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dv a n ce ub li ca ti on A r ti c l e Clue on ocean redox condition from trace element and rare earth element (REE) composition of iron formation and carbonate rocks from the late Paleoproterozoic Morar Formation, Gwalior Group, central India

Trace element and rare earth element (REE) composition of iron formation and carbonate rocks from the Morar Formation, Gwalior Group, central India provides valuable information on the redox condition of late Paleoproterozoic Ocean. Facies types of iron formation suggest deposition in various oceanic environments ranging from shoreface – beach to subtidal shelf settings, whereas carbonates belong to shallow and deep subtidal settings. La/Nd values between 0.57 and 25, MREE enrichment and small negative (0.69) to positive (1.46) Ce anomaly in iron formation suggest a strati ﬁ ed character for the Gwalior Sea with development of shallow transitional redoxcline. Whereas deep sea is interpreted as near anoxic and ferruginous, the shallow sea was not very high in dissolved oxygen (DO 2 ) either. A suboxic to mild oxic shallow sea condition (DO 2 ≥ 0.2 µM) is interpreted allowing Mn (II) oxidation and Ce sequestration. Carbonates, however, do not register any geochemical signature of redoxcline possibly because of the depositional setting either close to or below the redoxcline.


INTRODUCTION
The Proterozoic Eon (2500-541 Ma, Gradstein et al., 2012) encompasses more than 40% of the Earth's history and represents a unique time period. The non-reversible changes that the planet witnessed in the time period include amalgamation and breakup of supercontinents with growing importance of Phanerozoic-style plate tectonics (Nelson et al., 1999), dramatic oceanic and atmospheric changes on global scale (Hoffman et al., 1998), biological evolution leading to the advent of multicellular life (Knoll, 1994), and major changes in the upper crustal composition (Taylor and McLennan, 1997). None of these changes fol-lowed uniform and steady character, and especially Paleoproterozoic (2500-1800 Ma) and late Neoproterozoic (800-541 Ma) time period witnessed major biological, tectonic, climatic, and atmospheric changes. The billion-year interval (1800-800 Ma) mostly encompassing the Mesoproterozoic time is marked by tectonic, climatic, and evolutionary stability (Holland, 2006;Mukherjee et al., 2018). Nevertheless, amongst the above-mentioned breakthrough changes, the most debated one is oxygenation of Proterozoic atmosphere and hydrosphere, since it is intimately tied up with appearance and proliferation of life. It is generally believed that atmospheric O 2 partial pressure (P O 2 ) rose modestly for the first time (from <0.001 to >0.1% of present atmospheric level; PAL) at the Great Oxidation Event (GOE, 2.45Ga; Brocks et al., 1999) and further waited patiently for about two billion years to rise to a higher level in Neoproterozoic Oxygenation event (NOE; Shields-Zhou and Och, 2011). The idea, however, is contested in recent studies by Lyons et al. (2014) and Mukhopadhyay et al. (2014), wherein evidence for oxygenic photosynthesis and whiffs of rising atmospheric oxygen is put forward well before the GOE. In fact, dynamic scenarios that have emerged in recent studies include possibility of a much higher P O 2 in early stage of the GOE, which is followed by a deep plunge to a lower level (¹0.1% PAL) in next billion years until it rose again at the NOE (1000-541 Ma; Planavsky et al., 2015).
