# Rare earth elements as indicators of post-Marinoan (~635 Ma) paleoceanographic changes from the Amazon Craton

Renan dos Santos, Afonso C.R. Nogueira, Gabriel Uhlein, Fábio Pereira, Guilherme Romero, Pierre Sansjofre, Peter Crockford, Laurane Fogret, Stefan Lalonde

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## To cite this version:

Renan dos Santos, Afonso C.R. Nogueira, Gabriel Uhlein, Fábio Pereira, Guilherme Romero, et al. Rare earth elements as indicators of post-Marinoan (~635 Ma) paleoceanographic changes from the Amazon Craton. Precambrian Research, 2024, 413, pp.107575. 10.1016/j.precamres.2024.107575. hal-04820626

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1 Rare earth elements as indicators of post-Marinoan

2 (~635 Ma) paleoceanographic changes from the Amazon

3 Craton

4 Renan F. dos Santosª, Pierre Sansjofreᵇ, Afonso C.R. Nogueiraªᶜ, Peter

5 W. Crockfordᵈ, Gabriel J. Uhleinᵉ, Laurane Fogretᵇ, Fábio S. Pereiraª, Guilherme

6 R. Romerof, Stefan V. Lalondeg

7 ª Programa de Pós-Graduação em Geologia e Geoquímica, Faculdade de Geologia,

8 Instituto de Geociências, Universidade Federal do Pará, Rua Augusto Corrêa s/n CEP

9 66075-110, Belém, PA, Brazil

10 ᵇ UMR7590 Institute for Mineralogy, Physics of Materials and Cosmochemistry,

11 Museum National d'Histoire Naturelle, Sorbonne Universite, 75005 Paris, France

12 ᶜ Research Productivity of CNPq, Brazil

13 ᵈ Department of Earth Sciences, Carleton University, Ottawa ON, Canada K1S 5B6

14 ᵉ Universidade Federal de Minas Gerais, Programa de Pós-Graduação em Geologia,

15 CPMTC-IGC, Campus Pampulha, Belo Horizonte, MG, Brazil

16 ᵈ Department of Geology Sedimentar e Ambiental, Instituto de Geociências,

17 Universidade de São Paulo, São Paulo, Brazil

18 CNRS - UMR6538 Laboratoire Geo-Ocean, European Institute for Marine Studies,

19 Université de Bretagne Occidentale, 29280, Plouzané, France

20
21 ABSTRACT

22 The extensive deglaciation linked to snowball Earth climate elapsed in terminal 23 Neoproterozoic caused dramatic changes in the ocean's chemistry, exceptionally recorded 24 in the cap carbonate sequence. One of the most spectacular examples is the post-Marinoan 25 Puga cap carbonate (~635-622 Ma), exposed in the Amazon Craton associated with 26 expressive sea level changes related to the glacial-isostatic adjustment (GIA) resulting in 27 continuous mixed-waters, coastal upwelling succeeded by long-term transgression under 28 greenhouse climax. To better understand the complex dynamics during this period, we 29 conducted a comprehensive multiproxy study encompassing sedimentological, rare earth 30 elements, and stable isotopic analysis. The Rare Earth Element + Yttrium (REY) patterns 31 in the Puga cap carbonate reflect the marine conditions aftermath of the aftermath of the 32 snowball Earth. Specifically, the Y/Ho ratios in the cap dolostones consistently fall below 33 modern seawater values interpreted as seawater composition dilution due to the meltwater 34 influx from the continents during deglaciation. However, superchondritic Y/Ho ratios (up 35 to 70) and positive Eu anomalies (Eu/Eu* up to 3.6) are observed in the basal portion of 36 the cap dolostone. These results suggest the upwelling of hypersaline seawater with some 37 contribution from contemporaneous hydrothermal fluids in shallow coastal water. These 38 findings are interpreted as evidence of ocean destratification related to upwelling, ocean 39 circulation, and glacioisostatic adjustment. The cap dolostone occurs with minimal 40 siliciclastic contributions, consistent with a siliciclastic starvation model. It could be 41 attributed to the persistent shallow-water conditions in the coastal zone, significantly 42 influenced by glacial isostatic adjustment (GIA). This process led to continuous uplift, 43 resulting in a shallow shelf isolated from direct oceanic processes. Barrier-inland facies 44 were not observed in the studied succession, suggesting that irregular diamicton 45 morphologies may have created protected areas against waves. This scenario maintained 46 a consistently shallow water column and followed a significant landward shoreline 47 migration depositional pattern. These dramatic changes have been assumed to have 48 resulted from the rapid recovery of primary productivity during deglaciation, with the 49 proliferation of microbial communities and induced dolomite precipitation. The δ¹³C 50 isotopic signature is consistent with these interpretations and with the values found for 51 the Ediacaran in other cap carbonates. Mg concentration was constant during restricted 52 paleoenvironmental conditions on dolomitic platforms. The rapid rise in sea level led to 53 the dispersion of Mg ions, leading to the definitive interruption of dolomicrite 54 precipitation, abruptly replaced by seas supersaturated in CaCO₃ with massive aragonite 55 precipitation coinciding with the increase in detrital components under the influence of a 56 long-term transgression in the initial Ediacaran.

57 Keywords: cap dolostone; starvation siliciclastic, plumeword; restriction; upwelling

58 1 INTRODUCTION

59

60 The transition from the extreme Snowball Earth climate state to more temperate 61 conditions represents a pivotal interval (~720-635 Ma) of Earth's geologic history. These 62 drastic changes were manifested in marine sediments through the dramatic transformation 63 from a globally frozen ocean to an ice-free one (Hoffman et al. 2017, 2021). Significant
perturbations triggered large changes to sea level during the pre-, syn-, and post-glacial stages, primarily driven by the effects of glacial eustasy (global mean sea level), glacial-isostatic adjustment (GIA), and the thermal expansion of the oceans under greenhouse conditions (Allen and Hoffman, 2005; Benn et al. 2015; Creveling and Mitrovica, 2014; Hoffman et al. 2017, 2021, Mitchell et al. 2021, Rame and Marotzke, 2022). During the advance of the syn-deglacial transgression, meltwater production and surface warming created stable density stratification (Yang et al. 2017). Salinity below the meltwater layer, which encompassed the majority of the global ocean at this time, has been estimated to be three times higher than the modern ocean, creating a stable stratification that may have lasted thousands to hundreds of thousands of years (Shield, 2005; Liu et al. 2014; 2018; Yang et al. 2017; Hoffman et al. 2017). The GIA and oceanic circulation were crucial factors for the end of stratification during the post-melting stage (Yang et al., 2017; Rame and Marotzke, 2022).

Globally deposited cap carbonate sequences record a critical interval documenting the transition out of the Marinoan Snowball Earth (Hoffman et al. 2017). These sequences are recognized as the Cryogenian-Ediacaran boundary and consist of basal Ediacaran cap dolostones, which are conformably overlain by cap limestones. These associations exhibit striking similarities across continents and have been globally correlated through geochronological and geochemical means (e.g., Halverson et al., 2005; Corsetti and Lorentz, 2006; Hoffman et al., 2017; Yu et al., 2022).

Current cap carbonate formation models have been proposed to explain the genesis of the Marinoan cap dolostones. In one model, it is suggested that increased seawater alkalinity played a crucial role in the precipitation and widespread dolomitization observed during cap dolostone deposition. These factors include the overturning of deep anoxic ocean waters (Grotzinger and Knoll, 1995; Hoffman et al.,
1998; James et al., 2001), intense chemical weathering (Hoffman et al. 1998; Higgins and Schrag, 2003), Hydrothermal alkalinity model: alkalinity increases linked to extensive hydrothermal activity along basin-bounding faults (Gernon et al., 2016) gas-hydrate destabilization and massive methane release (Jiang et al. 2003, Jiang et al. 2006; Kennedy et al. 2008; Wang et al. 2008), alkanilinity increases associated to microbial activity (e.g., sulfate reduction) (Nogueira et al. 2003, 2007, 2022; Font et al. 2006) and meltwater plumes (Shields, 2005; Yang et al. 2017; Liu et al. 2014; Santos et al. 2021).

Despite the widespread distribution and intense study of these post-Marinoan sequences and the various proposed models, there has yet to be a consensus regarding their genesis. One aspect that remains to be understood is the microfacies changes from dolostone to limestone. This texture and factory composition transition is key to unraveling the intricate processes during the initial Ediacaran. The precipitation mechanism of cap dolostones remains a subject of intense debate. One reason for these ongoing debates is the variability between units; for example, many units record enigmatic barite horizons (e.g., Bao et al. 2008; Crockford et al. 2016; 2018), while many others do not. Despite this variability, to date, two main hypotheses have been proposed. The first suggests that dolostones formed directly, retaining their primary characteristics and reflecting the chemistry of the original depositional environment (Kennedy, 1996; Shields, 2005; Nogueira et al., 2003; Nedeléc et al., 2007; Font et al., 2006; Santos et al. 2021, 2023). On the other hand, the second hypothesis proposes that the dolostones underwent secondary dolomitization, wherein an aragonite or calcite precursor was transformed into dolomite (Fabre et al. 2013; Higgins et al. 2018; Ahm et al. 2019; Gammon et al. 2012; Gammon, 2012).

