 of 23 of 16 samples during sample analysis. The standard delta notation was used as the per mil (‰) difference from the Vienna Peedee Belemnite standard: The TOC content was determined at ALS Minerals Co., Ltd. (Guangzhou, China). Organic carbon was separated and filtered through a porous crucible after the samples were digested with dilute hydrochloric acid. The crucible was then cleaned with deionized water and dried, and the organic carbon content was quantitatively detected using an infrared sensor. The analytical precision was better than 5% for TOC. BAUX-CS4, GGC-08, GGC-09, and TOC-CS1 were used as standard materials to estimate the reliability of data.

In Situ REY Analysis Using LA-ICP-MS
We selected dolomite from MTC phosphorite and phosphorus-bearing dolostone for in situ REY analysis, which was performed using a RESOlution S-155 laser ablation system coupled to an Agilent 7900 ICP-MS at the Institute of Geochemistry, CAS, Guiyang. The laser system is a 193 nm excimer gas laser (RESOlution, Fyshwick, ACT, Australia). Helium was used as a carrier gas to enhance the transport efficiency of the ablated material. A beam of 42 µm and a frequency of 4 Hz were used. The counting time for background analysis was 20 s and the counting time for sample analysis was 40 s. The glass standards NIST610 and NIST612 were used as external calibration standards. We used the average Ca concentration of dolomites in the samples as the internal standard. The data error was less than 10%.

In Situ Major Element Analysis
In-situ major elements were obtained at the State Key Laboratory of Ore Deposit Geochemistry at the Institute of Geochemistry, CAS, Guiyang. We selected dolomites from thin sections coated with Au for major element analysis, which was performed using JXA8530F-plus EPMA produced by Jeol Ltd., Tokyo, Japan. The accelerating voltage and current were 25 kV and 10 nA, respectively. The analysis time was 30 s and the beam spot diameter was 6 µm during the elemental signal collections. Because of the influence of matrix effects, including atomic number effect (Z), absorption effect (A), and fluorescence effect (F), all data was calibrated by the ZAF method. The elements obtained were Mg, Ca, Mn, and Fe. We used dolomite as a standard material. The detection limit was 0.01%, and the analytical error was 1%-2%.

Results
The compositions of the C isotopes and concentrations of TOC of the Zhijin phosphorus rock series are shown in Table 1. The δ 13 C carb , δ 18 O carb , δ 13 C org , and TOC of the phosphorus rock series range from −3.54‰ to 0.69‰, from −13.52‰ to −1.83‰, from −33.85‰ to −26.34‰, and from 0.07% to 1.27%, respectively. Our data reflects a large decoupling between the δ 13 C carb and δ 13 C org curves from the MTC profile, whereas there are coupled curves between δ 13 C carb and δ 13 C org from the MSC profile [18] (Figure 4). Meanwhile, the δ 13 C carb in the MTC profile is heavier than that in the MSC profile, and the δ 13 C org in the MTC profile is slightly lighter than that in the MSC profile ( Figure 4). Except for data spots influenced by diagenesis, the remaining samples do not show positive correlations between δ 13 C carb and δ 18 O carb values ( Figure 5a). δ 13 C carb does not have any correlation with TOC from the MTC and MSC profiles ( Figure 6a). The TOC content of the MTC profile is higher than that of the MSC profile ( Figure 6b). The ∆ 13 C value of the MTC profile is higher than that of the MSC profile (Figure 6c and d). The ∆ 13 C from the MTC profile is not correlated with δ 13 C carb , but there is a positive correlation in the MSC profile ( Figure 6c). Both profiles exhibit a negative correlation between δ 13 C org and ∆ 13 C (Figure 6d).  