More dynamic and debated oxygenation history has been proposed for the Proterozoic hydrosphere. While many researchers (Canfield, 1998;Scott et al., 2008;Reinhard et al., 2013 and many others) support stratified ocean model with near oxic surface water and euxinic and/or ferruginous deep water in late Paleoproterozoic and Mesoproterozoic, the other view Tang et al., 2016) advocates relatively low oxygen concentration (<0.2 µmol/l) in shallow seawater till mid-Proterozoic (~1.54 Ga). Oxygenation of deep-ocean is also controversial. From disappearance of iron formation in rock record, Holland (1984) and Holland and Beukes (1990) contemplated beginning of deep ocean oxygenation at~1.8 Ga, which did not get support from studies on biomarkers of sulphur bacteria (Brocks et al., 2005) and Mo, Sr, Cr, and S isotope compositions in black shales (Arnold et al., 2004) and pyrites (Canfield, 1998(Canfield, , 2004Poulton et al., 2004). Low trace element and phosphorus (P) concentrations in marine black shales allowed researchers to believe a continuation of the sulphidicanoxic deep ocean throughout the Mesoproterozoic and parts of the Neoproterozoic until the P O 2 approached modern levels after~1 Ga. A third view (Slack et al., 2007;Tang et al., 2016) claims suboxic condition with low concentrations of dissolved O 2 but no H 2 S in the deep ocean based on REE and redox-sensitive Ce signature in hydrothermal silica-iron oxide sediments in open marine volcanogenic massive sulfide deposits.
All these studies considered rare earth element data, particularly Ce anomalies, as a robust proxy for shallowmarine redox condition. In fact, REE composition of iron formation and non-skeletal carbonates potentially preserve original oceanic signature, if the water-rock ratio between diagenetic-hydrothermal-metamorphic fluids and sedimentary rocks is low (Bau and Dulski, 1996;Planavsky et al., 2010). Also, Ce is considered a unique redox proxy because in addition to its trivalent oxidation state, like other REEs, it also exists in tetravalent oxidation state depending on the environmental oxidation condition. However, post-depositional reductive dissolution of iron oxide can significantly mobilize LREE including Ce to pore water and result in disturbance of the REE signature.
With this understanding, a number of recent studies have focused on iron formations as proxy for the emergence of oxygen in the Earth system: its initial production by at least 2.7 Ga (Kendall et al., 2010), its spread through the atmosphere between 2.45 and 2.32 Ga, the GOE (Farquhar et al., 2011), and a plunge in next billion years (mid-Proterozoic, 1.8-0.8 Ga; Reinhard et al., 2013;Tang et al., 2016). In this contribution, we use the geochemical proxy such as Ce anomaly, REE + Y and trace element signatures from iron formations and non-skeletal carbonates from the Gwalior basin, central India to throw light on possible ocean redox structure during the late Paleoproterozoic time.

Basin geology
Unconformably overlying the Bundelkhand gneiss in central India, rocks of the Gwalior Group (maximum thickness~1.8 km) are exposed over an area of 2400 km 2 (Ramakrishnan and Vaidyanadhan, 2010) (Figs. 1a and 1b). The Gwalior Group is subdivided into two formations; lower Par Formation and upper Morar Formation, the boundary between the two is marked by the presence of a meters-thick volcanic and volcaniclastic unit (Fig. 1c). The siliciclastic Par Formation, restricted all along the southern margin of the basin, is made up of medium to coarse-grained sandstones, fanglomerate/conglomerate, heterolithic sandstone-shale-carbonaceous shale; sandstones, by far, represent the largest volumetric proportion. From geochemical characters of clastics from the Par Formation, Absar et al. (2009) inferred severe weathering in the immediate basement of the basin. Conformably overlying the Par Formation, rocks of the Morar Formation (maximum thickness~800 m) represent dominant parts of the outcrop area of the Gwalior basin and composed mainly of iron formation and limestone with subordinate fine-grained argillaceous sediments, viz., ferruginous shale, black siliceous shale, and kaolinitic clay.