Rare earth elements and yttrium (REE+Y) have been extensively utilized in paleoenvironmental and paleoceanographic studies to provide insights into the aftermath
of Snowball Earth events (Font et al. 2006; Nedelec et al. 2007, Huang et al. 2009; Meyer et al. 2012; Wang et al. 2014; Viera et al. 2015; Hohl et al. 2015, 2017; Caxito et al. 2018; Verdel et al. 2018; Wu et al. 2019; Zhao et al. 2021; Cui et al. 2021; Ren and Li, 2023).
These elements, when analyzed in ancient non-skeletal carbonates such as microbialites, serve as proxies for reconstructing paleoenvironmental conditions and offer valuable information about past seawater conditions (Schere and Suitz, 1980; Shaw and Wasserburg, 1985; Webb and Kamber, 2000; Kamber and Webb, 2001; Nothdurft et al. 2004; Huang et al. 2009; Ling et al. 2013). The incorporation of REE+Y into carbonate minerals is controlled by various physicochemical processes, such as water chemistry, pH, and temperature (Wood, 1990; Nozaki and Alibo, 2003; Chen et al., 2015; Bolhar et al., 2015; Rahlf et al. 2021; Zhao et al. 2022). Consequently, the utility of marine carbonates as recorders of oceanographic conditions hinges upon the preservation quality of these carbonates, well-established diagenetic constraints, and a comprehensive understanding of contamination by noncarbonate phases.

Abundant carbonate deposits with well-preserved primary textures occur in the Southern Amazon Craton associated with the Marinoan Puga cap carbonate. Previous studies have predominantly focused on microbial carbonate petrology, sedimentology, and δ¹³C and δ¹⁸O geochemical values of carbonates (Nogueira et al. 2003; 2022; Font et al. 2006; 2010; Sansjofre et al. 2011; Soares et al. 2020; Romero et al. 2020; Santos et al. 2021, 2023). However, despite well-established geological constraints, this succession lacks a comprehensive multiproxy study to constrain the post-Marinoan conditions and test previous carbonate precipitation models, thereby refining our understanding of the factors driving the changes in the carbonate factory over this critical time interval. The evaluation of the Puga cap carbonate exposed in the southern part of the Amazon Craton represents a unique opportunity to achieve this end due to its exceptional preservation,
largely unaffected by severe late-stage diagenetic modification (Sansjofre et al. 2014; Soares et al. 2020; Santos et al. 2021). Therefore, a comprehensive multiproxy study incorporating sedimentological aspects (microfacies and facies), rare earth elements (e.g., Y/Ho, Eu/Eu* and Ce/Ce* anomalies), and isotopic composition of δ¹³C and δ¹⁸O may provide the needed insights into the post-Marinoan conditions in this region. For example, the impact of meltwater and stable stratification on ocean circulation in this region and the potential influence of siliciclastic starvation on precipitation and carbonate factory changes (from dolostone to aragonite) remain to be evaluated.

147

## 2. GEOLOGICAL BACKGROUND

149

The focus of this study is the Puga cap carbonate, which is in the Southern Amazon Craton and is associated with the basal deposits of the intracratonic Araras-Alto Paraguai Basin installed during the Late Cryogenian (Fig. 1 A and B). (Nogueira et al. 2003; 2019; Gaia et al. 2017; Afonso and Nogueira, 2018; Santos et al. 2020; Soares et al. 2020; Romero et al. 2020; Santos et al. 2021,2023). This basin overlies crystalline and metamorphic basement rocks of the Cuiabá Group that characterize the Paraguai Belt (Nogueira et al. 2019; Santos et al. 2020). The basal fill of this basin is composed of Cryogenian rocks of the Puga Formation (pelites, sandstones, dropstones, and glacial diamictites) overlain by earliest Ediacaran carbonate rocks of the Araras Group (Fig. 1C), which consists of four formations: 1) Mirassol d'Oeste (Cap dolostone), composed of dolomitic layers interpreted as a shallow platform environment; 2) Guia (Cap Limestone), composed of carbonates and shales, interpreted as a moderately deep platform; 3) Serra do Quilombo, composed of shallow to moderately deep dolomitic platforms; and 4) Nobres, consisting of tidally influenced dolomites (Nogueira and Riccomini, 2006;
164 Nogueira et al. 2003; 2007; 2019;2022; Rudnitzki et al. 2016; Romero et al. 2020; 2022;
165 Soares et al. 2013; 2020; Santos et al. 2021).

166 The cap dolostone comprises stratiform stromatolites with gypsum
167 pseudomorphs, giant stromatolites with tubestone, and wave-dominated dolostone facies
168 (peloidal dolomudstone with parallel laminations; peloidal dolopacktone with quasi-
169 planar laminations, and peloidal dolopacktone with megaripples beds). The cap limestone
170 is predominantly composed of limestones with megaripples, wavy laminations, quasi-
171 planar and crystal fans, and subordinate laminated shale (for more details, see Nogueira
172 et al. 2003; 2018; Sansjofre et al. 2014; Soares et al. 2020; Romero et al. 2020; Santos et
173 al. 2021,2023). Post-Marinoan cap carbonate deposition is generally constrained through
174 U-Pb zircon dating and more recent Re-Os dating to between 635.2±0.5 Ma to 632.3±5.9
175 Ma respectively (Condon et al., 2005; Rooney et al., 2015). Consistent with this are ages
176 of 627±32 Ma for cap dolostone and 622±33 Ma for the cap limestone obtained through
177 Pb-Pb dating, indicating a lower Ediacaran age (Babinski et al. 2006; Romero et al. 2013)
178 and an age of 623 ± 3 [9] Ma U-Pb in situ in seafloor crystal fans was obtained recently
179 by Carvalho et al. (2023). This latter age is interpreted as a new constraint on the transition
180 from shallow to intermediate burial diagenesis of the Puga cap carbonate (Fig. 1C). The
181 isotopic values of δ¹³C from -5 ‰ on the hood to -2 ‰ and ⁸⁷Sr/⁸⁶Sr between 0.7074 and
182 0.7089 (Fig. 1C) are compatible with global curves (e.g., Otavi Group, Namibia)
183 (Nogueira et al. 2019).

## 3. MATERIALS AND METHODS

184 In total, 97 samples were analyzed from the cap dolostone (Mirassol d'Oeste
185 Formation) and cap limestone (Guia Formation) of the Puga cap carbonate, southern
186 Amazon craton (Fig.1). Of these 97 samples, 37 were sampled via outcrop from the
187 Terconi section in the Terconi quarry and 60 were sampled from the Tangará section with
188
189 41 samples being taken from drill cores and 19 samples from outcrops (Fig.3 and Fig.4).

190 The samples followed the stratigraphic proposed and facies/microfacies criterion
191 presented by Nogueira et al. 2003; Sansjofre et al. 2014 and Santos et al. 2021,2023.

192 3.1Trace elements

193 **3.1.1 REE geochemistry - acetic acid leaching**

194 In total, 42 samples of the drillcore JOR, which captures the cap dolostone from
195 the Tangará section, were prepared for major/minor/trace/REE + Y concentrations to
196 avoid REE contributions from non-carbonatic components (e.g., clays and oxides)
197 following the digestion protocol of Rongemaille et al. (2011). Approximately 20 mg of
198 powder was weighed into acid-cleaned 2.5 mL microcentrifuge tubes, and 0.5 mL of
199 trace-metal grade acetic acid at a concentration of 0.85 N (i.e., 5%) was added and left to
200 dissolve the samples overnight at room temperature. Samples were centrifuged, and 0.25
201 mL aliquots of the supernatant were diluted in 4.75 mL PFA-distilled nitric acid prepared
202 to a concentration of 0.45 N HNO3 (i.e.,2%) with 10 ppb indium as an internal standard
203 to control instrumental drift. Analyses were conducted on a Thermo Scientific
204 ELEMENT XR high-resolution inductively coupled plasma mass spectrometer (HR ICP-
205 MS) of the Pôle Spectrométrie Océan (PSO) at Institut Universitaire Europe'en de la Mer,
206 Université de Bretagne Occidentale, France. Using low, medium, and high-resolution
207 modes allowed for measuring major/minor/trace/REE + Y concentrations at a constant
208 dilution factor during a single analytical session. Repeat analyses of the standard
209 reference material CAL-S (N = 15) showed 1r precision of 7-9%, with results within 10%
210 of published values. Concentrations were determined by external calibration with a
211 synthetic, matrix-matched calibration solution.
212

### 3.1.2 REE geochemistry - 2 N nitric acid leaching

213

214

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216

217

218

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220

221

222

223

224

225

In addition, 56 outcrop samples were analyzed, leaching 37 from the Terconi section and 19 from the Tangará section using total bulk dissolution. Rock fragments free of apparent fractures or recrystallization were carefully selected and crushed in an agate mortar. Chemical analysis was performed by ICP-AES (major and minor elements) and ICP-MS (trace elements) at the spectrochemical laboratory (SARM) of the Centre de Recherches en Pétrographie et Géochimie in Vandœuvre-les-Nancy. Samples were prepared by fusion with LiBO₂ (980 °C during 60 min), followed by dissolution in an HNO₃ (1 mol L⁻¹)-H₂O₂ (~0.5% v/v)-glycerol (~10% v/v) mixture to obtain a diluted solution (Carignan et al. 2001). Precision and accuracy were better than 1% of the measured value (mean 0.5%) for major-minor elements and 8% for trace metals, as checked by international standards and analysis of replicate samples (Carignan et al. 2001).

226

227

## 3.2 Stable isotopes (δ¹³C and δ¹⁸O)

228

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232

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234

235

236

In total, 44 samples of the drill core JOR, which captures the cap dolostone from the Tangará section, were prepared for stable isotope analyses and were performed at the National Museum of Natural History in Paris at the SMIMM Platform. The samples were first ground into a powder in an agate mortar. CO₂ was extracted by dissolution with 100% H₃PO₄ at 70 °C under a primary vacuum. Stable carbon and oxygen isotopic compositions of the product CO₂ were measured using a Kiel IV device. Standards CA21 – Marble – NBS19 were calibrated to the δ¹³C sample/ref data. These standards have been calibrated relative to V-PDB using two NBS19 and other international standards. The
results are presented in δ-notation relative to the international Standards V-PDB for δ¹³C and δ¹⁸O. The external reproducibility for δ¹³C and δ¹⁸O measurements is 0.1‰ and 0.15‰, respectively (1σ).