The concentrations of major elements, trace elements, and REY of the phosphorus rock series from the MTC profile are shown in Table 2 and Table S1. The ∑REY values of phosphorus-bearing dolostones, phosphorus dolostones, and phosphorites are 137-558 ppm (average = 330 ppm), 481-990 ppm (average = 676 ppm), and 654-1969 ppm (average = 1477 ppm), respectively. In this study, the REY data is normalized by PAAS [27]. The The concentrations of major elements, trace elements, and REY of the phosphorus rock series from the MTC profile are shown in Tables 2 and S1. The ∑REY values of phosphorus-bearing dolostones, phosphorus dolostones, and phosphorites are 137-558 ppm (average = 330 ppm), 481-990 ppm (average = 676 ppm), and 654-1969 ppm (average = 1477 ppm), respectively. In this study, the REY data is normalized by PAAS [27]. The Ce/Ce * of the samples has no correlation with either ∑REY or Dy N /Sm N (Figure 5c,d), where Ce/Ce * = Ce N /(La N * Pr N ) 1/2 . The Y/Ho of samples exhibits a negative correlation with Al 2 O 3 , Th, Sc, and Zr, respectively (Figure 7). PAAS-normalized REY patterns of phosphorites exhibit positive La anomalies, negative Ce anomalies, positive Gd anomalies, and HREE depletions ( Figure 8). With an increase in the phosphorus content, ∑REY and TOC also increase in the MTC profile (Figure 9a,b). The ∑REY and TOC values of MSC phosphorites are lower than those of MTC phosphorites (Figure 9a,b). In situ chemical compositions of the dolomites from phosphorus-bearing dolostones and phosphorites are shown in Table 3. The laser signal line graphs of representative samples are shown in Figure S1. Furthermore, the average Ca concentration in dolomite is used as the internal standard in the software processing of laser signals, and the data are shown in Table 4. The dolomites from the phosphorus-bearing dolostones and phosphorite formations have an average ∑REY of 37.5 ppm and 75.5 ppm, respectively. The PAASnormalized REY patterns of dolomite cements in phosphorites exhibit more obvious HREE (Er-Lu) depletion and positive Gd anomalies than those of dolomites in phosphorusbearing dolostones, where Gd/Gd * = Gd N /(Eu N * Tb N ) 1/2 (Table 3, Figure 8).

Diagenetic Evaluation
It has been suggested that the primary stratigraphic signal can be modified by meteoric diagenesis based on previous δ 13 C-δ 18 O data of carbonate samples, resulting in extremely low δ 13 C carb and δ 18 O carb values (<11‰), and a positive correlation between them [43][44][45][46]. Therefore, diagenetic influences must be evaluated before discussing the stratigraphic characteristics of δ 13 C carb and δ 18 O carb [44,[46][47][48][49][50]. In the binary figure of δ 13 C carb and δ 18 O carb (Figure 5a), the gray samples may have undergone short-term diagenetic alteration; therefore, they were excluded from the carbon isotope composition curves. The remaining samples do not exhibit a positive correlation between δ 13 C carb and δ 18 O carb , indicating that they were less influenced by meteoric diagenesis (Figure 5a).
The PAAS-normalized REY patterns of typical phosphorites (e.g., Meishucun phosphorites) during the early Cambrian Period exhibit obviously positive La and Gd anomalies, and negative Ce anomalies, as well as HREE depletion characteristics [51,52], which are thought to be influenced by diagenesis. Due to this influence, Ce anomalies exhibit a positive correlation with the ∑REY, and a negative correlation with Dy N /Sm N ratios [51]. However, it is doubtful that the ∑REY in marine phosphorites can reach the REY concentrations of the Zhijin phosphorites only by diagenesis. For example, the ∑REY of lower Cambrian Meishucun and Soltanieh phosphorites under the influence of diagenesis is typically less than 500 ppm and 250 ppm, respectively [28,37]. Furthermore, there are no obvious δ 18 O carb differences between the Zhijin samples with an ∑REY of below 1000 ppm and those with an ∑REY of above 1000 ppm (Figure 5b). This result indicates that although diagenesis caused changes in δ 18 O carb in the profile, its contribution to REY enrichment was insignificant. The Ce anomalies also have no correlation with the ∑REY or Dy N /Sm N ratios in Zhijin samples (Figure 5c,d), which indicates that other factors contributed to the REY enrichment.