Geochronological framework
Two phases of dykes within the Bundelkhand Granite Massif, basement for the Gwalior basin, are dated as of 2150 and 2000 Ma through 40 Ar/ 39 Ar systematics (Mallikharjuna Rao et al., 2005). Geochronological data of the mafic rocks present within the basin have yielded dates of 1830 ± 200 Ma (Rb-Sr isochron; Ramakrishnan and Vaidyanadhan, 2010) and 1854 ± 7 Ma (U-Pb zircon concor-A d v a n c e P u b l i c a t i o n A r t i c l e dia; Deb et al., 2002). The lower age limit for the Gwalior Group is constrained at 1.83 Ga (cf. Dubey, 1930). Absar et al. (2009Absar et al. ( , 2010

Iron formation
Limited availability of outcrop caused serious constraint on the field-based characterization of iron formation. Outcrops are dominantly isolated and exposed as railway/road cuttings and quarry sections. Six different facies types identified based on bed geometry, primary sedimentary structures, lithological affinity (chert-, carbonate-or shale-associated), and mineralogy (hematite/ magnetite/Fe-chlorite). Figure 2 illustrates facies types and interpretation of paleogeographic settings that vary from shoreface-beach to subtidal distal shelf setting. Facies types MBF-1 and 2 record depositions in the wavedominated shallow marine setting ranging from the beach-upper shoreface to distal shoreface. Plane-laminated tabular ferruginous sandstones with low-angle truncation in facies MBF-1 (Fig. 3a) resemble high energy sheet flow sandstones described from upper shorefacebeach setting (Clifton, 2006), whereas muddy and wave ripple laminated (Fig. 3b) lenticular sandstones of MBF-2 bear signatures for distal shoreface within fair-weather A d v a n c e P u b l i c a t i o n A r t i c l e wave base. The occurrence of wave ripples with wavelength: amplitude ratio of 8.5:1 and chevron cross-stratification in facies MBF-2 bear indication for the operation of fair weather oscillatory flow (Sarkar et al., 1996;Chakraborty and Paul, 2008). In either of these facies types, iron can be noticed both in an autochthonous and reworked form. Quartz and hematite constitute dominant mineralogy for both MBF-1 and 2 with subordinate magnesioferrite in MBF-1 and rare alumino-phosphate, i.e. berlinite in MBF-2. The occurrence of hummocky cross-stratification (HCS) and gutter (Figs. 3c and 3d) suggests storm-induced sedimentation in facies MBF-3. Absence of any signature indicating fair weather reworking suggests depositional environment between fair weather and storm wave base, possibly in inner shelf setting. The presence of microbial (stromatolitic) structures in facies MBF-4 (Figs. 3e and 3f ) suggests the intertidal to the supratidal environment. However, absence of any signature indicating emergence and desiccation, viz. mud crack rules out supratidal origin; instead, the occurrence of small domal stromatolite columns with wavy laminations within intercolumnar spaces bear indication for the low-energy shallow subtidal setting. Iron phases in facies MBF-3 and 4 are represented dominantly by hematite with magnetite with magnesioferrite in subordinate volume. The occurrence of clay phases like montmorillonite also noticed in some samples.
The highest percentage of Fe compared to other facies types distinguishes the laminated chert-carbonateiron formation facies (MBF-5) from other facies types. The hallmark character for this facies is repetitive and regular alternations of mm-to centimeters-thick jasper and iron-rich layers laterally persisting for 100s of meter without any appreciable change in thickness (Fig. 3g). Hematite, magnesioferrite, and goethite with negligible quartz characterizes this facies. Significant lateral persistence of laminae without any outcrop-scale disruption and very low quartz content suggest low-energy distal depositional setting like the subtidal mid-shelf setting below the storm wave base. The occurrence of magnesioferrite [Mg(Fe 3+ ) 2 O 4 ], a spinel group mineral that forms a solid solution with magnetite, in a significant amount implies more anoxic condition for this facies than previously described facies types. Facies MBF-6 is characterized by alternation with fissile splintery shale and without any current or wave structure and represents the most distal setting with the highest bathymetry of the Morar Sea. Magnetite with subordinate hematite constitutes iron phases in this facies. A d v a n c e P u b l i c a t i o n A r t i c l e The Fe-rich bands of iron formation, in general, are submicron to centimeter-scale in thickness. Fe mineral phases associated with carbonate and chert (facies MBF-3, 4, and 5) are dominantly hematite and subordinately magnetite, while those alternated with shale (facies MBF-5 and 6) are Fe oxide/hydroxide recrystallized to hematite (cf. Chakraborty et al., 2015). At cases, Fe-rich bands entirely comprise of euhedral to subhedral crystals of martite, pseudomorph of magnetite (Fig. 4a) and hematite grains exhibit skeletal texture (Fig. 4c). The skeletal structure of hematite is potentially formed at the late diagenetic stage by oxidation of residual iron carbonates that might have escaped early silicification (Lascelles and Tsiokos, 2015). In addition, a few samples of massive iron formation have micro-fractures filled with goethite (colloform nature) (Fig. 4b). Fe-rich bands are sometimes disrupted by cross-cutting veins of calcite and chert (Fig.  4d).