240

## 4. RATIONALE FOR THE USE OF RARE EARTH ELEMENTS + YTRIUM (REE+Y) AS PALEOCEANOGRAPHIC INDICATORS

243

Ancient nonskeletal carbonates, including microbialites, can record ambient seawater rare earth elements and yttrium (REE+Y) concentrations, making them excellent potential archives for paleoenvironment proxies (Schere and Suitz, 1980; Shaw and Wasserburg, 1985; Webb and Kamber, 2000; Kamber and Webb 2001; Nothdurft et al. 2004; Huang et al. 2009; Ling et al. 2013; Hodskiss et al., 2021). However, their utility is predicated on the ability to control for noncarbonate contamination, as well as the impacts of diagenesis (Bolhar and Van Kranendonk, 2007; Rongemaille et al., 2011; Zang et al. 2015; Van Kranendonk et al. 2003; Bolhar et al. 2004; Nothdurft et al. 2004; Olivier and Boyet, 2006; Bolhar et al. 2015). In this study, we primarily utilize Y/Ho ratios, Ce anomalies, and Eu anomalies to explore the post-Marinoan record of the Amazon craton.

The Y/Ho ratio is a geochemical parameter distinguishing between continental and marine influence in marine sediments (Bau et al. 1996; Nozaki et al. 1997; Wang et al. 2014). Yttrium (Y) is more abundant in continental rocks, released during weathering, and transported by rivers. Holmium (Ho) is associated with marine waters (Lawrence et al. 2006). Consequently, the Y/Ho ratio of estuaries shows an increase from crustal values of ~27 to typical marine high values of >40, together with the increasing salinity towards the open ocean (Lawrence and Kamber., 2006). REE + Y is relatively flat in freshwaters with typical Y/Ho values ≈ 26 (Taylor and McLennan, 1985; Kamber et al., 2005;
Lawrence et al., 2006; Bolhar and Van Kranendonk, 2007; Babechuk et al., 2015). Values between ≈ 36 and 26 typically relate to inland waters or mixed seawater-freshwater systems (Webb and Kamber, 2000; Van Kranendonk et al., 2003; Bolhar et al., 2004; Shields and Webb, 2004; Lawrence). Similarly, contamination by detrital components typically overestimates carbonate concentrations of elements such as Al, Ti, Th, and Zr and will reduce Y/Ho ratios to values of ≈27, which is similar to the continental crust and much lower than expected seawater Y/Ho (Webb and Kamber, 2000; Nothdurft et al. 2004; Uhlein et al. 2021).

Ce anomalies are a valuable geochemical tool to assess past redox conditions in sedimentary environments. Cerium (Ce) exists in two oxidation states: Ce (III) and Ce (IV). Ce (III) is preferentially incorporated into sediments in reducing conditions, leading to a higher Ce anomaly, Ce/Ce* >1. Conversely, Ce (IV) is more stable and retained in the sediments under oxic conditions, resulting in a lower Ce anomaly, Ce/Ce* < 1. Ce anomalies in sedimentary records have been widely used to reconstruct marine redox conditions over time (Kamber and Webb, 2001; Bolhar et al., 2004; Bau and Alexander, 2006; Bolhar and van Kranendonk, 2007; Pourret et al., 2008; Planavsky et al. 2010; Ling et al. 2013; Hood and Wallace, 2015; Rodler et al. 2016; Sahoo et al. 2016; Wallace et al. 2017; Zhang and Shield, 2022).

The Eu/Eu* anomaly is a useful indicator for investigating changes in redox and hydrothermal-related parameters (Bau, 1991; Klinkhammer et al., 1994; Barrat et al., 2000; Lawrence et al., 2006; Alexander et al., 2008; Parsapoor et al. 2009). Under reducing conditions, such as hydrothermal systems influence of high-temperature fluids associated with fumaroles, Eu (II) is preferentially incorporated into sediments, resulting in higher Eu/Eu* anomalies. In contrast, Eu (III), under oxidizing conditions, led to lower Eu/Eu* anomalies. High Eu anomalies are absent from the general characteristics of
287 marine carbonates. The Eu/Eu* ratio in sedimentary records allows us to infer past
288 variations in hydrothermal activity or the presence of Eu in detrital sediments, especially
289 within calcic plagioclase. This insight significantly enhances our understanding of its
290 influence on sediment geochemistry and the resulting shifts in redox conditions over time.
291 (Michard, 1989; Klinkhammer et al., 1994; Haas et al., 1995; Bau and Dulski, 1999;
292 Douville et al., 1999; Barrat et al., 2000; Dai et al., 2016; Liu et al., 2016 Cui et al. 2019).

293 **5. RESULTS**

294 5.1 Major and Trace Elements

295 Trace element data from the Calcário Tangará and Terconi Quarry sections (Fig.
296 2 and Fig. 3) are listed in Table 1. They are described below, following the stratigraphic
297 proposed by Nogueira et al. 2003 and Santos et al. 2021,2023. The cap dolostone
298 (Mirassol d'Oeste Formation) was subdivided from base to top into four microfacies:
299 stratiform doloboundstones with gypsum pseudomorph, domic doloboundstones with
tubestone, and wave-dominated dolostone facies (peloidal dolograinstone/dolomudstone
300
301 with quasi-planar laminations and low-angle truncations, and peloidal dolograinstone
302 with megaripples beds) (see Soares et al. 2020, Santos et al. 2021,2023 for more details).

303 The measured REE+Y data of all Puga cap carbonate samples were normalized
304 to those of post-Archean Australian shale (PAAS) by McLennan, 1989. Anomalies of La,
305 Ce, and Eu were calculated using the equations following (1) Lawrence et al. (2006) and
306 (2) Frimel 2009. Overall, REE+Yttrium (ΣREE+Y) patterns from the Puga cap carbonate
307 are similar in both sections (Fig. 4 and Fig. 5). The Puga cap carbonate shows a
308 progressive enrichment of REEs up section, which is lower than those in the PAAS by
309 less than two orders of magnitude, in the Tangará SectionΣREE+Y: 1.58–60.41 ppm,
310 average = 5.79ppm, n = 61; Fig. 4) and the Terconi Section (ΣREE+Y: 1.4–97.6 ppm,
311 average = 11.4. ppm, n = 37; Fig. 5).
312 **5.1.1. Calcário Tangará Quarry**

313 **5.1.1.1 Cap dolostone from Tangará Quarry**

314 **5.1.1.1.1 Microbialites - Tangará Quarry**

315

316 In the Calcário Tangará Quarry section (Fig. 2), the microbialites
317 (doloboundstones stratiform stromatolites with gypsum pseudomorph cement and
318 doloboundstones domic stromatolites with tubestone associated) of the cap dolostone
319 showed Mg/Ca ratios ranging from 0.6 to 1.0 (average = 0.8 ± 0.18; n = 42), and low
320 concentration of Sr from 22. to 87 ppm (average = 50 ± 14 ppm) (Fig. 2). The stratiform
321 doloboundstones with gypsum pseudomorphs cement have low concentrations of
322 ΣREE+Y ranging from 1.6 to 27.04 ppm (average = 4 ± 4.7 ppm; n = 32), showed a flat
323 (REE + Y)SN with a depleted LREE pattern (average (Pr/Yb)SN = 1.0 ± 0.3; n= 33) (Fig.
324 6A, C and E). Negative La anomalies, La/La* = 0.8 to 2.2, average = 0.9 ± 0.3). Ce
325 anomalies are mostly weakly negative, with Ce/Ce* = 0.4 to 1.0 (average = 0.9 ± 0.2),
326 and 3) strong positive Eu anomalies, Eu/Eu* =1.0 to 3.2 (average = 1.5 ± 0.5). These
327 facies record slightly super-chondritic Y/Ho ratios (average = 32 ± 10; n = 32), with some
328 sample reaching Y Y/Ho values up to 70 (super-chondritic) (Fig. 7). The Y/Ho ratio of
329 these facies showed an absence of correlation factors (R² = 0.023); Th (R² = 0.0017); Ni
330 (R²= 0.0024), and Pb (R² = 0.063).

331 The domic doloboundstones with tubestone (giant stromatolites) exhibit ΣREE
332 +Y concentrations between 3.0 to 34.8 ppm (average = 3.7 ± 0.97; n=10) and have a flat
333 (REE + Y)SN (Fig. 4 B) with a depleted LREE pattern (average (Pr/Yb)SN = 0.7 ± 0.97;
334 n =10) (Fig. 6A, C, and E), occurring with 1) Weakly negative Ce anomalies, Ce/Ce*
335 =0.8 to 1.1, (average = 0.9 ± 0.1); 2) Positive Eu anomalies, Eu/Eu*= 1.0 to 1.9, (average
336 =1.3 ± 0.3); and 3) negative La anomalies, La/La*= 0.8 to 1, (average = 0.9 ± 0.1). The
337 Y/Ho ratio in this location was found to be between 27 and 30 (average = 29 ± 1) (Fig. 7)
338 and shows an absence of correlation factor with Zr (R2 = 0.2); Th (R2= 0.3); Ni (R2=
339 0.24), and Pb (R2= 0.30).