Controls on δ 13 Ccarb
Previous studies have reported decoupling between δ 13 Ccarb and δ 13 Corg, and good correlations between Δ 13 C and δ 13 Ccarb from the early Cambrian stratum of the Yangtze platform [13], which is mainly attributed to the variation in δ 13 Ccarb caused by diagenetic alteration [55,56], degradation of organic matter [18], and overturn of anoxic seawater [57]. In Section 5.1., we first excluded untrue δ 13 Ccarb caused by meteoric diagenesis. In the remaining real isotopic data, δ 13 Ccarb is not correlated with TOC in either the MTC or MSC profile (Figure 6a), implying that degradation of organic matter had no effect on δ 13 Ccarb. Furthermore, the contemporary profiles from shallow to deep water facies exhibit a gradually lighter δ 13 Corg in the Yangtze block during the early Cambrian Period [18,19]. The spatial variation of δ 13 C values between different sedimentary facies is attributed to the existence of the vertical δ 13 C gradient in the paleo-ocean [50]. Ocean stratification is considered to be a key controlling factor for the vertical δ 13 C gradient between shallow and deep water [8,20,21]. Owing to the influence of transgression or enhanced upwelling current, 13 C-depleted deep water input resulted in a negative δ 13 C excursion of up to −6.9‰ [18,58,59], and formed a strong correlation between Δ 13 C and δ 13 Ccarb in contemporary MSC profiles (Figure 6c). However, the MTC profile exhibits stable δ 13 Ccarb (approximately 0‰) and no correlation between Δ 13 C and δ 13 Ccarb (Figure 6c).
The carbonates of Dengying Formation during the late Ediacaran Period suffered from dissolution and erosion in different degrees due to the regression, resulting in unconformity between the Ediacaran and Cambrian Periods, and formation of paleo-karst depression [60,61]. A previous study also suggested that the thickness of the Zhijin phosphorus deposit was controlled by the paleo-karst topography of the Dengying Formation; meanwhile, the samples of this study were deposited in the basal carbonate depression of the Dengying Formation [62]. This resulted in the 13 C-depleted upwelling current during

Controls on δ 13 C carb
Previous studies have reported decoupling between δ 13 C carb and δ 13 C org , and good correlations between ∆ 13 C and δ 13 C carb from the early Cambrian stratum of the Yangtze platform [13], which is mainly attributed to the variation in δ 13 C carb caused by diagenetic alteration [55,56], degradation of organic matter [18], and overturn of anoxic seawater [57]. In Section 5.1., we first excluded untrue δ 13 C carb caused by meteoric diagenesis. In the remaining real isotopic data, δ 13 C carb is not correlated with TOC in either the MTC or MSC profile (Figure 6a), implying that degradation of organic matter had no effect on δ 13 C carb . Furthermore, the contemporary profiles from shallow to deep water facies exhibit a gradually lighter δ 13 C org in the Yangtze block during the early Cambrian Period [18,19]. The spatial variation of δ 13 C values between different sedimentary facies is attributed to the existence of the vertical δ 13 C gradient in the paleo-ocean [50]. Ocean stratification is considered to be a key controlling factor for the vertical δ 13 C gradient between shallow and deep water [8,20,21]. Owing to the influence of transgression or enhanced upwelling current, 13 Cdepleted deep water input resulted in a negative δ 13 C excursion of up to −6.9‰ [18,58,59], and formed a strong correlation between ∆ 13 C and δ 13 C carb in contemporary MSC profiles (Figure 6c). However, the MTC profile exhibits stable δ 13 C carb (approximately 0‰) and no correlation between ∆ 13 C and δ 13 C carb (Figure 6c).
The carbonates of Dengying Formation during the late Ediacaran Period suffered from dissolution and erosion in different degrees due to the regression, resulting in unconformity between the Ediacaran and Cambrian Periods, and formation of paleo-karst depression [60,61]. A previous study also suggested that the thickness of the Zhijin phosphorus deposit was controlled by the paleo-karst topography of the Dengying Formation; meanwhile, the samples of this study were deposited in the basal carbonate depression of the Dengying Formation [62]. This resulted in the 13 C-depleted upwelling current during the transgression not changing the δ 13 C carb of the MTC profile due to the restrictions of paleo-karst topography. Furthermore, the Zhijin area had a higher primary productivity due to heavier δ 13 C carb in the MTC profile compared to the contemporary MSC profile (Figure 4). the transgression not changing the δ 13 Ccarb of the MTC profile due to the restrictions of paleo-karst topography. Furthermore, the Zhijin area had a higher primary productivity due to heavier δ 13 Ccarb in the MTC profile compared to the contemporary MSC profile (Figure 4).