Carbonate
Carbonates within the Morar Formation occur at two discrete stratigraphic levels, viz. i) at basal part of the Formation (Bela-ki-Bauri section); either as microbial limestones with occasional domal growth in alternation with volcaniclastics or as discrete lenticular limestone beds in alternation with dark grey shale and ii) at middle part of the Formation (Utila section); centimeter-to decimeterthick tabular beds of massive or plane-laminated micritic limestone in alternation with light-grey limy shale (Figs. 5a, 5b, and 5c).
Limestones of these sections are dominantly micritic/ sucrosic in grain size (Figs. 5a and 5b) and calcitic in mineralogy. Mosaics of sparry calcite are present in patches or within late-stage veins (Fig. 6c); the maximum spar size measured is~1000 µm. Dolomite grains, wherever present, are in the form of individual rhomb and volumetrically never exceed >6% in any of the observed thin sections. Only rarely, incipient replacement of micritic groundmass by cryptocrystalline silica (Fig. 6d) is noticed. Carbonates of the Bela-ki-Bauri section are sometimes mixed with volcaniclastics and basaltic melt resulting in replacement of micritic groundmass (Figs. 6e and 6f ). The occurrence of pyrite within the micritic groundmass can be noticed in patches (Fig. 6g).

MATERIAL AND METHODS
Iron formations (37 samples) and carbonates (29 samples) were cleaned with Milli-Q water and powder to >200 mesh size. Whole-rock geochemical analyses for major and trace elements were carried out by the X-ray fluorescence (XRF) technique using Rigaku ZXS systems equipped at the Hiroshima University, Japan. At the Hiroshima University, X-ray generated by a 3kW Rh-W dual anode tube was radiated on fused bead samples. The detailed routine procedure is illustrated in Saha et al. (2016). The analytical precision (1σ) of major elements measured in XRF ranges between 0.001 and 0.14 wt%, whereas that for minor elements ranges between 0.68 and 2.39 ppm, as outlined in Kanazawa et al. (2001). Rare earth compositions of the carbonate samples were meas- The analyzing procedure was as described by Kamei (2016). Although REEs are commonly trivalent (+3), Ce and Eu can also occur in +4 and +2 states in addition to +3 state depending on the redox condition of their carrying fluid (Alibo and Nozaki, 1999;Hohl et al., 2015). Variations in shale (Post-Archean Australian Shale; Nance and Taylor, 1976)-normalized (SN) REE pattern commonly expressed either as anomalies from the pattern (i.e., for Ce and Eu) or as the elemental ratio (e.g., Yb (SN) /Pr (SN) etc.). However, care has to be taken in assessing true negative Ce (SN) anomalies from apparent Ce anomalies, which may arise because of La overabundance in seawater. To resolve the issue, Bau and Dulski (1996) took the help of an additional chemical parameter i.e., Pr/Pr*. In absence of any chemical reason in seawater that may result in Nd or Pr anomalies, the presence of a true Ce anomaly should lead to Pr/Pr* ≠ 1 (Bau and Dulski, 1996;Fei Ling et al., 2013). Keeping (Bau and Dulski, 1996;Lawrence et al., 2006;Ling et al., 2013;Tong et al., 2016).

