Oceanic Redox State During the Early Cambrian: Insights From Mo‐S Isotopes and Geochemistry of Himalayan Shales

The Precambrian‐Cambrian (Pc‐C) boundary marks significant biological, atmospheric, and oceanic changes. These changes include extinction of the Ediacaran fauna, initiation of complex lifeforms, and oxygenation of the atmosphere and oceans. In this contribution, elemental and Mo‐S isotopic compositions of organic‐rich shales overlying the Pc‐C boundary from the Tal Formation, Lesser Himalaya, have been investigated. These datasets are used to reconstruct past oceanic redox state and sulfate concentrations. The principal component analysis of the elemental dataset identifies six major factors, with factors associated with organic matter and sulfide phases accounting for about half of the total variance. Iron speciation and Mo/U data suggest that the shales were deposited in anoxic and ferruginous deep water conditions. The δ98Mo data (1.5 ± 0.2‰) and their mass balance calculations indicate that the areal extent of sulfidic waters and pyrite burial rates were about 2–4 times higher during the Pc‐C transition than in the modern ocean. The pyrite‐δ34S values (3.6–8.3‰) for the Tal shales are isotopically heavier compared to modern‐day sedimentary pyrites (∼−21‰). Calculations involving earlier‐reported δ34S values for early Cambrian seawater and our measured pyrite‐δ34S data estimate the seawater sulfate concentration (8 ± 3 mM) during their deposition. This sulfate value for the Tal basin is higher than that reported for the late Neoproterozoic ocean (<5 mM), attributable to increasing oxygen availability and continental supply during this time. The observed basinal conditions and high terrestrial input may have influenced metazoan diversification.


Geological Background
The Proterozoic to early Cambrian sedimentary succession of the Kumaon-Garhwal region in the Lesser Himalaya consists of autochthonous to para-autochthonous sedimentary sequences occurring between the Main Central Thrust (MCT) and the Main Boundary Thrust (MBT; Valdiya, 1980;Figure 1a).Based on lithological and geochronological constraints, these sequences are classified into older autochthonous Damtha and Tejam Groups, and the younger para-autochthonous Krol belt succession.
The Krol belt succession consists of the Jaunsar and Mussoorie Groups.The Jaunsar Group represents sedimentation on a shallow marine progradational shelf affected by occasional storms and transgressive events (Ghosh et al., 2016;Pant & Goswami, 2002).The occurrence of detrital zircons of ∼0.8 Ga and the age of the overlying Blaini Formation (∼692 Myr; Figure 1b) provide constraints for a Neoproterozoic age of the Jaunsar Group (Célérier et al., 2009;Etienne et al., 2011).The Jaunsar Group comprises phyllites and siltstones of the Chandpur Formation at the base and the orthoquartzites-slates of the Nagthat Formation.The Jaunsar Group is conformably overlain by the glaciomarine-arenaceous-argillaceous sequence of the Blaini-Infra Krol-Krol and Tal successions.The Blaini Formation consists of diamictite units separated by shale-sandstone alternations (Brookfield, 1987).A maximum depositional age of 692 ± 18 Myr (Pb-Pb detrital zircon age) has been provided for the Blaini diamictites (Etienne et al., 2011).The uppermost unit of the Blaini Formation is the Infra Krol succession, consisting of carbonaceous shales and pyritiferous phosphates deposited in an euxinic lagoonal environment (Valdiya, 2015).The Infra Krol succession grade into the shallow marine carbonates and siliciclastics of the Krol and Tal units (Valdiya, 1980).The Krol Formation is dominated by argillaceous limestones, marls, and slates in the lower parts (Krol A, B, C) and microbial dolomites (Krol D and E) in the upper parts, representing a supratidal to shallow-subtidal environment (Valdiya, 1980(Valdiya, , 2015)).The overlying Tal Formation contains a sequence of shale-chert-phosphorite assemblages.This sequence was deposited in a barred basin with an open ocean connection where restricted circulation and basin upwelling prevailed (Bhargava et al., 2021;Mazumdar et al., 1999;Valdiya, 2015).This succession is characterized by assemblages of small shelly faunal genera, such as Anabarites, Protohertzina, and Maldeotaia (Bhatt et al., 1985;Brasier & Singh, 1987).Chemo-(carbon and oxygen isotopes), bio-(acritarchs, small shelly fauna, trilobites, and brachiopods) and chrono-(535 ± 11 Myr; Singh et al., 1999) stratigraphic evidence mark the Pc-C boundary near the Krol-Tal transition (Aharon et al., 1987;Azmi, 1983;Jiang et al., 2003).An older Rb-Sr age of 626 ± 13 Myr (Sharma et al., 1991) for the shales has also been reported for the Lower Tal, which may have been influenced by sediment provenance age.