Controls on δ 13 Corg
The Δ 13 C caused by photosynthesis ranges from 28 to 32 during geological time [4]. The MSC profile exhibits a coupled change between δ 13 Corg and δ 13 Ccarb (Figure 4), recording a typical photosynthetic characteristic with stable Δ 13 C values of approximately 28‰ [18]. Meanwhile, the δ 13 Corg of the MSC profile was also affected by the 13 C-depleted upwelling current [18], which resulted in a strong negative correlation between Δ 13 C and δ 13 Corg (Figure 6d). However, the MTC profile exhibits decoupled δ 13 Corg and δ 13 Ccarb curves (Figure 4). The Δ 13 C also exhibits a strong negative correlation with δ 13 Corg, and the MTC profile has higher Δ 13 C and slightly lower δ 13 Corg values than those of the MSC profile (Figure 6d). Owing to the restriction of paleo-karst topography, the MTC profile records relatively stable δ 13 Ccarb values. The decoupling between δ 13 Ccarb and δ 13 Corg should be attributed to variation of δ 13 Corg.
In addition to photosynthesis, δ 13 Corg can also be influenced by clastic organic carbon [12,15], a large DOC reservoir buffering below the oxygen chemocline [5,6], post-sedimentary processes [63], and chemoautotrophic organisms or methanotrophic bacteria [4,64]. First, the δ 13 C values of most modern terrigenous organic matters range from −23‰ to −33‰ [65], which may result in decoupled δ 13 Corg and δ 13 Ccarb curves. However, land plants were scarce during the Ediacaran−Cambrian transition, until their explosive evolution during the Silurian led to a diversity of photosynthetic organisms on Earth [66,67]. Furthermore, Y has much lower marine-particle reactivity, and longer residence time than Ho [68,69]. Compared with other REY, the removal efficiency of Y from seawater is also relatively low [68,69]. Because of these phenomena, seawater has a higher Y/Ho ratio than

Controls on δ 13 C org
The ∆ 13 C caused by photosynthesis ranges from 28 to 32 during geological time [4]. The MSC profile exhibits a coupled change between δ 13 C org and δ 13 C carb (Figure 4), recording a typical photosynthetic characteristic with stable ∆ 13 C values of approximately 28‰ [18]. Meanwhile, the δ 13 C org of the MSC profile was also affected by the 13 C-depleted upwelling current [18], which resulted in a strong negative correlation between ∆ 13 C and δ 13 C org (Figure 6d). However, the MTC profile exhibits decoupled δ 13 C org and δ 13 C carb curves ( Figure 4). The ∆ 13 C also exhibits a strong negative correlation with δ 13 C org , and the MTC profile has higher ∆ 13 C and slightly lower δ 13 C org values than those of the MSC profile (Figure 6d). Owing to the restriction of paleo-karst topography, the MTC profile records relatively stable δ 13 C carb values. The decoupling between δ 13 C carb and δ 13 C org should be attributed to variation of δ 13 C org .
In addition to photosynthesis, δ 13 C org can also be influenced by clastic organic carbon [12,15], a large DOC reservoir buffering below the oxygen chemocline [5,6], postsedimentary processes [63], and chemoautotrophic organisms or methanotrophic bacteria [4,64]. First, the δ 13 C values of most modern terrigenous organic matters range from −23‰ to −33‰ [65], which may result in decoupled δ 13 C org and δ 13 C carb curves. However, land plants were scarce during the Ediacaran−Cambrian transition, until their explosive evolution during the Silurian led to a diversity of photosynthetic organisms on Earth [66,67]. Furthermore, Y has much lower marine-particle reactivity, and longer residence time than Ho [68,69]. Compared with other REY, the removal efficiency of Y from seawater is also relatively low [68,69]. Because of these phenomena, seawater has a higher Y/Ho ratio than continental crust [68,69], with seawater and terrestrial sources having Y/Ho ratios of~60 and~28, respectively [28]. Meanwhile, the influence of terrestrial sources can be tested using immobile elements (e.g., Th, Sc, and Zr) [70]. Although the samples were influenced by terrestrial sources as the Y/Ho ratios of samples exhibit weakly negative correlations with Al 2 O 3 , Th, Sc, and Zr, respectively, the samples' Y/Ho ratios (45−59) are closer to that of seawater (~60) (Figure 7). The PAAS-normalized flatly REY pattern of the upper continental crust is also quite different from that of the Zhijin phosphorite with HREE depletion (Figure 8) [27]. These results suggest that seawater-sourced REY played an important role in Zhijin phosphorites. Meanwhile, ∑REY increases with increasing organic matter