Iron formation
Out of 37 iron formation samples from different facies types, 13 samples were analyzed for major and trace elements and REE, while the rest of the samples were analyzed only for REE (Tables 1 and 2). Jasper samples recorded constantly high SiO 2 (92.25-99.02%) and relatively low∑Fe 2 O 3 (3.27-7.09%), while SiO 2 content showed a wider range (40.49-83.61%) and higher ∑Fe 2 O 3 (13.88-57.44%) in Fe-rich band samples. Except for two samples (BIF-Gw-1 and BIF-Gw-32), iron formation samples showed low concentrations of lithophile (Al, Ti, Zr, Th, and Sc) elements (3.62-44.05 wt%; av. 20.69 wt%), ∑REE (15.26-96.21 ppm), and Th (0.09-4.59 ppm) and co-variations between Ce/Ce * , Pr (SN) / Yb (SN) , and Y/Ho ratios with Th concentration were insig- A d v a n c e P u b l i c a t i o n A r t i c l e A d v a n c e P u b l i c a t i o n A r t i c l e nificant (Figs. 7a-7d). These features are the characteristics of low clastic input (Planavsky et al., 2010;Tang et al., 2016). Irrespective of facies types, MnO concentration is uniformly low (0.03-0.11%) and MnO/Fe 2 O 3 ratios range from 9 × 10 −4 to 1 × 10 −2 , which is lower than those recorded from hydrothermal iron hydroxide deposits of modern oxygenated oceans (e.g., Mid Atlantic Ridge precipitates; German et al., 1991German et al., , 1995Slack et al., 2007).  (Table 2 and Fig. 7e). Except for a few samples, Y/Ho ratios show relatively high values (33-50), which is close to the modern seawater values (~40-80; Nozaki et al., 1997). Only a couple of samples (BIF-Gw-5, BIF-Gw-6, BIF-Gw-14, and BIF-Gw-16) records Y/Ho values of 23.55-28.61, which is close to shale or crustal value (~27; Nozaki et al., 1997) (Table 2). Leaving aside samples with high La (marked by '*' in Table 2), La/La * and Gd/Gd * values mostly show positive anomalies (i.e., values larger than 1). Several samples show La/La * and Gd/Gd * values larger than 1.2 (grey area in Fig. 7f ), which is the characteristics of modern and Archean seawater (cf. Viehmann et al., 2015). Iron formation samples (irrespective of jasper or Fe-rich band) record a wide variation in Ce anomaly ranging from marginally negative to small positive, i.e., 0.59-1.46 (Figs. 7e and 9a).  Redox reconstruction of hydrosphere from carbonate REE data demands assessment for any effect of diagenesis, dolomitization or detrital contamination. Although it is believed that diagenesis does not alter REE pattern of carbonates in a large way (Webb et al., 2009), Ce enrichment, Eu depletion, and decrease in Dy (SN) / Sm (SN) ratio potentially occur during diagenetic alteration (Ling et al., 2013). Cross-plots of Ce/Ce * with Eu/Eu * , Dy (SN) /Sm (SN) , and ∑REE (Figs. 8a-8c) do not show any significant correlation (r 2 values 0.197, 0.029, and 0.045, respectively), suggesting insignificant diagenetic effect in the REE patterns of studied carbonate samples (Shields and Stille, 2001).
Shale-normalized REE patterns of carbonates record small positive Ce anomaly of variable degrees (Fig. 9b); the magnitudes of anomaly are 1.03-1.19 at the Utila section and 1.08-1.18 at the Bela-ki-Bauri section. Most Gwalior carbonate samples show La (SN) /Nd (SN) values >1 (Table 3) and MREE-enrichment. It is noteworthy that carbonate samples uniformly record small positive Ce anomaly unlike the modern and Archean seawater (strong negative or absence of Ce anomaly; Derry and Jacobsen, 1990;Planavsky et al., 2014).
Three samples with substantial silica replacement (Fig. 6e) show a somewhat different trend: they are characterized by high ΣREE (Fig. 8c) and La/Sm, La/Dy, La/ A d v a n c e P u b l i c a t i o n A r t i c l e A d v a n c e P u b l i c a t i o n A r t i c l e Nd values ≥1. To consider the possible influence of silicification on REE, these data are excluded from later discussion.