Samples and Analytical Methods
Black shale samples from the sequence overlying the Pc-C boundary in the Lesser Himalaya were collected for this study (Figures 1a-1c).The samples belonging to the Tal Formation were collected from a road-cut section at the Maldeota village (30°21′50.17″N,78° 8′30.11″E).These shale samples are mostly undeformed and least metamorphosed.During sampling, care was taken to collect fresh samples with no observable surficial weathering.Fine laminations with disseminated pyrite grains (mm sized) were observed in hand specimens of the shales.For geochemical comparison, a few shale samples from an underlying (Jaunsar) Group of Neoproterozoic age (Negi et al., 2022;Richards et al., 2005) were also collected from a road-cut section (30°30′50.40″N,77°59′49.20″E).The collected samples from both sections were powdered and homogenized (to <100 μm size) using an agate mortar and pestle.Splits of the powdered samples were used for geochemical and Mo-S isotopic analyses.

Geochemical Analyses
For major oxides, about 0.5 gm of powdered samples (n = 22), along with Li 2 B 4 O 7 -LiBO 2 flux in a fixed (1:17) proportion, were fused at 1050°C.The fused beads were measured using a wavelength-dispersive X-Ray Fluorescence Spectrometer (Bruker S8 Tiger); intensities of characteristic wavelengths were quantified to constrain the elemental concentrations.Along with the shale samples, two USGS rock standards (BHVO-2 and BCR-2) were also measured to establish the accuracy (±1%) of these measurements, whereas five replicate samples were measured to check the data reproducibility (±1%).Trace elemental analyses were carried out following the analytical methodology of Samanta et al. (2022).Briefly, about 100 mg of powdered samples were completely dissolved  Rai & Singh, 1983) and (b) generalized stratigraphic details of Neoproterozoic-Cambrian succession of the outer Lesser Himalaya (after Ghosh et al., 2016;Hughes, 2016).The proposed Pc-C boundary at the base of the Tal Formation (Bhargava et al., 1998;Tarhan et al., 2014) and tentative age information for the succession (Etienne et al., 2011;Richards et al., 2005) are also shown.(c) Lithological, carbon isotopic (Mazumdar & Banerjee, 2001), and chronological (* Singh et al., 1999; # Sharma et al., 1991; and $ Tripathy, 2011) details for the Mussoorie phosphorite mine section.
using HF-HNO 3 -HCl acids in a microwave digestion system (Anton Paar Multiwave Pro).The concentrations of trace elements in the digested solutions were measured using a quadrupole ICPMS instrument (Thermo iCAP-Q).Additionally, digested solutions of the USGS reference material (BHVO-2) and replicate samples (n = 8) were also measured for data accuracy (∼4%) and precision (∼3%).The enrichment factor for element (X) with respect to its upper continental crust (UCC) composition (Rudnick & Gao, 2003) is computed as X EF = (X/Al) sample /(X/Al) UCC .The Ce-anomaly of the samples are also computed (Taylor & McLennan, 1985) as follows: where, the subscripts, sample and PAAS, stand for the concentrations in the sample and Post-Archean Australian shales, respectively.
Total nitrogen (TN), carbon (TC), and sulfur (TS) concentrations were determined using an elemental analyzer (Elementar vario Pyro Cube).For this, about 30 mg of samples were combusted at 1150°C, and the amounts of CO 2 , N 2, and SO 2 gases were analyzed to determine the concentrations (Samanta et al., 2022).The precision of these measurements, based on replicate analyses (n = 8), was found to be ±3%.Further, the carbonate (and hence, total inorganic Carbon (TIC)) content of the shale samples was measured using a UIC CO 2 coulometer.For this, ∼25 mg of the sample was treated with ortho-phosphoric acid, and the amount of CO 2 released was quantified.The total organic carbon (TOC) content of the shales was calculated as the difference between total carbon and inorganic carbon concentrations (TOC = TC − TIC).

Iron Speciation and DOP
We have adopted the analytical methodology of Poulton and Canfield (2005) to evaluate the iron speciation of the shale samples.For this, about 100 mg of sample powders were sequentially treated with sodium acetate, sodium dithionite, and ammonium oxalate solutions to extract iron associated with carbonate (Fe carb ), oxides (Fe ox ), and magnetite (Fe mag ) phases, respectively.The pyrite-bound Fe (Fe py ) was determined using a Cr-reduction approach (Canfield et al., 1986), where the amount of precipitated Ag 2 S from the sample was gravimetrically quantified.The amount of highly reactive iron pool (Fe HR ) is calculated by adding the Fe content of the individual extractions (Fe HR = Fe carb + Fe ox + Fe mag + Fe py ).Additionally, the reactive iron pool (Fe R ) was determined by treating samples with boiling 12N HCl (Berner, 1970).The Fe content of all the extracts was determined on an Atomic Absorption Spectrometer (ThermoFisher iCE 3500) using a standard calibration approach.The Fe-extracts for the reference material WHIT (Alcott et al., 2020) were measured to check the data accuracy (±7%; n = 8).Replicate extractions of sample aliquots yielded the measurement reproducibility (±5%; n = 7).The Degree of Pyritization (DOP; Berner, 1970) value for these samples is computed as follows: (2)