content in the Zhijin phosphorus deposit (Figure 9a,b). Therefore, the terrigenous organic matter in the Zhijin profile probably did not play a significant role. Second, if a large-sized DOC reservoir was preserved in Zhijin seawater, δ 13 C carb should have exhibited dramatically negative excursions, while δ 13 C org would have been largely stable due to the buffering of DOC [7]. In fact, a completely opposite result is observed with a stable δ 13 C carb (~0‰) and a significantly fluctuating δ 13 C org in the Zhijin profile ( Figure 4). The profiles from the shelf to basin facies exhibit gradually lighter δ 13 C org during the Ediacaran-Cambrian transition [18], which also suggests that the δ 13 C org of the Zhijin area was not influenced by DOC buffering. Furthermore, the accelerated removal of organic matter from surface seawater, seafloor ventilation, and oxygen uptake also resulted in the elimination of suspended organic carbon reservoirs in seawater during the Ediacaran-Cambrian transition [7,71]. Third, the thermal degradation of sedimentary organic matter can change δ 13 C org values during diagenesis, which will result in heavier δ 13 C org values in residual organic matter due to preferential mobilization of the 12 C [72]. Previous research, however, demonstrated that the δ 13 C org variation of sedimentary organic matter was relatively small during diagenesis [65]. Although increased thermal alteration can decrease the H/C ratio of organic matter, the thermal decomposition has little effect on the carbon isotope composition of kerogen when the H/C ratio is greater than 0.2, otherwise, the kerogen is richer in 13 C [73]. Therefore, the H/C ratio of kerogen in nearly contemporary strata within the Yangtze Plate is usually >0.2, demonstrating that the δ 13 C of organic matter was not severely altered by post-sedimentary processes [74,75]. continental crust [68,69], with seawater and terrestrial sources having Y/Ho ratios of ~60 and ~28, respectively [28]. Meanwhile, the influence of terrestrial sources can be tested using immobile elements (e.g., Th, Sc, and Zr) [70]. Although the samples were influenced by terrestrial sources as the Y/Ho ratios of samples exhibit weakly negative correlations with Al2O3, Th, Sc, and Zr, respectively, the samples' Y/Ho ratios (45−59) are closer to that of seawater (~60) (Figure 7). The PAAS-normalized flatly REY pattern of the upper continental crust is also quite different from that of the Zhijin phosphorite with HREE depletion ( Figure 8) [27]. These results suggest that seawater-sourced REY played an important role in Zhijin phosphorites. Meanwhile, ∑REY increases with increasing organic matter content in the Zhijin phosphorus deposit (Figure 9a,b). Therefore, the terrigenous organic matter in the Zhijin profile probably did not play a significant role. Second, if a large-sized DOC reservoir was preserved in Zhijin seawater, δ 13 Ccarb should have exhibited dramatically negative excursions, while δ 13 Corg would have been largely stable due to the buffering of DOC [7]. In fact, a completely opposite result is observed with a stable δ 13 Ccarb (~0‰) and a significantly fluctuating δ 13 Corg in the Zhijin profile ( Figure 4). The profiles from the shelf to basin facies exhibit gradually lighter δ 13 Corg during the Ediacaran-Cambrian transition [18], which also suggests that the δ 13 Corg of the Zhijin area was not influenced by DOC buffering. Furthermore, the accelerated removal of organic matter from surface seawater, seafloor ventilation, and oxygen uptake also resulted in the elimination of suspended organic carbon reservoirs in seawater during the Ediacaran-Cambrian transition [7,71]. Third, the thermal degradation of sedimentary organic matter can change δ 13 Corg values during diagenesis, which will result in heavier δ 13 Corg values in residual organic matter due to preferential mobilization of the 12 C [72]. Previous research, however, demonstrated that the δ 13 Corg variation of sedimentary organic matter was relatively small during diagenesis [65]. Although increased thermal alteration can decrease the H/C ratio of organic matter, the thermal decomposition has little effect on the carbon isotope composition of kerogen when the H/C ratio is greater than 0.2, otherwise, the kerogen is richer in 13 C [73]. Therefore, the H/C ratio of kerogen in nearly contemporary strata within the Yangtze Plate is usually >0.2, demonstrating that the δ 13 C of organic matter was not severely altered by post-sedimentary processes [74,75].    [27]. Organic matters from Niutitang formation data were obtained from [40]. The REY data of modern seawater is magnified 10 6 times [76]. For secondary production, the carbon isotopic fractionation caused by chemoautotrophic organisms can reach −35‰, while the carbon isotopic fractionation produced by marine phytoplankton is about −28‰ [77]. Methanotrophic bacteria take up carbon from 13 C-depleted methane, causing their biomass to be 13 C-depleted by up to −15-−40‰ compared to photosynthetic organic matter [78,79]. Previous studies revealed that redox stratification developed during the Ediacaran-Cambrian transition with oxic surface seawater and anoxic/euxinic deep seawater [20]. The whole-rock Fe isotopes (~0‰) from the MSC phosphorus deposit indicate that these phosphorites were deposited under an oxic seawater environment [52]. However, the whole-rock Fe isotopes (~0-0.45‰) of the Zhijin phosphorus deposit indicate a fluctuating oxic-suboxic seawater environment [52]. The redox environment of Zhijin seawater recorded by Fe isotopes is consistent with the weakened hydrodynamic condition recorded by petrography during the Zhijin phosphorus rock series sedimentary (see Section 2.2.). The TOC content of the MTC profile is higher than that of the MSC profile ( Figure 6b); meanwhile, we can obviously see that organic matter and pyrite coexist in the Zhijin sample (Figure 3k), also indicating that the Zhijin  [27]. Organic matters from Niutitang formation data were obtained from [40]. The REY data of modern seawater is magnified 10 6 times [76].  [27]. Organic matters from Niutitang formation data were obtained from [40]. The REY data of modern seawater is magnified 10 6 times [76]. For secondary production, the carbon isotopic fractionation caused by chemoautotrophic organisms can reach −35‰, while the carbon isotopic fractionation produced by marine phytoplankton is about −28‰ [77]. Methanotrophic bacteria take up carbon from 13 C-depleted methane, causing their biomass to be 13 C-depleted by up to −15-−40‰ compared to photosynthetic organic matter [78,79]. Previous studies revealed that redox stratification developed during the Ediacaran-Cambrian transition with oxic surface seawater and anoxic/euxinic deep seawater [20]. The whole-rock Fe isotopes (~0‰) from the MSC phosphorus deposit indicate that these phosphorites were deposited under an oxic seawater environment [52]. However, the whole-rock Fe isotopes (~0-0.45‰) of the Zhijin phosphorus deposit indicate a fluctuating oxic-suboxic seawater environment [52]. The redox environment of Zhijin seawater recorded by Fe isotopes is consistent with the weakened hydrodynamic condition recorded by petrography during the Zhijin phosphorus rock series sedimentary (see Section 2.2.). The TOC content of the MTC profile is higher than that of the MSC profile ( Figure 6b); meanwhile, we can obviously see that organic matter and pyrite coexist in the Zhijin sample (Figure 3k), also indicating that the Zhijin For secondary production, the carbon isotopic fractionation caused by chemoautotrophic organisms can reach −35‰, while the carbon isotopic fractionation produced by marine phytoplankton is about −28‰ [77]. Methanotrophic bacteria take up carbon from 13 C-depleted methane, causing their biomass to be 13 C-depleted by up to −15-−40‰ compared to photosynthetic organic matter [78,79]. Previous studies revealed that redox stratification developed during the Ediacaran-Cambrian transition with oxic surface seawater and anoxic/euxinic deep seawater [20]. The whole-rock Fe isotopes (~0‰) from the MSC phosphorus deposit indicate that these phosphorites were deposited under an oxic seawater environment [52]. However, the whole-rock Fe isotopes (~0-0.45‰) of the Zhijin phosphorus deposit indicate a fluctuating oxic-suboxic seawater environment [52]. The redox environment of Zhijin seawater recorded by Fe isotopes is consistent with the weakened hydrodynamic condition recorded by petrography during the Zhijin phosphorus rock series sedimentary (see Section 2.2.). The TOC content of the MTC profile is higher than that of the MSC profile ( Figure 6b); meanwhile, we can obviously see that organic matter and pyrite coexist in the Zhijin sample (Figure 3k), also indicating that the Zhijin samples were deposited in a relatively reductive environment compared to the MSC profile. The fluctuating oxic-suboxic environment of Zhijin seawater may indicate a spatially heterogeneous redox environment during the early Cambrian Period [80].