340 5.1.1.1.2 Wave-dominated dolostone facies Tangará Quarry

341

342 The wave-dominated dolostone facies (peloidal dolomudstone with parallel
343 lamination, peloidal dolopacktone with quasi-planar lamination, and peloidal
344 dolopacktone with megaripples beds) show Mg/Ca ratios ranging from 0.6 to 0.9,
345 (average = 0.6 ± 0.11; n=9) and with low Sr concentrations between 62.75 and 173 ppm
346 (average= 86.4 ± 34 ppm) are similar to those of the basal section (Fig. 2). the
347 concentrations of the detrital component towards the top of the cap dolostone, with Al =
348 12.34 to 37.88 ppm (average = 25 ± 9.7), Zr =7 to 20 ppm (average =13 ± 5 ppm) and Th
349 = 0.59 to 1.77 (average = 1.18 ± 0.5 ppm) (Fig. 2). The transition from cap dolostone to
350 cap limestone record the decreased purity toward the top of the Puga cap carbonate. The
351 wave-dominated dolostone facies exhibit ΣREE +Y concentrations between 12.3 to 39.02
352 ppm (average = 26 ppm ± 8.3; n =10), have a flat (REE + Y)Sn (Fig. 4C), with a depleted
353 LREE pattern (average (Pr/Yb) Sn = 1.0 ± 0.1) (Fig. 6A, C, and E)), and show the
354 following attributes: 1) Absence of La anomalies La/La* =0.8 to 0.9 (average = 0.84 ±
355 0.03). 2) weakly negative Ce anomalies Ce/Ce* =0,8 (average = 0.78 ± 0.02), 2). Eu
356 anomalies are small to absent, Eu/Eu* = 0,9 to 1,0 (average=0.96 ± 0.04). Y/Ho = 32 to
357 34 (average = 33.5 ± 0.6) (Fig. 2) and absence of correlation with Zr (R²= 0.085); Th
358 (R²=0.16); Ni (R²=0.096); Pb (R²=0.39).

359 5.1.1.2 Cap Limestone Tangará Quarry

360 The cap limestone shows higher Sr concentrations between 549 to 1115 ppm
361 (average = 634 ± 196; n =10). The cap limestone is also associated with an elevated
presence of detrital components compared to the cap dolostone, such as Al = 17.15 to 45.42 ppm (average = 34.3 ± 9.3), Zr = 18 to 48 ppm (average = 28 ± 10), and Th = 1.14 to 2.68 ppm (average = 1.86 ± 0.48) (Fig. 2). Cap limestone has (REE + Y) SN patterns (Fig. 4 D) characterized by a higher concentration of ΣREE +Y, ranging from 34.9 to 60.4 ppm (average = 47.2 ± 8.1; n = 10) with a depleted LREE pattern (average (Pr/Yb) SN = 0.87 ± 0.07) (Fig. 6A, C, and E), and show the following attributes: 1) negative La anomalies, La/La* = 0.8 to 0.9 (average = 0.79 ± 0.06); 2) weakly negative Ce anomalies Ce/Ce* = 0.7 to 0.8, (average = 0.76 ± 0.02); 3) Eu anomalies are minor to absent, Eu/Eu* = 0.9 to 1.0, (average = 1.04 ± 0.05). The Y/Ho ratio exhibits a super-chondritic value when contrasted with the cap dolostone between 33 to 38 (average = 35 ± 1.6) (Fig. 7) and occurs with an absence of Zr (R²= 0.069); Th (R²= 0.35); Pb (R²= 0.10), and Ni (R²= 0.13).

## 5.1.2. Terconi Quarry

### 5.1.2.1 Cap dolostone from Terconi Quarry

#### 5.1.2.1.1 Microbialites Terconi

The microbialites (stratiform doloboundstones with gypsum cement and domic doloboundstones with tubestone) of the cap dolostone in the Terconi quarry section are similar to those of the Calcário Tangará Section, as follows: 1) Mg/Ca ratio ranging from 0.4 to 0.6, (average= 0.6 ± 0.05; n = 18); 2) low Sr concentrations of between 40 to 60 ppm (average = 49 ppm ± 6; n = 17); 3) Zr = 1.5 to 12.6 ppm (average = 3.5 ± 3.0); 4) Th = 0.03 to 0.75 (average = 0.19 ± 0.17); and 5) Al = 1.69 to 42.30 ppm (average= 7.9 ± 9.7) (Fig. 3). The microbialites at the basal section occurs with low ΣREE +Y concentrations, ranging from 1.4 to 6.6 ppm (average = 5.80 ppm ± 3.4; n = 18) and have
flat (REE + Y)SN patterns (Fig. 5A) with a depleted LREE pattern (average (Pr/Yb)SN = 0.82 ± 0.15; n=17) (Fig. 6B, D, and F), and show the following attributes:1) weakly negative Ce anomalies Ce/Ce* = 0.7 to 1.1 (average = 0.93 ± 10), 2) Eu anomalies are small to absent Eu/Eu* = 0.9 to 1.2 (average = 1.08 ± 0.04), and 3) weakly negative La anomalies, La/La* = 0.7 to 1.0, (average = 0.94 ± 0.09). The Y/Ho ratio exhibits a super-chondritic value when contrasted with the medium values of the microbialites from Tangará da Serra, ratios between 32 to 49 ppm (average = 35.2 ± 0.09) (Fig. 3). These ratios showed an absence of correlation with (R² = 0.085); Th (R² = 0.16); Ni (R² = 0.096), and Pb (R² = 0.39).

### 5.1.2.1.2 Wave-dominated dolostone facies Terconi

The wave-dominated dolostone facies (e.g., peloidal dolopacktone with megaripples beds) has Mg/Ca = 0.5 to 0.6 (average = 0.6 ± 0.01 n= 8) and low Sr concentrations similar at the base of the section, between 63 and 87 ppm (average = 59 ± 8.4 ppm). In these facies, there is an increase in detrital components compared to the basal microbialites, such as Al = 10.59 to 83.65 ppm (average = 36.6 ± 24.1); Zr = 3.14 to 27.13 ppm (average = 12.6 ± 9.25), and Th = 0.24 to 1.54 ppm (average = 0.92 ± 0.55) (Fig. 3).

These facies occur with higher ΣREE +Y concentrations than those of microbialites, ranging from 8.32 to 42.26 ppm (average = 15.6 ppm ± 12.5; n = 8) and have flat (REE + Y)SN patterns (Fig. 5B) with a depleted LREE pattern (average (Pr/Yb)SN = 0.9 ± 0.1; n = 8) (Fig. 6B, D, and F), and show the following attributes:1) weakly negative Ce anomalies, Ce/Ce* = 0.8 to 0.9, (average = 0.87 ± 0.02), 2) Eu anomalies are minor to absent, Eu/Eu* = 0.9 to 1.1 (average = 0.95 ± 0.06), and 3) weakly negative La anomalies, La/La* = 0.7 to 1.1, (average = 0.96 ± 0.05) and slightly super-chondritic
410 Y/Ho ratios between 32 to 34 (average = 34.2 ± 1.6) (Fig. 8B) and showed an absence of
411 a correlation between Zr (R² = 0.14), Th (R² = 0.03), Ni (R² = 0.0002) and Pb (R² = 0.074).

412 5.1.2.2 Cap Limestone Terconi Quarry

413

414 The cap limestone shows higher Sr concentrations between 549 to 1115 ppm
415 (average = 427 ± 347 ppm; n = 12). The cap limestone is also associated with an elevated
416 presence of detrital components compared to the cap dolostone, such as Al = 10.59 to
417 83.65 ppm (average = 71.5 ± 39 ppm), Zr = 63.13 to 194.71 ppm (average = 35.14 ± 17.4
418 ppm) and Th = 1.5 to 3.8 ppm (average = 1.8 ± 0.7 ppm) (Fig. 3). Cap limestone has (REE
419 + Y) SN patterns (Fig. 5C) characterized by a significantly higher concentration of ΣREE
420 +Y, ranging from 36,4 to 97,6 ppm (average = 46.3 ± 17 ppm, n =12) with a depleted
421 LREE pattern (average (Pr/Yb) SN = 0.9 ± 0.05, n=12) (Fig. 6B, D, and F), and show the
422 following attributes;1) weakly negative Ce anomalies, Ce/Ce* = 0.9 (average = 0.9 ±
423 0.01); 2) Eu anomalies are small to absent, Eu/Eu* = 0.9 to 1.0 (average = 0.964 ± 0.03),
424 and 3) weakly negative La anomalies La/La* = 0.8 to 1.1 (average = 0.86 ± 0.07). Y/Ho
425 ratios range from 31 to 37 (average = 33 ±2) and show a low correlation between Zr (R²
426 = 0.29); Th R² = 0.47, Ni (R² = 0.34) and Pb (R² = 8.10⁻⁶).

427 5.2 Carbon and oxygen isotope stratigraphy in the Tangará Quarry cap dolostone

428

429 The δ¹³C and δ¹⁸O results of the Calcário Terconi Quarry section are presented in
430 Fig. 8 and detailed in Table 2. The isotopic composition is relatively homogeneous. In
431 the dolostone basal (stratiform doloboundstones with gypsum pseudomorphs), δ¹³C
432 values from -4.1 to -8.2 (average = -4.8% ± 1.1 n = 33) and δ¹⁸O values from -4.1% to -
433 7.3% (average = -5.1 % ± 0.7 n = 33) occur, followed by δ¹³C values from -4.8‰ to
434 -5.7‰ (average = -5.0‰ ± 0.3 n = 10) and δ¹⁸O values from -3.7% to -4.9% (average
435 = -4.7‰ ± 0.4 n = 10) in the domestic doloboundstones with tubestone associated. Our data
436 complete the isotopic curve of the dolostone succession (Fig. 7) made by Sansjofre et al.
437 2011 and Nogueira et al. 2019 show an absence of covariation between δ¹³C and δ¹⁸O
438 (R²=6.10⁻⁵) isotope compositions. The isotopic trends are similar to those obtained for
439 Puga cap carbonates from the southern Amazon craton (Terconi, Fig. 7B) and Calcário
440 Tangará Quarry (Fig. 7A) by Nogueira et al. 2003;2019; Font et al. 2006;2010; Sansjofre
441 et al. 2011, which resemble similar trends for other Ediacaran carbonate platforms.