Molybdenum Isotopic Measurements
Mo isotopic analyses of shale samples were carried out following the analytical approach of Goswami et al. (2022).For this, the shale samples were completely dissolved by treating with HF-HCl-HNO 3 acids (cf.Section 3.1), and the solution was measured for its Mo concentration using a Q-ICPMS instrument.Based on its Mo concentration, a known amount of the sample solution was mixed with an appropriate amount of 97 Mo-100 Mo double spike solution (Rudge et al., 2009;Siebert et al., 2001) and allowed for sample-spike equilibration.The pure Mo fraction of the sample-spike mixture was extracted using anion exchange chromatography.The Mo isotopic composition of the samples was measured on a Thermo Neptune Plus MC-ICPMS in a low-resolution mode.The sample introduction was done through an APEX-HF (Elemental Scientific Inc.) desolvating nebulizer system.All seven Mo isotopes, 92 Mo, 94 Mo, 95 Mo, 96 Mo, 97 Mo, 98 Mo, and 100 Mo, were measured using a static multi-collection mode.
In addition, the signals of 91 Zr and 99 Ru were constantly monitored to ensure no isobaric interferences.The measured data were corrected offline for instrumental mass-bias using a MATLAB code (Goswami et al., 2022;Rudge et al., 2009) to obtain the Mo isotopic abundances of samples.Finally, the Mo isotopic composition is reported in a delta notation relative to NIST SRM 3134 with a positive offset of +0.25‰ (Nӓgler et al., 2014): The internal precision of the measurements was better than 0.04‰ (2σ μ ), and the external precision of Mo isotopic analyses during the course of analyses was better than 0.05‰ (2σ).The total analytical blank for this procedure was ∼1.5 ng (n = 3), which is lower by two orders in magnitude than the total Mo processed, and hence, no blank corrections were made.

Sulfur Isotopic Measurements
To measure sulfur isotopic composition of pyrite phases (Gill et al., 2011)

Results
Geochemical and Mo-S isotopic data for the Himalayan shales (Tal Formation) overlying the Pc-C boundary are provided in Tables 1 and 2. For comparison, these data for a Neoproterozoic shale sequence from an underlying (Jaunsar) Group are also included.Concentrations for most of the major oxides for the Tal shales match well with those reported for the UCC (Rudnick & Gao, 2003; Figure 2).The total carbon in these shales, on average, is composed of ∼60% of organic and ∼40% of inorganic carbon.The average TOC for these shales is 1.3 ± 0.1%, about an order of magnitude higher than that of the Neoproterozoic shales (0.14 ± 0.06%; Table 1) from the Himalaya.The total nitrogen (TN) concentrations (0.092 ± 0.003%) of these samples exhibit a good correlation with TOC (r = 0.43, p < 0.05) and Al (r = 0.71, p < 0.001).The (atomic) TOC/TN ratios for the Pc-C shales vary between 15.2 and 17.8, with an average value of 16 ± 1.These TOC/TN ratios, although higher than the Redfield ratio (∼6), are consistent with those reported for shales with similar depositional ages (e.g., Cremonese et al., 2014;Kikumoto et al., 2014).The average phosphorous concentration in these samples is 668 μg/g.P concentrations show a good correlation with Ca (r = 0.49, p < 0.05), Mg (r = 0.65, p < 0.01), and Ti (r = 0.50, p < 0.05).
The total Fe content of the samples ranges from 3.4% to 3.7%, with an average value of 3.5 ± 0.1%, which is lower than the UCC (3.9%; Rudnick & Gao, 2003).The average Fe/Al ratio (0.48 ± 0.01) for these samples is lower than that reported for Paleozoic oxic marine shales (0.53 ± 0.11; Raiswell et al., 2008) and modern marine sediments (0.55 ± 0.11; Clarkson et al., 2014).Iron, manganese, potassium, and other trace elements (V and Cu) show a positive correlation with the Al concentrations.Most of the trace elements, except Mo and U, show limited variations, with enrichment factors of ≤1 (Figure 2c).The Mo concentrations of these shales vary between 13 and 17 μg/g, whereas U concentrations vary between 5.1 and 8.3 μg/g.The U concentrations correlate well with the TOC (r = 0.45, p < 0.05).In contrast, the Mo concentrations show a good correlation with Mn and Al, with no significant correlation with S. Molybdenum isotopic (δ 98 Mo) values of the Tal samples range from 0.99 to 1.89‰ with an average value of 1.47 ± 0.21‰ (n = 11; Table 2).Excluding two samples, the δ 98 Mo values exhibit minimal variations with a mean value of 1.48 ± 0.07‰, which is ∼0.8‰ depleted than modernday seawater value (2.3 ± 0.1‰; Goswami et al., 2022;Nakagawa et al., 2012).The δ 98 Mo values of Tal shales are comparable to those reported across the Pc-C transitions (Figure 3a).However, these values are systematically depleted than that reported for shales deposited in sulfidic environments during the Pc-C boundary (∼1.9‰;Wille et al., 2008) and early Cambrian (∼2.1‰ at ∼535 Myr; Wen et al., 2011) period.

Discussions
Elemental and isotopic compositions of organic-rich shales serve as reliable proxy for deep water redox state, primary productivity, and nutrient cycling (Tribovillard et al., 2006;Tripathy et al., 2014).The Tal Formation, as mentioned in an earlier section, is mainly composed of chert-phosphorites, carbonates, and sandstones with the occurrence of a shale horizon above the Pc-C boundary (Figure 1).Chemo-stratigraphic investigation of this section has already been reported (Banerjee et al., 1997) and is compiled in Figure 4.This study only focuses on Mo-S isotopic and trace elemental distributions for the shale horizon.Geochemical data of these shales and their principal component analyses were employed to constrain the major carrier phases of redox-sensitive elements.The intensity of deep water anoxia and seawater sulfate content was evaluated using molybdenum and sulfur isotopic data.