We can clearly see pyrites located on the surface of fluorapatite under secondary electron mode (Figure 3l). The previously reported values of ∆ 34 S sulfate-pyrite range from 4.45‰ to 18.63‰ from the Zhijin phosphorites, suggesting fractionation produced by bacterial sulfate reduction [81]. The polymetallic Ni-Mo sulfide layer of the Niutitang Formation has a Re-Os age of 537 ± 10 Ma from the Zhongnancun profile of Zunyi, South China, which is consistent with the age of the Ni-Mo sulfide layer formed immediately above the Zhijin phosphorus deposit (521 ± 5 Ma) [42,82]. As a result, the formation age of black shale in the Niutitang Formation of Zunyi, South China, is almost the same as that of the Zhijin phosphorus deposit. The negative Ce anomalies of organic matter in black shales indicate that most of the organic matter came from primary productivity in surface seawater [83]. When the organic matter in surface seawater sank to the interface between seawater and sediment, 12 C was re-released into bottom seawater or pore water due to the anaerobic oxidation process of organic matter (bacterial sulfate reduction). Chemoautotrophic organisms or methanotrophic bacteria in the water column assimilated recycling 12 C, which resulted in decoupled δ 13 C carb and δ 13 C org curves in the Zhijin profile. Although the DOC concentration of modern shallow water columns close to the chemocline can reach 280 µM, the organic matter would be degraded by bacteria within several months [84,85]. The Zhijin area controlled by paleo-karst topography was less affected by marine circulation ventilation, and the formation of suboxic seawater in the Zhijin area was more conducive to organic matter recycling by chemoautotrophic organisms or methanotrophic bacteria. Furthermore, the statistical results of paired δ 13 C carb and δ 13 C org from geological history exhibit that organic matter with ∆ 13 C values of >32‰ is associated with chemoautotrophic biomass contribution [4]. The ∆ 13 C values of Zhijin samples can reach 34, indicating that organic matter in the Zhijin area was contributed to by chemoautotrophic biomass. Moreover, the statistical data from geological history shows that the intermittently decoupled δ 13 C carb and δ 13 C org from the Yangtze platform are associated with chemoautotrophs or methanotrophs [13]. Therefore, the decoupled δ 13 C carb and δ 13 C org of the MTC profile were associated with chemoautotrophic organisms or methanotrophic bacteria in the water column.

Effect of Organic Matter on REY Enrichment
The PAAS-normalized REY pattern of modern seawater exhibits a positive La anomaly, a negative Ce anomaly, and HREE enrichment [34,76] (Figure 8). The apatite from modern deep-sea mud directly inherits from seawater information [33]. However, the phosphorite from geological history does not exhibit the REY pattern that is consistent with modern seawater [86,87]. The PAAS-normalized REY patterns of these phosphorites exhibit positive La anomalies, positive Gd anomalies, negative Ce anomalies, and HREE depletions [86,87]. It is suggested that there is no clear temporal gap between HREE enrichment and HREE depletion in marine phosphate, wherein the HREE depletion characteristics of some phosphorites can be explained by diagenesis [51,86]. Because diagenesis can result in preferential enrichment of MREE [51]. In fact, the PAAS-normalized REY pattern of seawater during the geological history should be consistent with that of modern seawater [86]. The HREE depletion of marine phosphate that was not affected by diagenesis may be attributed to REY exchange with a non-clastic component [86].
Although REY was derived from seawater in Zhijin phosphorite, modern seawater has an extremely low REY concentration (~30 ppt) [76]. Previous studies had concluded that REY in apatite was derived directly from pore water [32,33,89,90]. Therefore, the REY cycling between seawater and pore water is important. The REY concentration in the pore water of marine sediment is 1-2 orders of magnitude higher than that in seawater due to REY release from host matter influenced by early diagenesis [25]. Meanwhile, the PAAS-normalized REY pattern of pore water will change from HREE enrichment to HREE depletion [34]. It is interesting to note that REY is preferentially scavenged by organic matter and then released into pore water to generate a pattern of HREE depletion [24,25]. The REY pattern of phosphate can also inherit those of organic matter [24,25,[91][92][93]. The REY patterns of HREE depletions and positive Gd anomalies from the Zhijin phosphorites are also similar to those of nearly contemporary kerogen [40] (Figure 8). In Section 5.2.2., we believe that organic matters from surface seawater participated in the process of bacterial sulfate reduction at the interface between seawater and sediment. Before organic matters sink to the interface between seawater and sediment, they can scavenge REY from seawater, leading to the first accumulation of REY in organic matters. When bacterial sulfate reduction occurred at the interface, the anaerobic oxidation of organic matters resulted in the re-release of REY into the bottom seawater or pore water, which eventually led to REY enrichment in the pore water. Moreover, REY concentration can reach 500 ppm in organic matter [40]. In particular, the ∑REY of dolomite cement in phosphorite is higher than that of dolomite in the phosphorus-bearing dolostone in the MTC profile (Table 3). These results suggest that the degradation of organic matter caused the high REY abundance in pore water.