DISCUSSION
Fe-oxides at circum-neutral pH record REE patterns of seawater, whereas Mn oxides preferentially scavenge Ce and HREE and hence, register enrichment of those elements with respect to seawater (De Carlo et al., 2000;Ohta and Kawabe, 2001). Modern oxygenated seawater records strong negative Ce anomaly but no such anomaly is recorded when the seawater is anoxic (German and Elderfield, 1990;Byrne and Sholkovitz, 1996). In fact, the absence of any Ce anomaly in Archean and early Paleoproterozoic iron formations suggests the reducing character of basins. Both small negative (Derry and Jacobsen, 1990;Bolhar et al., 2005) and consistent small positive Ce anomalies (Slack et al., 2007) are recorded from late Paleoproterozoic hematitic iron formations. However, it is generally believed that~1.9 Ga iron formations, deposited after the initial rise of atmospheric oxygen (GOE), commonly show positive Ce anomalies (Planavsky et al., 2010).
Marginal negative to small positive Ce anomaly in iron formation (0.69-1.46) bears strong indication towards mildly oxidizing to reducing condition. Whereas negative Ce anomalies point towards an oxidizing condition in the shallow sea, positive Ce anomalies indicate that there was redox cycling of Ce in the subtidal environment of the Gwalior Sea. A stratified water column with a transitional redox zone must have existed so that Ce(IV) scav-    Bau andDulski, 1996 andLawrence et al., 2006). Note: Ce anomaly ranging from marginally negative to marginally positive. (b) Ce/Ce* versus Pr/Pr* plot for carbonate samples. P u b l i c a t i o n A r t i c l e enged by Mn-Fe oxides, organic matters and clay minerals in oxic-suboxic shallow sea went through reductive dissolution below the redoxcline, resulting in accumulation of reduced Ce(III) (Byrne and Sholkovitz, 1996;Slack et al., 2007). Chemical proxies like very low MnO concentration (0.03-0.11%), low to very low MnO/Fe 2 O 3 ratio and a variable light to heavy REE (Pr (SN) /Yb (SN) ) ratio including samples (12 out of 33 total samples analyzed; Table 2) recording values ≥1 point towards reductive Mn-Fe oxide dissolution across the redox boundary (De Carlo and Green, 2002;Tang et al., 2016). A possible modern analog may be the Black Sea where small negative to positive Ce anomaly is interpreted as a result of the stratified water column in terms of concentration of oxygen, H 2 S, dissolved and particulate Mn (German et al., 1991). The rare presence of sulfide in association with iron formation and dominant hematite mineralogy rules out anoxic, sulfidic ambient water conditions within the Gwalior Sea. Instead, a suboxic to near anoxic and ferruginous condition is interpreted below the redoxcline of Gwalior sea taking into consideration large variation in (La/Nd) SN value (0.57-25) and Y/Ho ratio (23.55-49.82). Since rapid Mn(II) oxidation can happen in in-situ O 2 concentration as low as 0.2 µM (Schippers et al., 2005), we interpret that the ambient dissolved concentration of O 2 (DO 2 ) in Gwalior deep sea was below 0.2 µM. From a study in the suboxic (microaerobic) modern ocean, Edwards et al. (2003) have shown prolific bacterial mediation in catalyzing the oxidation of Fe 2+ for their primary energy source. Observation of hematitic filament in jasper from Jerome Mines deposit, central Arizona, USA prompted Slack et al. (2007) to envisage the bacterial role in hematitic chert and iron formation in suboxic late Paleoproterozoic seawater. Except for the preliminary evidence for convex-up growth (Fig. 4b), the bacterial role in Gwalior iron formation remains to be investigated in detail. However, sedimentological evidence like the restriction of algal iron facies (MBF-IV) in the shallow subtidal environment above the fair-weather wave base allows us to infer shallow redoxcline in the Gwalior Sea (cf. Dai et al., 2004;Tang et al., 2015Tang et al., , 2016. This is in tandem with the proposition of Bellefroid et