Trace Elements and Iron Speciation Data
Trace element abundances in organic-rich shales are mostly regulated by their redox-dependent scavenging from the water column and subsequent enrichment in underlying sediments.For instance, redox-sensitive elements (e.g., U, Mo, V, and Cr) remain mostly soluble in oxygenated conditions and get removed by forming organo-metallic complexes/sulfides in reducing conditions (Tribovillard et al., 2006).For the Tal shales, only selected trace elements (Mo, U, and Zn) exhibit high (>1) enrichment factors (Figure 2).In addition to authigenic enrichment, the abundances of these elements in marine sediments depend on continental input, hydrothermal supply, post-depositional alteration, hydrographic restriction, and particulate shuttling (Algeo & Maynard, 2008;Algeo & Rowe, 2012;Tribovillard et al., 2006).We have carried out a principal component analysis (PCA) of the geochemical dataset using the PAST (Paleontological Statistics;v. 4.08;Hammer et al., 2001) software to better constrain the major mechanisms affecting the trace element distribution.The PCA identifies six principal components (PC) with Eigen values > 1, which accounts for ∼86% of the observed data variance (Table S3 in Supporting Information S1).
Factor loadings for different elements from these components are shown in Figure 5.The PC-1, which explains ∼26% of the total variance, has high factor loadings for TOC, TN, Al, K, Mn, and several redox-sensitive trace elements (V, Cr, Co, Ni, Zn, Mo, Th, U).The observed elemental grouping possibly suggests an authigenic source for these major and trace elements.It is intriguing to observe high loading for Al, which mostly reflects lithogenic sources, in this factor.High TOC-trace element association and low loading for other lithogenic indices (such as Si and Na) establish that these Al loadings may also be linked to hydrogenous processes.We, therefore, hypothesize that an appreciable amount of Al in this factor may also be linked to Al adsorption onto clay surfaces (Timothy & Calvert, 1998).The PC-1 shows low loading for sulfur and hence may not account for the removal of elements in euxinic conditions.In contrast, the PC-2 is characterized by high S loading and explains ∼20% of the variance.This factor also has high loadings for Fe, Co, and Cu, which are usually taken up by pyrites or other sulfide phases during diagenesis (Morse & Luther, 1999).The PC-3 (variance ∼15%) has high loadings for TIC, Ca, Sr, and P, suggesting a carbonate phase association.The association of P with Ca and TIC for this component also hints at the presence of carbonate-fluorapatite minerals, which can form authigenically due to P-supersaturation in pore waters (Ruttenberg & Berner, 1993).The PC-4 (variance ∼14%) shows high factor loading for Si, Ti, Fe, K, and Mg.These elements are typically found enriched in clastic sediments, and hence, the component can be identified as continental input to the basin.In addition to these four major principal components, which together explain ∼75% of the data variance, two more components relating to macronutrients (PC-5 with variance ∼6% for Si, Ni, K, TIC, TOC, and TN) and carbonate-rich clastic fluxes (PC-6 with variance ∼5% for TIC, Th, Al, Si, Ti, and S) are observed.The PCA thus gives an indication that the elements associated with the PC-1 and PC-2 can provide better insights into reconstructing the redox state of the basin.For comparison, these data for a Neoproterozoic shale sequence (Chandpur Formation) from Lesser Himalaya have also been included.#A, B, and C stand for different powdering aliquots of the shale sample.Table 1 Continued and renewal time of bottom water (Algeo & Lyons, 2006;Algeo & Tribovillard, 2009).Uranium enrichment occurs in the sediments when the soluble uranyl tricarbonate ions (UO 2 (CO 3 ) 3 4− ) diffuse into the sediments and get either adsorbed or precipitated as uraninite and related oxides (UO 2 , U 3 O 7 , U 3 O 8 ; Klinkhammer & Palmer, 1991;Zheng et al., 2002).In the case of molybdenum, the aqueous Mo species (molybdate (MoO 4 2− )) gets converted to particle reactive thiomolybdate species (MoO x S 2− (4-x) , x = 0 to 3) in sulfidic conditions and gets scavenged by sulfur-rich organic matter and iron sulfides (Erickson & Helz, 2000;Tribovillard et al., 2004).The U enrichment exceeds that of Mo in oceanic basins with suboxic bottom water, leading to a lower Mo/U ratio in sediments compared to that of the overlying seawater.During highly reducing and occasional sulfidic conditions, the relative enrichment of Mo matches that of U, leading to a sediment Mo/U ratio comparable to the seawater value (Algeo & Tribovillard, 2009).Additionally, Mo can adsorb onto Mn-Fe-oxyhydroxides formed in oxic waters and get transported to the sediment-water interface.In bottom water conditions where O 2 content falls below 10 μM (Calvert & Pedersen, 1996), reductive dissolution of these oxides can also enhance Mo burial in the sediments.To assess this proposition, we have evaluated the Ce anomaly for the Tal shale samples.This The uncertainties for Mo isotopic data are reported as 2σ μ .b The areal coverage has been computed using the reported f oxic , f interm , and f sulfidic values.rare earth element gets adsorbed onto Mn-Fe-oxyhydroxides in oxic layers, which subsequently can get released during their reductive dissolution in anoxic conditions, and lead to its enrichment in bottom waters/pore waters and chemical sedimentary phases (German & Elderfield, 1990;Tostevin et al., 2016).The Tal shales show insignificant Ce-anomaly (Ce/Ce* = 1.01 ± 0.08) and hence record no appreciable signature for reductive dissolution.However, this proposition based on cerium data cannot be ascertained as this anomaly may be influenced by both terrigenous and hydrogenous components in shales.Hence, we have also evaluated the co-variation between Mo EF and U EF (Figure 6a) to better constrain the depositional condition.The Tal shales display a positive correlation between Mo EF and U EF (r = 0.52, p < 0.05), and these data fall closer to the modern seawater Mo/U value, suggesting a reducing depositional condition for the basin.The Mo/TOC ratios of these shales (12 ± 1) are close to the Framvaren Fjord (9 ± 2; Algeo & Lyons, 2006), which is a semi-restricted and stratified basin with a bottom water renewal time of ∼125 years (Algeo & Rowe, 2012).Strong Mo EF -U EF correlations and  (Lyons et al., 2009) and molybdenum isotopes (Poulson et al., 2006) in sedimentary systems are also depicted.Mo EF < 25 for the Tal shale hint at an anoxic bottom-water condition with sulfidic condition restricted to pore waters (Kendall et al., 2017;Scott & Lyons, 2012).