442 6. DISCUSSION

443 6.1 Assessment of detrital input to Puga cap carbonate deposits

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Previous microfacies studies have suggested that the Puga cap carbonate sequence
within the southern Amazon Craton occurs with low detritic content in the matrix,
especially the cap dolostone (Nogueira et al. 2003; 2019; Font et al. 2004; Sansjofre et al.
2014; Soares et al. 2020; Santos et al. 2021). To assess the impact of detrital contributions
on cap dolostone geochemistry (REE+Y anomalies), we evaluated the data using two
dissolution techniques: total bulk rock dissolution and an acetic acid leach. The cap
dolostone exhibited high purity in both sections, Tangará Quarry (Fig. 2) and Terconi
Quarry (Fig. 3); the results are presented below. Our data suggest a clear covariance
between the increase in insoluble particles (Zr, Al, Th abundances) vs. ΣREE+Y. Zr
concentrations are an excellent marker of siliciclastic contamination, as they are abundant
in different detrital materials, and contamination will drive carbonate [Zr] far above
typical values of ≈4 ppm (Frimmel, 2009). Similarly, [Th] > 1 ppm may indicate
significant contributions from insoluble residues and thus are considered to be another
457 marker of detrital contamination (Webb and Kamber, 2000; Zhao and Zheng, 2014,
458 2017).

459 In this study, binary diagrams were generated to determine the correlation between
460 different element ratios and element concentrations to assess the degree of diagenetic
461 alteration and contamination by siliciclastic materials on ΣREE+Y distribution patterns
462 (Kamber and Webb, 2001; Nothdurft et al. 2004; Bolhar and Van Kranendonk, 2007;
463 Allwood et al. 2010; Wang et al. 2014; Zhao and Zheng, 2014, and Uhlein et al. 2021).
464 Despite the apparent influence on ΣREE+Y, there is a lack of anti-correlation, for
465 example, between Y/Ho ratios, which would be a predicted behavior in carbonates
466 significantly contaminated by terrigenous particles. The increase in the insoluble content
467 of the cap dolostone is observed over the transition from microbialite facies to wave-
468 dominated dolostone (Figs. 2 and 3). Conservative mixing lines between end members
469 are based on Y/Ho vs. (Zr, Th, Al, and ΣREE + Y) (Fig. 8) data using calcite cement from
470 Nothdurft et al. (2004) and PAAS from Taylor and McLennan (1985). The calcite cement
471 by Nothdurft et al. 2004 followed the criterion of purity and not its early diagenetic
472 history.

473 When assessing the impact of detrital materials on total bulk rock dissolution
474 analyses, we find that the cap dolostone is contaminated by 1%-5% detrital particles (Fig.
475 8A and 8B) (Fig. 8E and 8F). However, Y/Ho vs. Al suggests up to <1% contamination
476 (Fig. 8C and 8D), while Y/Ho vs. Th indicates up to 10% contamination (Fig. 8G and
477 8H). Our data display a minimal contribution of insoluble particles in the cap dolostone,
478 especially within the microbialites intervals, and the limestone succession of the cap
479 records the maximum detrital contribution. When analyzing the covariance of detrital
480 components between Y/Ho and Zr and Th, for example, low correlation is demonstrated
throughout the succession in total bulk rock dissolution. Another critical factor is that the anomalies remain unaltered, as Y/Ho, Ce/Ce, and Eu/Eu show similarities throughout the entire succession; we further discuss these observations in detail below. However, when applying conservative mixing lines between end members based on Y/Ho vs. Zr, Th, Al, and ΣREE + Y (Fig. 8), the cap limestone displays up to 30% contamination (Fig. 8A and 8B) (Fig. 8C, E and F), while Y/Ho vs. Al (Fig. 8B, C and D) is up to 1% detrital contribution, and Y/Ho vs. Th indicates up to 30% contamination (Fig. 8G and 8H).

## 6.2 Diagenetic influences on the Puga cap carbonate

Petrographic observations reveal that the matrix of the cap dolostone is predominantly composed of dolomicrite. It consists of replacement micritic pervasive dolomite with irregular intercrystalline boundaries (non-planar) and exhibits bright red to orange luminescence in cathodoluminescence. Texturally, a predominance of the grumose pattern is observed, manifested in diffuse peloids clustered in no individualized micro aggregates or isolated grumous. In rare instances, these peloids partly resemble shrubs, like those described by Frasier and Corsetti (2003). Additionally, replacement dolomite is present in all facies, forming a fine to medium planar-e to planar-s crystal mosaic with sizes ranging from 0.1 to 0.2 mm. The gypsum pseudomorph filling the fenestrae is well-developed. It has sizes ranging from 0.5 to 0.8 mm, occupying the majority of fenestral porosities within the rock, with its occurrence limited to the base of the section. The cement exhibits a fibrous habit, characterized by euhedral crystals that display two distinct growth forms: 1) fascicular and 2) radial/radiaxial, typical of swallowtail twinning. Subordinately, Fe-rich calcites occur mainly in fenestral, interpeloidal, and vug porosity. Furthermore, post-depositional alteration of the cap dolostone primarily occurred through processes such as dolomitization of gypsum,
neomorphism of dolomicrite, development of vugs and intercrystalline porosity, euhedral
dolomite precipitation, synsedimentary fracturing, calcite and quartz precipitation,
chemical compaction, Fe-Mn-oxide substitution, and precipitation (for more details, see
Soares et al. 2020, and Santos et al. 2021; Santos et al. 2023). The diagenetic processes
that the cap limestone was subjected to are suggested to have occurred in the following
sequence: pyrite precipitation, calcite cementation, neomorphism of micrite and crystal
fan, partial crystal fan dissolution, dolomitization, chemical compaction, fracturing,
ferrous dolomite and calcite precipitation, pyrite (pseudomorph) and micrite replacement
with Fe-oxide and replacement of rhombohedral dolomite by Mn-oxides (Soares et
al.2020; Santos et al. 2023).

The paragenetic relationships of the Puga cap carbonate demonstrate consistent
evidence of early diagenetic processes, with lesser influence from later mesodiagenetic
events (Soares et al., 2020; Santos et al., 2023). The cap dolostone shows a more
significant impact of syndepositional cementation, which may have enhanced the
preservation of the cap dolostone (Santos et al. 2021, 2023). This exceptional
preservation is observed in the well-preserved primary texture and what is interpreted to
be primary d¹³C and d¹⁸O signals from the new and previously published data (Nogueira
et al. 2003, 2019; Font et al. 2006, 2010; Sansjofre et al. 2011; Soares et al. 2020, Santos
et al. 2021, 2023). We discuss possible influences of late fluid-flow interaction in the Puga
cap carbonate using mobile trace elements, δ¹³C, and δ¹⁸O isotopic compositions to
evaluate syn- or post-diagenetic fluid-rock interactions (Derry, 2010; Brand and Veizer,
1981; Derry et al. 1989; Banner and Hanson, 1990; Jacobsen and Kaufman, 1999;
McArthur, 1994; Hohl et al. 2015). The overprint of these fluids can mask the pristine
geochemical compositions and thus must be thoroughly monitored (Hohl et al.
2015; 2022). All samples analyzed present negative δ¹⁸O values without systematic
correlations with δ¹³C values, which is suggestive of the preservation of primary signals (Fig. 9). The low correlation (R² = 6 x 10⁻⁵) between δ¹³C and δ¹⁸O suggests that the carbon isotope system was not significantly disturbed (Folling and Frimmel, 2002). Postdepositional fluids can modify d¹³C and d¹⁸O values (Jacobsen and Kaufman, 1999; Higgins et al., 2018; Crockford et al., 2021). Many mechanisms have been proposed to explain this close covariance in altered δ¹³C and δ¹⁸O, including early authigenesis in shallow marine sediments (Cui et al. 2017) and post-depositional processes such as interaction with meteoric water and diagenetic fluids during burial diagenesis (Taylor 1977; Banner and Hanson 1999; Knauth et al. 2017; Kennedy, 2009; Zao et al. 2020; Derry, 2010). Therefore, the δ¹³C and δ¹⁸O signals of the occurrences of the southern region of the Amazon Craton are interpreted as pristine carbonate, although we acknowledge that the application of additional diagenetic tracers such as δ⁴⁴Ca, δ²⁶Mg, or δ⁵⁷Li would potentially be insightful (Higgins et al., 2018; Ahm et al., 2018; 2019; Crockford et al., 2021). In most samples analyzed, Mg/Ca ratios averaged 0.8 ± 0.18, suggesting that dolomitization partially influenced the succession (Derry et al. 1994). However, the cross plot of δ¹⁸O vs. Mg/Ca shows a low correlation R² = 0.0182 (Fig. 9), suggesting that late dolomitization processes did not change the values of δ¹⁸O (Lu and Meyers, 1998). This low correlation between Mg/Ca vs. δ¹⁸O and δ¹³C vs. δ¹⁸O suggests that isotopes may have maintained pristine signals. Likewise, the low correlation of the Mg/Ca ratio with the Mn/Sr and Fe ratios (R² = 0.015) indicated that the trace elements did not change during dolomitization (Chen et al. 2022).