al. (2018). Considering atmospheric P O 2 values in early and middle Proterozoic between 0.1 to 1% PAL (Liu et al., 2016) and efficient water-atmosphere exchange in a wave-/tide-agitated shallow sea, Bellefroid et al. (2018) estimated oxic layer thickness in late Paleoproterozoic ocean shallower than 50 and 90 m, respectively. Iron speciation, sulfur, molybdenum isotope, and biomarker studies from many late Paleoproterozoic and Mesoproterozoic basins point towards initiation of widespread deep water sulfidic condition around 1.84 Ga (Poulton et al., 2004;Brocks et al., 2005;Sarkar et al., 2010). However, from sulfur isotope composition and model for rate-dependent sulfur isotope change in marine carbonate-associated sulfides (CAS) Kah et al. (2004) calculated marine sulfate levels only 5 to 15% of modern values for more than 1 Ga after initial oxidation of Earth's biosphere. A suboxic to the near anoxic condition that is equally compatible to account for most of the geochemical results may be a more acceptable alternative global deep ocean model (Arnold et al., 2004;Slack et al., 2007;Singh et al., 2018). Sulfidic condition possibly remained restricted within intracratonic basins and shallow shelves those had limited communication with the global ocean (Slack et al., 2007).
Geochemical signatures from carbonates are not found in full agreement with geochemical proxies from BIF. Alike most pre~1.54 Ga carbonates (Planavsky et al., 2010;Tang et al., 2016;Bellefroid et al., 2018), Gwalior carbonates, irrespective of facies type, record absence of negative Ce anomaly (1.03-1.19) and suggest minimum Ce redox cycling. Geochemical signatures in support of the presence of a redoxcline could not be found from carbonates. It is pertinent to mention here that from sedimentological evidence it is interpreted that carbonates in the Morar sea were deposited at shallow (Belaki-Bauri) or deep (Utila) subtidal setting, which may be either close to the redoxcline or below it. La/Nd values >1 and uniform positive Ce anomaly in carbonates bears indication of consistent low concentration of dissolved O 2 in Gwalior deep sea.
Since carbonates are not the efficient carrier of redox-sensitive elements during deposition (Tang et al., 2016) and the negative Ce anomaly in BIF remains well evident, it may be argued that the Morar coastline was mildly oxidizing only where the dissolved concentration of O 2 raised above 0.2 µM favoring Mn(II) oxidation. A subtle drop in DO 2 (<0.2 µM) across the redoxcline resulted in Gwalior deepwater suboxic to near anoxic and ferruginous. This is consistent with previous studies (Poulton et al., 2010;Luo et al., 2014;Zhang et al., 2018) that suggest a very low surface water oxygenation in late Paleoproterozoic and early Mesoproterozoic oceans.

CONCLUSIONS
1. REE + Y signature including small negative to positive Ce anomaly in BIF bears indication for a stratified water column in late Paleoproterozoic Gwalior sea with mildly oxidizing shallow water and reducing ferruginous deep sea. Absence of sulfide in association with BIF and dominant hematite mineral-A d v a n c e P u b l i c a t i o n A r t i c l e ogy rule out the sulfidic condition and instead allowed the proposition of suboxic to near anoxic condition with P O 2 < 0.2 µM in Gwalior deep sea. 2. The dissolved concentration of oxygen (DO 2 ) was not so high in the shallow water of the Gwalior Sea either. The shallow Gwalior sea possibly was of mildly oxidizing/ suboxic nature that could allow limited Mn(II) oxidation and Ce sequestration. Redoxcline was shallow and the absence of any significant negative Ce anomaly does not support a welloxygenated condition even in shallow Gwalior coastline. 3. Gwalior carbonates do not record any signature for redoxcline. Carbonates in Gwalior basin belong to either shallow subtidal or deeper shelf below wave base depositional setting and record La/Nd values >1, MREE enrichment and uniform positive Ce anomaly.