Table 2 Mo Elemental and Isotopic (δ 98 Mo) Data for the Tal Shales Along With the Estimated Mo Flux Fractions (f oxic , f interm , f sulfidic ) and Areal Extent (A oxic , A interm , A sulfidic ) for Oxic, Intermediate, and Sulfidic Sinks
Iron speciation data for shales also provide additional insight into bottom-water anoxic conditions (Poulton & Canfield, 2005).This approach relies on the abundance of reactive iron and pyrites in the shales, which depend on the basin anoxia and euxinia, respectively (Poulton & Canfield, 2005;Raiswell et al., 2018).The Fe-speciation data of the Tal shales show that about one-third of the sedimentary Fe is of pyritic nature, whereas the iron associated with carbonate (∼11%), oxide (∼3%), and magnetite (∼2%) phases are minimal.Figure 6b compares the  Fe HR /Fe T and Fe py /Fe HR ratios for the Tal shales.The threshold values for various depositional settings shown in Figure 6b are from the earlier-reported data for modern and Phanerozoic marine sediments (e.g., Poulton & Raiswell, 2002;Raiswell & Canfield, 1998).These compiled data show that the Fe HR /Fe T ratios are generally low (below 0.2) for oxic conditions, which increase to >0.38 in anoxic settings (Poulton & Canfield, 2011).High sedimentation rates or transfer of Fe HR to poorly reactive silicates during diagenesis can yield intermediate ratios (0.22-0.38;Poulton & Canfield, 2011).The Fe py /Fe HR ratios increase significantly in euxinic deposition due to the pyritization of reactive iron species and hence are higher (>0.7;März et al., 2008) in sulfidic water conditions.The iron speciation data for the Chandpur shales, with lower Fe HR /Fe T (∼0.16) and Fe py /Fe HR (∼0.06) ratios, indicate their deposition in modern-like oxic depositional settings.In contrast, the Tal shales have an average Fe HR /Fe T value of 0.48, which is greater than the oxic threshold (Figure 6b).These shales also show a range of Fe py /Fe HR values from 0.63 to 0.70 with an average ratio of 0.67, which fall close to the ferruginous-euxinic transition threshold.Further, these samples have an average DOP value of 0.51, falling in the cut-off range for anoxic-ferruginous deposition (0.45-0.75).Consistent with the Mo-U trend, the iron speciation data for the Tal shales also confirm an anoxic and ferruginous depositional environment.