Mn/Sr and Sr/Ca variations can have two causes: postdepositional alteration or paleoenvironmental changes. However, our data demonstrated an anti-correlation. The Mn/Sr vs. δ¹⁸O of the cap dolostone showed an absence correlation (R² = 0.017), and the Sr/Ca vs. δ¹⁸O occurred with R² = 0.43, while Sr/Ca vs. Mg/Ca occurred with R² = 0.40.
Despite the low correlation, our data have a more significant relationship with
syndepositional dolomitization (Sr/Ca and Mg/Ca) than post-stabilization changes in
dolomite (Late diagenesis). Therefore, Mn/Sr and Sr/Ca ratios may result from specific
paleoceanographic conditions or the development of alteration processes by diagenetic
fluids (Asmerom et al., 1991; Derry et al., 1992; Kaufman et al., 1993; Kaufman and
Knoll, 1995; Melezhik et al. 2001; Folling and Frimmel, 2002; Halverson et al. 2007).
Thus, the high Mn/Sr ratios obtained in many cap dolostones may have been caused by
increased Mn concentrations in bottom waters or anoxic porewaters. (Hohl et al. 2017;
Chen et al. 2022). Another possible explanation for the high Mn concentration in cap
carbonates is that Fe-rich hydrothermal fluids in the coeval anoxic or suboxic deep ocean
(Yoshioka et al. 2003; Nédélec et al. 2007; Shen et al. 2005; Font et al. 2006; Hurtgen et
al. 2006).

The relationship between our petrographic and isotopic data suggests that
diagenetic processes may have occurred due to fluid (possibly seawater) rock interactions
during synsedimentary diagenesis or shallow burial (Nogueira et al. 2003; Font et al.
2004; Gilleaudeau et al. 2018; Soares et al. 2020; Santos et al. 2021, 2023). The evidence
of early diagenesis demonstrates an early stabilization of the Puga cap carbonate. This
evidence is confirmed in the dating of aragonite crystals fans from the cap limestone,
which indicates a U-Pb age of 623 million years (Carvalho et al. 2023), suggesting that
post-depositional alteration occurred within the first ≈10 million years. In contrast,
besides the majority of cap carbonate having the same textural and compositional
framework, the diagenetic history seems to be regional and completely dependent on the
tectonic, thermal, and burial history of the depositional basins where they were deposited
(Soares et al. 2020, Ahm et al., 2019, Santos et al., 2023). However, we suggested that the
cap carbonate succession preserves its syndepositional signature. We show that most
post-depositional modifications occurred early, and there is no evidence of effective overprints in the succession. This exceptional preservation qualifies the Puga cap carbonate succession as an excellent candidate for tracing the paleoceanographic dynamics after the Marinoan glaciation.

## 6.2.1 Diagenetic overprint interference in REE+Y distributions

Previous diagenetic evaluations of the Puga cap carbonate (Soares et al. 2020; Santos et al. 2021) demonstrate a predominance of early diagenesis process, which may have increased the potential preservation of cap dolostone. Diagenetic studies of REE+Y distributions in ancient carbonates suggest that ΣREE+Y patterns may be stable even during relatively intense diagenesis (Webb and Kamber, 2000; Liu et al. 2019) because ΣREE+Y concentrations are low in diagenetic fluids (10⁻⁶ to 10⁻⁴ ppm) (Sholkovitz et al. 1989; Banner and Hanson, 1990; McLennan et al. 2003), and the high permeability and porosity of sedimentary carbonates are strongly reduced by calcite cement formation during early diagenesis (Tanaka and Kawabe, 2006).

We analyzed Ce and Eu variations within binary diagrams. Ce and Eu were chosen for investigation due to their high sensitivity to fluid mobilization processes. Ce³⁺ has the potential to oxidize to Ce⁴⁺, while concurrently, Eu³⁺ can undergo reduction to Eu²⁺ (Brookins, 1989). Diagenetic overprinting of carbonate deposits will lead to REE enrichment and a decrease in (Dy/Sm). They would show strong covariance between Ce anomalies and (Dy/Sm) SN (German and Elderfield (1990), Shields and Stille, 2001). They could show a positive correlation between Eu/Eu* anomalies and (Pr/Sm) SN (Bolhar and Van Kranendonk, 2007). However, Ce/Ce* anomalies would not covary with (Dy/Sm) SN (Fig. 11), and anticorrelation between Eu/Eu* anomalies and (Pr/Sm) SN suggesting low interference of postdepositional diagenetic fluids in the ΣREE+Y of the
carbonates of the Puga cap carbonate. The Puga cap carbonate sequence shows strong Eu/Eu* anomalies ranging up to 3.2 (average =1.56 ± 0.31). Eu anomalies could be an analytical artifact due to interference by high Ba concentrations during ICP-MS measurements (Jiang et al. 2007). Our data show a low correlation with Ba (R² = 0.35). Also, the Eu/Eu* anomalies strongly correlate with Gd and Sm (R² = 0.9), which has nothing to do with isobaric interference with Ba. These positive Eu anomalies can be explained by plagioclase dissolution. Deng et al. (2017) and Langarica et al. (2020) discussed the influences of these variables on Eu anomalies and the flattening of the REE+Y patterns. However, this is not the case in the present study since we demonstrated in section 5.1 the purity of cap dolostone. Our data show low covariance between Eu/Eu anomalies and Zr, Al, and Th. Another possible explanation for the anomalies could be syn- and postdepositional hydrothermal processes because the hydrothermal fluids are a potentially significant secondary flow of REE+Y to the ocean, which carry flat or heavy rare earth (HREE)-depleted REE+Y signal and excessive concentrations of europium (Bau and Dulski, 1996; Douville et al. 1999; Bau e Dulski, 1999; Douville et al. 1999; Barrat et al. 2000; Cui et al. 2019). Furthermore, the previous discussion noted the low covariance of δ¹³C vs. δ¹⁸O and the absence of a correlation of Mg/Ca vs. δ¹⁸O (Fig. 9), which suggests minor modification after dolomitization, which agrees with the diagenetic studies by Soares et al. (2020) and Santos et al. (2023). These observations suggest that REE+Y patterns are most consistent with syn or penecontemporaneous processes before cap dolostone stabilization.

## 6.3 Constraints on the addition of meltwater in shallow platforms during the aftermath of the Marinoan glaciation

The REE + Y patterns of the Puga cap carbonate have distinct shale-normalized patterns compared to modern marine environments, which have uniform LREE depletion
and HREE enrichment, positive La anomalies, negative Ce anomalies, and high Y/Ho ratios (Bau and Dulski, 1996; Webb and Kamber, 2000; Kamber and Webb, 2001; Van Kranendonk et al. 2003; Bolhar et al. 2004; Shields and Webb, 2004; Zhang et al. 2016; Tepe e Bau, 2016). The Y/Ho ratio can differentiate between successions influenced by freshwater vs. seawater (Bau et al., 1996; Nozaki et al., 1997; Wang et al., 2014; Hodgskiss et al., 2021). As a result, marine carbonates often exhibit elevated values, i.e., they are considered superchondritic with Y/Ho ratios > 36 (Bau et al., 1996; Zhang and Nazaki, 1996; Nozaki et al., 1997). Conversely, REE + Y ratios remain relatively constant in freshwater environments, typically hovering around Y/Ho ≈ 26 (Lawrence et al., 2006; Bolhar and Van Kranendonk, 2007). Intermediate ratios are commonly associated with inland waters or mixed systems (Webb and Kamber, 2000; Van Kranendonk et al., 2003; Bolhar et al., 2004; Shields and Webb, 2004; Lawrence and Kamber, 2006). However, it has been frequently reported that marine carbonates may exhibit distinct REE + Y patterns compared to seawater, primarily due to the input of freshwater into shallow seawater (Frimmel, 2009; Zhao et al., 2009; Bolhar et al., 2015). Except for Ce, the typical marine rare earth element features can thus be attributed to estuarine or mixing water processes (Lawrence and Kamber, 2006).
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The cap dolostone occurs with Y/Ho ratios ranging from 24.50 to 76.67. Our Y/Ho values in the Tangará and Terconi sections are consistent with a mixture of waters leading to an intermediate salinity, for example, in the basal microbialite (stratiform doloboundstones with gypsum pseudomorphs) with an average value of 32 ± 10; n = 32.

However, intermittent elevated ratios (up to 70) of typical seawater are observed, implying sporadic episodes of salinity increase (Lawrence and Kamber, 2006). The domestic doloboundstone stromatolites with tubestone occur with values between 26 and 33 (average = 29 ± 1), suggesting the mixture of waters (Bolhar and Van Kranendonk, 2007).

The wave-dominated dolostone facies (e.g., Peloid dolopacktone with quasi-planar lamination) occur with Y/Ho = 32 to 34 (average = 33.5 ± 0.6) close to marine values coastal. These ratios are common in glacial meltwater and seawater, characterized by generally low Y/Ho values relative to terrigenous REE + Y patterns (Frimmel, 2009; Zhao and Zheng, 2010). These patterns are widely recorded in cap dolostone deposits (REFS).