Compared to dolomites from phosphorus-bearing dolostones (Er N /Lu N = 0.512-0.843, Gd/Gd * = 0.816-1.19), dolomite cements from phosphorites exhibit more obvious HREE depletion (Er-Lu) and positive Gd anomalies (Er N /Lu N = 0.957-1.16, Gd/Gd * = 1.11-1.41) (Figure 8). Thus, even though REY from pore water was ingested into fluorapatite, dolomite cements partially retain the geochemical characteristics of pore water. Furthermore, the Zhijin phosphorite has a higher ∑REY and TOC than that of the MSC profile (Figure 9a,b). These results imply that organic matter played a significant role in the REY enrichment process in the Zhijin phosphorus deposit ( Figure 10).
played an important role in REY enrichment. The PAAS-normalized REY patterns of dolomites in the Zhijin phosphorus rock series are consistent with that of modern seawater (Figure 8), also indicating that seawater-sourced REY is in Zhijin samples.
Although REY was derived from seawater in Zhijin phosphorite, modern seawater has an extremely low REY concentration (~30 ppt) [76]. Previous studies had concluded that REY in apatite was derived directly from pore water [32,33,89,90]. Therefore, the REY cycling between seawater and pore water is important. The REY concentration in the pore water of marine sediment is 1-2 orders of magnitude higher than that in seawater due to REY release from host matter influenced by early diagenesis [25]. Meanwhile, the PAASnormalized REY pattern of pore water will change from HREE enrichment to HREE depletion [34]. It is interesting to note that REY is preferentially scavenged by organic matter and then released into pore water to generate a pattern of HREE depletion [24,25]. The REY pattern of phosphate can also inherit those of organic matter [24,25,[91][92][93]. The REY patterns of HREE depletions and positive Gd anomalies from the Zhijin phosphorites are also similar to those of nearly contemporary kerogen [40] (Figure 8). In Section 5.2.2., we believe that organic matters from surface seawater participated in the process of bacterial sulfate reduction at the interface between seawater and sediment. Before organic matters sink to the interface between seawater and sediment, they can scavenge REY from seawater, leading to the first accumulation of REY in organic matters. When bacterial sulfate reduction occurred at the interface, the anaerobic oxidation of organic matters resulted in the re-release of REY into the bottom seawater or pore water, which eventually led to REY enrichment in the pore water. Moreover, REY concentration can reach 500 ppm in organic matter [40]. In particular, the ∑REY of dolomite cement in phosphorite is higher than that of dolomite in the phosphorus-bearing dolostone in the MTC profile (Table 3). These results suggest that the degradation of organic matter caused the high REY abundance in pore water.

Conclusions
The Zhijin area had higher primary productivity due to heavier δ 13 C carb in the MTC profile compared to the contemporary MSC profile. The decoupling of paired C isotopes results mainly from the contribution of chemoautotrophic organisms or methanotrophic bacteria.
The REY pattern of the Zhijin phosphorite is similar to that of contemporary organic matter. The ∑REY of dolomite cement in the Zhijin phosphorite is higher than that of dolomite in the phosphorus-bearing dolostone. Compared to dolomite in the phosphorusbearing dolostone, dolomite cement from phosphorite exhibits more obvious HREE depletion (Er-Lu) and a positive Gd anomaly, which records partial information of pore water. It can be concluded that the degradation of organic matter increased the REY concentration of pore water, causing REY enrichment in Zhijin phosphorite and producing REY pattern transition from HREE enrichment to HREE depletion.
Supplementary Materials: The following are available online at https://www.mdpi.com/article/ 10.3390/min12070876/s1, Figure S1: The laser signal line graphs of representative samples. (a) The laser signal line graph of dolomite cement-4 from Table 3. (b) The laser signal line graph of dolomite-2 from Table 3, Table S1: The concentrations of major elements (%), trace elements (ppm), and REY (ppm) of the Zhijin samples.