Molybdenum Isotopic Constraints
The molybdenum isotopic signature of organic-rich shales has been widely used to reconstruct past seawater δ 98 Mo value and areal extent of redox state (e.g., Arnold et al., 2004;Chen et al., 2015;Kendall et al., 2011).The seawater δ 98 Mo (modern-day value ∼ 2.3‰; Nakagawa et al., 2012) is mainly regulated by its input and removal pathways.The major sources of Mo are riverine (∼0.7‰;Archer & Vance, 2008) and hydrothermal (−3.5 to 2.06‰; McManus et al., 2002;Neely et al., 2018) inputs, whereas Mo removal from the oceans occurs through its association with Fe-Mn oxyhydroxides in oxic settings, and as thiomolybdate complexes in sulfidic waters.
The δ 98 Mo values of the Tal shales range from 0.99 to 1.89‰ with an average of 1.5 ± 0.2‰.These values are marginally lower than those reported for euxinic shales during this period (∼1.9‰;Wille et al., 2008;∼2.1‰;Wen et al., 2011).This difference is likely linked to isotopic fractionation between seawater and shale in anoxic and ferruginous depositional setting of the Tal shales (cf.Section 5.1.1;Figure 6b).Observations based on a limited dataset from modern anoxic basins hint that sediments deposited in weakly reducing basins and continental margins may display an isotopic depletion of ∼0.7‰ from the contemporaneous seawater value (Dickson et al., 2014;Eroglu et al., 2020;Poulson et al., 2006;Poulson Brucker et al., 2009).Considering this offset from our mean δ 98 Mo value of 1.5‰, a qualitative estimate of the seawater during the deposition of the Tal shales is found to be ∼2.2‰.This inferred seawater δ 98 Mo for the Pc-C interval is consistent with that reported earlier for other global sections (Figure 3a).The measured seawater Mo isotopic composition and related isotopic balance equations, hence, can be used to quantify the relative Mo removal proportions for different redox pathways.
A steady-state Mo isotopic balance approach has been adopted to quantify the areal extent of oxic, intermediate reducing (or Sulfidic At Depth; SAD), and sulfidic conditions for the early Cambrian ocean (Kendall et al., 2017).Towards this, it is assumed that the Mo supply via riverine input is balanced by its removal in oxic, intermediate reducing, and sulfidic settings.The related mass balance equations are provided below: Or, where f oxic , f interm , and f sulfidic represent the fraction of Mo removed to oxic, intermediate reducing, and sulfidic sinks, respectively.The δ 98 Mo input , δ 98 Mo oxic , δ 98 Mo interm, and δ 98 Mo sulfidic stand for the isotopic compositions associated with respective sources and sinks.The Δ oxic , Δ interm, and Δ sulfidic represent Mo isotopic fractionation factors for oxic, intermediate reducing, and sulfidic sinks, respectively.Here, δ 98 Mo SW − δ 98 Mo input ≈ δ 98 Mo samples since both the isotopic offset in anoxic settings and the assumed riverine input value are ∼0.7‰.
In this calculation, the δ 98 Mo river value has been assumed to be the same as that of the modern-day value (with ∼0.7‰; Archer & Vance, 2008;King & Pett-Ridge, 2018).Support for this proposition comes from the seawater 187 Os/ 188 Os ratio, which also depends on the degree of oxidative weathering on the continents.Available Os isotopic ratios show that the initial 187 Os/ 188 Os ratios for Pc-C shales (0.8-1.2; Fu et al., 2016;Jiang et al., 2007;Kendall et al., 2009;Singh et al., 1999;Zhu et al., 2013) are similar to the modern seawater 187 Os/ 188 Os ratio (∼1.05;Lu et al., 2017), indicating a modern-like oceanic budget for osmium and hence, a modern-like riverine composition and flux.This observation, however, may not be strictly valid as δ 98 Mo river values are regulated by multiple factors such as incongruent weathering, δ 98 Mo value of exposed rocks, oxidative weathering rate, and Mo adsorption, which may have varied in the past (Kendall et al., 2015).
Equations 5 and 6 involving observed (left side of the equation) and model (right side) parameters were constructed for the 11 samples analyzed for Mo isotopes during this study (Table 2).Here, the observed parameters stand for the δ 98 Mo values for the samples, whereas the Δ values (Δ oxic, Δ interm , and Δ sulfidic ) and Mo fractions (f oxic , f interm , and f sulfidic ) serve as the model parameters.These parameters are assumed to follow a log-normal distribution for the inversion.Further, modern-like values for δ 98 Mo river and seawater Mo concentrations have been assumed for the Tal basin.We have adopted an inverse model with a Quasi-Newton optimization algorithm to find the best-fit solution for the model parameters.The computational model used in this study has already been successfully employed to constrain solute sources for water (Danish et al., 2020;Goswami et al., 2014;Tripathy & Singh, 2010) samples.The model starts an iteration from a-priori values provided for the model parameters (Δ oxic = 3.00 ± 0.14‰; Barling & Anbar, 2004;Barling et al., 2001;Siebert et al., 2003; Δ interm = 0.70 ± 0.14‰, Poulson et al., 2006; Δ sulfidic = 0.50 ± 0.30‰; Arnold et al., 2004;Chen et al., 2015;Kendall et al., 2017) to yield corresponding a-posteriori values (Δ oxic = 2.99 ± 0.01‰, Δ interm = 0.67 ± 0.05‰, Δ sulfidic = 0.39 ± 0.11‰) that can explain the equations with least residual.Using the inversion outcomes, the fraction of Mo removed (f oxic , f interm , and f sulfidic ) can be related to the seafloor area represented for each sink (A) using the following equation.
where, A represents the seafloor area covered by the sinks (in cm 2 ), B is the Mo burial flux for each sink (in nmol cm −2 yr −1 ), and F input represents the Mo input flux (3.1 × 10 17 nmol yr −1 ; Miller et al., 2011).In this evaluation, the average Mo burial flux (B) for modern-day oxic (0.021 nmol cm −2 yr −1 ), intermediate (2.61 nmol cm −2 yr −1 ), and sulfidic (12.51 nmol cm −2 yr −1 ) oceanic settings was considered (Scott et al., 2008).The assumed burial fluxes and seawater Mo inventory may have varied in the past depending on the ocean redox state, continental Mo supply, and its removal intensities.These possible changes in these parameters may also introduce uncertainties into our calculation.
Results from the inversion show that the average Mo removal fractions for oxic (0.22-0.53), intermediate reducing (0.24-0.36), and sulfidic (0.23-0.42) fractions are comparable (Figure 7a).While the average oxic removal fractions (∼0.39) are close to those estimated for the modern-oceanic settings (f oxic ∼ 0.40), the average sulfidic removal fraction (∼0.31) is about four times higher than the modern oceans (f sulfidic ∼ 0.08; Kendall et al., 2009;Poulson Brucker et al., 2009;Scott et al., 2008).Evaluation of the relative seafloor extent during the Tal shale deposition (A oxic = 99.19%;A interm = 0.67%; A sulfidic = 0.14%) and its comparison with respect to modern values (A oxic = 98.94%;A interm = 1.03%;A sulfidic = 0.03%; Table 2) indicate enhanced Mo removal and sulfidic areal extent during the Pc-C transition.The high sulfidic removal of Mo in this anoxic basin is intriguing.One possible explanation for this is sub-surface conversion of molybdates to particle-reactive thiomolybdate complexes in pore waters with limited H 2 S.
The removal of seawater sulfate mainly occurs via the formation of sulfides (pyrites) and sulfates (gypsum).
Mass balance calculations for present-day values suggest that ∼10-45% of S gets removed through pyrite burial (Tostevin et al., 2014).Additionally, sulfur associated with organic matter and carbonate phases also serve as minor sulfate sinks.The burial of sulfide minerals is a consequence of sulfate reduction through biological pathways and the subsequent conversion of released H 2 S into pyrites.This microbial sulfate reduction (MSR) process involves preferential incorporation of lighter ( 32 S) isotopes with varying degrees of isotopic fractionation (≥40‰; Canfield, 2001;Canfield et al., 2010).Additionally, the δ 34 S of sulfides are also regulated by overlying water δ 34 S value, sulfate concentration, basin restriction, reactive iron availability, diagenesis, and biological uptake rates (Canfield, 2001;Fike & Grotzinger, 2010;Gomes & Hurtgen, 2013).In contrast to sulfides, the δ 34 S value of sulfate minerals (e.g., gypsum anhydrite, barites) mostly retain the seawater δ 34 S signature, with minimal (∼1.6‰) fractionation (Raab & Spiro, 1991).
Seawater sulfate concentration and its variation are mainly linked to oceanic oxygenation and continental oxidative weathering.Previous studies on seawater sulfate estimation have adopted several approaches, such as (a) mass balance modeling (Canfield & Farquhar, 2009;Fike & Grotzinger, 2008;Kah et al., 2004;Loyd et al., 2012), (b) isotopic discrimination between co-genetic sulfides and sulfates (Algeo et al., 2015;Song et al., 2014;Wu et al., 2015), (c) numerical modeling (Li et al., 2023;Zhu et al., 2021) or (d) direct estimations from halites (Brennan et al., 2004;Horita et al., 2002;Lowenstein et al., 2003).The isotopic balance approach requires information on multiple parameters, which include sulfate inventory and related input/output fluxes, rate of sulfur isotopic changes (∂δ ocean /∂t), (open/closed) diagenetic settings, and organic substrate type and reduction rates (Algeo et al., 2015;Song et al., 2014).In the absence of these data, we have utilized the MSR trend method (Algeo et al., 2015) for estimating the sulfate concentration of the Tal basin.This approach has been established by correlating modern-day aqueous sulfate concentration [SO 4 2− ] and isotopic fractionation (Δ 34 S sulfate-sulfide = δ 3 4 S sulfate -δ 34 S sulfide ) data.This empirical relation, therefore, is less sensitive to organic substrate type, strain-specific fractionation, and sulfate reduction rates (Algeo et al., 2015).The following equation (Algeo et al., 2015) has been used for sulfate estimation for the Tal basin: The calculation yields sulfate concentrations ranging from 4 to 15 mM for the Tal basin, with an average value of 8 ± 3 mM (Figure 8a).Our calculation relies on earlier-reported seawater δ 34 S value for this period (35 ± 3‰; Kampschulte & Strauss, 2004), as co-genetic sulfate δ 34 S data were not available.In an earlier study, Mazumdar et al. (2008) have reported δ 34 S for both sulfate and pyrites for the Tal basin.Using these literature data, an average ] SW concentration of ∼7 mM is computed, which is also similar to our estimates (∼8 mM).Further, these estimates match well with the compiled [SO 4 2− ] SW data (6 ± 5 mM; n = 111; Algeo et al., 2015) for 550-520 Myr (Figure 8b).These data (Figure 8b) during the Pc-C transition are systematically higher than that reported for the Precambrian.Higher [SO 4 2− ] SW can be linked to its increased continental supply via oxidative weathering during this period.Further, Figure 8b  ] SW across the Pc-C transition for the Yangtze platform (∼2-6 mM; Li et al., 2023) is lower than that obtained for the Tal basin (∼8 mM).As both the sites had similar paleogeographic locations (Jiang et al., 2003), these differences may have been linked to the restricted connection of the Tal basin (unlike the Yangtze) to the open ocean.More studies with precise information on basinal setting, seawater δ 34 S and paleogeography can constrain causative factors for spatial variability in sulfate concentrations.