They are commonly interpreted as reflecting the dilution of the seawater signal in shallow marine environments by the influx of meltwater derived from continents in the aftermath of the Marinoan glaciation (~635 Ma). The plume world hypothesis proposed that during the massive Marinoan deglaciation, a thick layer of floating freshwater plume was produced from continental ice melting, creating physical and chemical stratification of the oceans (Shields, 2005; Yang et al., 2017). Upwelling events (e.g., oceanic circulation) disrupt sea stratification (Liu et al., 2014; Rame and Marotzke, 2022). Our mineralogical and geochemical data support this interpretation. During these events, hypersaline deep water reaches shallow water, causing increased salinity, high values of Y/Ho (up to 71), and gypsum precipitation in the fenestral porosities of basal cap dolostone.
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### 6.3.1 Upwelling of hydrothermal and hypersaline fluids in the shallow platform

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The Puga cap carbonate exhibits mostly weakly negative Eu anomalies throughout both sequences; however, elevated values are documented at the base of the cap dolostone (microbialites level). High Eu anomalies are absent from the general characteristics of marine carbonates (Lawrence et al. 2006) and are generally regarded as strong indicators of high temperature-reducing fluids (Bau, 1991; Klinkhammer et al. 1994; Barrat et al. 2000; Alexander et al., 2008; Parsapoor et al. 2009). REE-depleted patterns with positive Eu anomalies indicate the contribution of hydrothermal fluids (Cui et al. 2019). Our data are consistent with positive Eu/Eu values, which are generally considered a sign of hydrothermal components in previous studies of post-Marinoan cap dolostones (Font et al. 2006; Nédélec et al. 2007; Huang et al. 2011; Meyer et al. 2013; Wang et al. 2014; Rodler et al. 2016; Hohl et al. 2017; Xiao et al. 2020).

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694

We tested the hydrothermal influence on cap carbonates using Fe/Ti vs. Al/ (Al + Fe + Mn) (Bostrom, 1973; Barrett, 1981). Deposits with a strong hydrothermal influence don't typically occur with high concentrations of Al, resulting in high Al/Ti ratios (Marchig et al. 1982). The contamination of these deposits by pelagic and terrigenous seabed sediments enriches them in components such as Ti and Al, resulting in a drastic reduction of Fe/Ti ratios and an increase in the Al/ (Al + Fe + Mn) ratio. High FeT/Al ratios but deficient concentrations of extractable pyrites suggest ferruginous rather than euxinic conditions (Huang et al. 2011). In addition, REE absorption by Fe oxides is typically associated with submarine hydrothermal plume activity (German et al. 1991; Ling et al. 2013).
Our data suggest a robust hydrothermal influence with respect to siliciclastic contamination (Fig. 10A). The Eu/Eu anomalies are strongly correlated with Fe/Al ratios (Fig. 10B). This relation of the increase in Fe accompanying the strong Eu anomalies reinforces the interpretation of a hydrothermal contribution for the Puga cap carbonate, similar to what was described by Jing Huang in 2011, which suggests a hydrothermal origin of elevated iron and manganese for the Doushantou cap carbonate from China. Furthermore, we assessed the hydrothermal impact by applying conservative mixing lines (see Fig. 10 D, E, and F). This approach aligns with the model proposed by Alexander et al. (2008), which considers the interaction of high T-hydrothermal fluids with seawater during sediment deposition using the MREE/HREE ratios. To achieve this, we integrated data from high T-hydrothermal sources (>350 °C, as per Bau and Dulski, 1999) and shallow seawater components (as per Alibo and Nozaki, 1999). Conservative mixing calculations between end members based on Sm/Yb vs. Eu/Sm (Fig. 10 D) indicate that a high-T hydrothermal fluid contribution of less than 1% could have been responsible for the anomalies recorded in the cap dolostone (Alexander et al. 2009). The Eu anomalies and the strong Sm/Yb and La/Yb ratios suggest interaction with hydrothermal syndepositional fluids (Fig. 10). The studied samples were plotted distant from the mixing line (between 0.1 and 1% hydrothermal fluids). This scenario suggests no influence of high-temperature fluids and that the depositional environment of carbonate rocks was distant from fumaroles. However, positive Eu anomalies obtained from Archean and Paleoproterozoic chemical sediments are interpreted as mixtures of seawater and hydrothermal fluids (Kamber and Webb, 2001; Frei et al., 2008; Hodgskiss et al., 2021). For the Ediacaran, positive Eu anomalies are typical for the deep anoxic ocean (Kamber and Webb, 2011; Bolhar et al., 2005; Allwood et al., 2010; Rodler et al., 2016). In addition, our interpretations are similar to recent interpretations for the origin of the
Doushantou cap carbonates, which are mainly based on δ¹³C, δ³⁴SCAS, δ¹⁷O_CAS, δ⁴⁴Ca, 720
⁸⁷Sr/⁸⁶Sr, rare earth, elements, and metal trace elements. These data suggest that they may 721
722 have been formed under the influence of deglacial meltwater, hydrothermal fluids, and
723 upwelling (Huang et al. 2009; Huang et al. 2011; Huang et al. 2013; Wang et al. 2014;
724 Hohl et al. 2017; Zhao et al. 2018; Wei et al. 2019a; Peng et al., 2022).

Carbonates associated with deep ocean brine are characterized by Fe and trace 725
metal enrichments with positive Eu anomalies (Meyer et al. 2012). Liu et al. (2013) 726
727 extensively discussed the potential origin of radiogenic ⁸⁷Sr/⁸⁶Sr in Marinoan cap
728 carbonates of the Nuccaleena Formation, South Australia. They found that the interaction
729 of the carbonates with brines or hydrothermal fluids could better explain these values, or
730 they may reflect values from glacial meltwater mixing with seawater. This evidence
731 strengthens previous interpretations that the Marinoan cap carbonate precipitated in a
732 mixture of coastal meltwater with saline Mg-rich deeper waters through the upwelling
733 process (Grotzinger and Knoll, 1995; Ridgwell et al., 2003; Shields, 2006; Yang et al.,
734 2017; Liu et al. 2014; 2018; Hoffman et al. 2017; Santos et al. 2021; Rame and Marotzke,
735 2022). Through modeling, a glacial hydrothermal flow during the Marinoan snowball
736 glaciation was demonstrated using a binary mixture of Ge/Si and Eu/Eu* values (Cui et
737 al., 2016). Marinoan deposits were accumulated during the drifting stage of post-rift
738 subsidence (Hoffman et al. 2017), and this hydrothermal flow during Snowball Earth
739 events could potentially enrich the oceans in certain elements in the aftermath of the
740 Marinoan glaciation.

The syndepositional hydrothermal event was unlikely, and these observations 741
align with our interpretations of rapid precipitation of the Marinoan cap carbonate 742
(Hoffman et al. 1998; Nogueira et al. 2003; 2019; Cui et al. 2016; Zhou et al. 2019; Santos 743
et al. 2021). Our data suggests some samples with typical seawater values (Y/Ho up to 744
745 70) others displaying strong indications of hydrothermal fluid interaction (Eu/Eu* up to 746 3.2). Nevertheless, most of our samples exhibit Y/Ho ratios typical of mixwater, 747 approximately ~ 33 (a combination of glacial meltwater and seawater). This finding aligns 748 consistently with interpretations of upwelling events during the precipitation of the cap 749 dolostone. These deep Cryogenian waters would have interacted with high-temperature 750 hydrothermal fluids responsible for the elemental enrichments (e.g., Eu, Fe, Mn). At this 751 time, deep waters would have interacted with high-temperature hydrothermal fluids 752 responsible for the elemental enrichments (e.g., Eu, Fe, Mn). This evidence is consistent 753 with the diachronous precipitation model for the Marinoan cap dolostone from the melt- 754 water layer during syn-deglacial transgression (Hoffman et al. 2017).

755 This upwelling interaction may have also lowered the Y/Ho values of the 756 dolostone samples in the basal succession, as hydrothermal fluids tend to possess low 757 Y/Ho ratios. (Allwood et al. 2010). In addition, as discussed in section 5.1, we 758 demonstrated the absence of interference of siliciclastic components at lower Y/Ho 759 values. The base of the cap dolostone shows a fluctuation in the Y/Ho values and high 760 values of Eu anomalies, suggesting a strong contribution from hydrothermal sources of 761 REE+Y to the cap dolostone of the southeast margin of the Amazon Craton.

762 6.4 Paleoredox indications

763 Ce/Ce* anomalies in the Puga cap carbonate indicate oxygen deficiency 764 throughout the section (Fig. 11), but these anomalies are weaker in the microbialite 765 interval. In modern seawater, REE+Y (Rare Earth Elements + Yttrium) can record the 766 evolution of seawater redox, and the negative Ce anomaly marks oxidized conditions 767 (Byrne and Sholkovitz, 1996; Webb and Kamber, 2000; Ling et al., 2013), while the 768 absence of anomalies is indicative of anoxic conditions (Lawrence et al., 2006). 769 Furthermore, soluble Ce³⁺ is easily oxidized to Ce⁴⁺, which is relatively insoluble in oxic
770 seawater (McLennan, 1989; Alibo and Nozaki, 1999; Pattan et al., 2005; Frimmel, 2009).

771 Therefore, the Ce anomaly in marine sediments distinguishes oxic from anoxic
depositional environments throughout the sedimentary record (Bau et al., 1996; de Baar
et al., 2018).

774 However, the scenario presented thus far demonstrates the influence of water
775 mixing, particularly a strong influence of meltwater. The Y/Ho ratio exhibited similar
776 behavior to estuarine dynamics (freshwater and saltwater). Except for Ce, typical marine
777 rare earth element features can be attributed to estuarine or mixing water processes since
778 Ce anomalies only form in salinities greater than 2% (the ocean average is 3.5%)
779 (Lawrence and Kamber, 2006). Therefore, only samples with a Y/Ho ratio ≥ 36 will be
780 considered for interpretation. Three samples fall into the oxygenated category, but most
781 of these eligible samples exhibit values suggesting dysoxic conditions. Our results have
782 been extensively tested for post-depositional changes; at least, the samples with high
783 Y/Ho ratios may reflect the redox conditions of our post-Marinoan scenario.