Ocean Biogeochemical Cycle of the Tal Basin During the Pc-C Transition
The above discussion establishes a sulfate concentration of ∼8 mM for the Pc-C ocean (Figure 8).Assuming modern-like fluxes for sources and sinks, this concentration corresponds to a residence time of ∼5 Myr for sulfate.This concentration and residence time are higher compared to those for the Neoproterozoic era but lower than the modern ocean (∼28 mM and ∼20 Myr).This comparison confirms high nutrient supply (compared to the Neoproterozoic) and reactive iron (Fe HR ) into the early Cambrian oceans due to intense oxidative weathering in the continents.A high nutrient flux could enhance oceanic productivity, leading to increased OM export to the seafloor and the subsequent expansion of sulfidic water masses in the ocean.Our modelling efforts utilizing Mo isotopes show elevated areal extent (∼3-4 times of modern area; Figure 7) for sulfidic waters in the global oceans, confirming this proposition.Further, the observed anoxic-ferruginous bottom water condition (Figure 6) in the Tal Basin could be a result of increased reactive iron supply sourced from either riverine supply (Raiswell & Anderson, 2005) or marine incursions into the shelf regions (Hughes et al., 2005;Mazumdar et al., 1999).Previous studies from the same section (Liu et al., 2021;Mazumdar et al., 1999), as well as global records (Hughes 2− ] for the Tal Basin using Δ 34 S sulfate-sulfide values using the MSR rate method (Algeo et al., 2015).Compiled data (gray circles) for modern aqueous systems and seawater sulfate concentration are also shown (Algeo et al., 2015).(b) Comparison of our estimated sulfate concentrations with compiled seawater sulfate data from other synchronous global sections (Algeo et al., 2015;Brennan et al., 2004;Canfield & Farquhar, 2009;Horita et al., 2002;Li et al., 2023;Lowenstein et al., 2003;Loyd et al., 2012;Ries et al., 2009;Wotte et al., 2012;Wu et al., 2015;Zhu et al., 2021Zhu et al., ). et al., 2005)), provide evidence for these processes.Fossil evidence of small shelly fauna, acritarchs, brachiopods, and trilobites in the overlying sections of the Lower Tal Formation further ascertains that environmental conditions in the Tal basin eventually became conducive for ecological expansion and sustained complex organisms.

Conclusions
Geochemical and Mo-S isotopic compositions of black shales overlying the Pc-C boundary (Tal Formation, Lesser Himalaya) have been investigated.These datasets have been used to constrain the redox state, seawater sulfate inventory, and areal extent of reducing conditions during this period.The combined data on Fe-speciation, trace elemental enrichment, and Mo isotopes indicate anoxic and ferruginous (iron-rich) deep water conditions in the shelf regions during the deposition of Tal shales.Mo isotopic data estimate higher areal extent of sulfidic waters compared to modern oceans, possibly due to increased net organic burial during the Pc-C transition.Our calculations based on pyrite-S isotopes yield a seawater sulfate concentration of 8 ± 3 mM.This concentration is higher compared to that of the Neoproterozoic ocean, confirming increased terrestrial input and oxygen in the ocean-atmospheric system during the Pc-C.These observed environmental conditions would have imparted a primary control on the subsequent expansion and diversification of complex life during the early Cambrian.

Figure 2 .
Figure 2. UCC-normalized (a) elemental ratios and (b and c) enrichment factors of selected trace elements for Pc-C (Tal Fm) and Neoproterozoic (Chandpur Fm) shales.Red dashed lines represent an enrichment factor of 1.

Figure 3 .
Figure 3. Frequency distribution of (a) δ 98 Mo isotopes and (b) Pyrite-δ 34 S of the Pc-C shales from the Himalaya.For reference, the global Mo and pyrite-S isotopic data for the early Cambrian (550-530 Myr; Farrell et al., 2021) are also shown.Expected isotopic fractionation factors for sulfur(Lyons et al., 2009) and molybdenum isotopes(Poulson et al., 2006) in sedimentary systems are also depicted.

Figure 5 .
Figure 5. Factor loadings for four major factors extracted from the principal component analysis of the geochemical dataset.

Figure 6 .
Figure 6.Co-variation plots of (a) U EF -Mo EF and (b) Iron speciation (Fe HR /Fe T and Fe py /Fe HR ) parameters for Pc-C (Tal Fm) and Neoproterozoic (Chandpur Fm) shales.For comparison, modern seawater Mo/U molar ratio (solid line) and compiled Fe-speciation data (Gray circle) for early Cambrian (550-530 Myr; Farrell et al., 2021) shales are also shown.

Figure 7 .
Figure 7. Ternary diagrams of estimated (a) Mo removal fluxes and (b) seafloor areal extents of oxic, intermediate, and sulfidic sinks (based on burial rates from Scott et al. (2008) and Asael et al. (2018)).Solid lines in the diagrams represent calculated values for different seawater δ 98 Mo values, whereas the (red-colored) dashed lines depict the calculation for modern seawater δ 98 Mo (2.3‰) value.
also exhibits large variability in the sulfate estimates from Tal and other global sections.For instance, the estimated [SO 4 2−