784 The interpretation of a dysoxic to anoxic water environment in the Puga cap
785 carbonate is consistent with other post-Marinoan deposits (Huang et al., 2009; Huang et
786 al., 2011; Wang et al., 2014; Lang et al., 2016; Rodler et al., 2016; Hohl et al., 2017;
787 Caxito et al., 2018; Wu et al., 2019; Hu et al., 2023; Ren and Li, 2023). In this sense, it
788 could reflect the biochemical conditions of this anomalous environment, as Ce anomalies
789 are known to be affected by environmental parameters such as pH, alkalinity, and
790 microbial activity (De Baar et al., 1985a; Alemão and Elderfield, 1990; Byrne and Kim,
791 1993; Pourret et al., 2008; Zhao et al., 2022). The Ce/Ce* values of the studied samples
792 (0.70–1.05, mostly below 0.95) could indicate a small amount of oxygen deficiency in
793 the water column (Fig. 11). Dysoxic waters would have likely provided even more
794 stressful conditions for microbial life after the Marinoan glaciation. Generally, oxidizing
conditions prevailed in the atmosphere and shallow seawater during the post-Marinoan phase (Fike et al., 2004; Sahoo et al., 2012; Sansjofre et al., 2014; Hohl et al., 2015; 2017; Wy et al., 20xx; Ader et al., 2014). Sansjofre et al. (2014), based on trace metals, suggested that a water column must have been essentially oxic during the cap dolostone precipitation, with anoxia only sporadically reaching the sediment-water column interface in the deepest parts of the platform on the Amazon craton. The interpretation of these authors differs in part from our findings, which indicate predominantly dysoxic water. However, other proxies must be applied to understand microbial communities' role and evaluate the reservoir effect due to water mixing.

## 6.5 CAP CARBONATE PRECIPITATION MODEL AND GLOBAL CORRELATION

The cap carbonate immediately precipitated over the Puga diamictite in a coastal-marine environment influenced by microbial mat proliferation during the transition from icehouse to greenhouse conditions (Nogueira et al. 2003; 2019; 2019; Font et al. 2006, 2010; Nédélec et al. 2007; Sansjofre et al. 2014; Santos et al. 2021, 2023). The deformed dolomite-diamictite contact indicates rapid accumulation of dolomicrite deposited over a still plastic diamicton linked to a regional scale event compatible with the glacial isostatic adjustments on the coastal zone, forming a substrate with irregular morphology (Nogueira et al. 2003; 2019). Rapid precipitation (≈10³ – 10⁴ years) for cap dolostones is consistent with our sedimentological evidence and other models, such as enhanced continental weathering, alkalinity addition, gas hydrate destabilization, and persistent freshwater surface layer (Shields, 2005; Hoffman et al., 2017; Crockford et al. 2021). Complete glacier retreat and ocean warming were succeeded by isostatic adjustments, characterized by an uplift of continental areas and eustatic subsidence on the coastal zone, concomitant with relative post-glacial sea-level rise (Creveling and Mitrovica, 2014) and during the advance of the syn-deglacial transgression, meltwater
production and surface warming created stable density stratification (Liu et al. 2014;2018; Yang et al. 2017; Hoffman et al. 2017). The diamicton substrate was colonized by microbial communities in stratified and shallow water, and the high organic productivity due to the maximum nutrient input from mixed water culminated in the giant biostromas built on the coast (Santos et al. 2021). The microbial communities that colonized the diamicton recorded not only the influences of the chemistry of the environment but also the effects of the hydrodynamics of the environment. According to Yang et al. (2017), water stagnation and high density led to the development of lamination (Santos et al. 2021).

The negative δ¹³C isotope signatures can be considered a consequence of microbial activity, producing organically derivatived ¹³C-depleted bicarbonate (Nedeléc et al. 2007). These interpretations are consistent with our δ¹³C data ranging from -8.2‰ to -4.1‰ (average = -4.9‰); as discussed earlier, our tests suggested values that express Early Ediacaran seawater chemistry, as well as previous data for Puga cap carbonate (Nogueira et al. 2003;2019; Fonte et al. 2006; Sansjofre et al. 2011) similar to Marinoan cap dolostone successions worldwide (Nedeléc et al 2007; Hoffman et al., 2021; Peng et al., 2022). Thus, the negative carbon isotope values may reflect the high productivity after the Marinoan glaciation in a dysoxic water column, with intense microbial activity on the paleocoast due to rising temperatures and significant inputs of nutrients provided by reactivation of ocean biogeochemical cycling and an increased hydrological cycle (Kunzmann 2012; Sansjofre et al. 2014; He et al. 2021). These conditions favored the proliferation and preservation of a giant biostroma laterally continuous for kilometers in the paleocoast (Nogueira et al. 2003;2022; Font et al. 2006; Romero et al. 2020; Santos et al.2021,2023).
During deglaciation, continental ice sheets trapped detrital materials (Hoffman et al. 2017; Yu et al. 2020). Estimating how long the sediments were trapped in the ice sheets in the distal zone before reaching the coastal area is impossible. However, rapid precipitation for cap dolostones is consistent with our sedimentological evidence. Our data show intense participation with microbial activity (giant stromatolites), an abrupt increase in alkalinity (cap dolostone-cap limestone transitions), plume world (Mixwater and gypsum cement), and siliciclastic starvation (pure carbonates). Rapid precipitation is consistent with other models, such as continental weathering alkalinity, gas hydrate destabilization, and plume world, as well as sedimentary features and geochemical models (~$10^3 – 10^4$ years) (Shields, 2005; Crockford et al., 2016; Hoffman et al., 2017; Yang et al., 2017; Crockford et al. 2021). However, a single model does not explain the precipitation of the Puga cap carbonate despite the similarity between different sequences of this age. A reason for this heterogeneity is that there is an endemic distinct preservation between locations, primarily related to the paleogeographical position that dictates the coastal/marine dynamics.

We have previously demonstrated the purity of cap dolostone, between 1 and 5% contamination, based on Y/Ho vs. Zr and Al (see item 5.1). These conditions of high density, up to 3 times higher than current rates, according to the model by Yang et al. 2017, may have controlled the sediment inflow to the shallow carbonate platform, indicated by the low content of detrital components in the cap dolostone (figs). The density between meltwater and seawater or even sediment starvation can explain the low content of detrital constituents. The starvation of sediments indicates precipitation of the cap dolostone over ~$10^5 – 10^6$ years (Spence et al. 2016; Wallace et al. 2019). A larger landward shoreline migration in the siliciclastic starvation model in cap deposition is not directly associated with final Snowball Earth conditions (Nordsvan et al. 2019). During
869 deglaciation, detrital materials were trapped in continental ice sheets (Hoffman et al. 2017; Yu et al. 2020). In modern shelves, the area starved of sediment is associated with detritic trapping that can occur at the terrestrial limit with the saline wedge, producing maximum estuarine turbidity (Swift, 1968; Kidwell, 1991; Geye et al. 2004; McHugh et al., 2010). The generation of hyperpycnal flows to transport fine sediments distributed in surficial plumes that only occur in very shallow waters can extend for thousands of kilometers (Geyer et al. 2004; Dowdeswell et al. 2015. The model by Yang et al. (2017) postulates a salinity three times higher than the current level. In this sense, the detrital content in meltwaters from the Marinoan deglaciation could be trapped during the linkage of freshwater and hypersaline seawater, forming extensive plumes with thousands of kilometers.

880 However, our geochemical data does not agree with a higher salinity during fast cap dolostone precipitation in mixwaters as previously demonstrated (Y/Ho ratios of the mixwater patterns). Another possible explanation for siliciclastic starvation could have arisen due to the restricted conditions with persistently shallow waters in the coastal zone.
881
882
883 This depositional scenario was strongly influenced by glacial isostatic adjustment (GIA) or rebound but affected only the initial meters of cap dolostone (e.g., Nogueira et al. 2003). This process led to continuous uplift in conjunction with a rapid post-glacial sea-level rise, maintaining a shallow shelf isolated from direct oceanic processes. Barrier-inland facies were not observed in the studied succession, indicating that irregular diamicton morphologies may have created protected areas against waves, promoting stratiform microbial proliferation and inducing dolomite precipitation. This scenario resulted in a consistently shallow water column, following a significant landward shoreline migration depositional pattern with a high potential for preserving the same
884
885
886
887
888
889
890
891
892
893

signature, devoid of detritic components, found for hundreds of kilometers along the
southern margin of the Amazon Craton.

894

895

The conditions outlined above likely lasted for approximately ~8 ka (Liu et al. 2014) or would have occurred on a scale of 10⁴ and 10⁵ years following a unidirectional mixing model associated with thermal expansion (Yang et al. 2017). Another hypothesis suggests only a few thousand years for the mixing waters based on a three-dimensional model considering oceanic circulation as a crucial factor for the end of stratification by meltwater (Rame and Marotzke, 2022). Although the glacial-isostatic adjustment (GIA) could be the driving force influencing the mixing of the waters and paleoceanography, the time spent developing this event depends on the ice cap thickness and melting velocity. The lack of terrigenous components in cap dolostone may be evidence of rapid precipitation (keep-up) before the main detritic flow on the coast of the Amazon craton or lasted at least as long as the stratified oceans.

896

897

898

899

900

901

902

903

904

905

906

## 6.6 Implications for the Carbonatic Factory Change

907

908

909

910

911

912

913

914

915

916

Water turbidity increases partially inhibit carbonate precipitation, causing marl deposition and suppressing microbial life, a prelude to marine stratification declines and thermal expansion (Romero et al. 2020; Santos et al. 2021; Nogueira et al. 2022). The mixing waters and massive detritic flow generated a geochemical change, causing CaCO3 over-saturation, which contributed to the precipitation of poloidal limestones with calcite crystal fans (pseudomorphs after aragonite) (Higgins and Schag, 2003; Nogueira et al. 2003, 2007, 2019; Hoffman et al. 2017,2021; Soares et al. 2020; Santos et al. 2021,2023). The cap limestone presents an abrupt increase in the Sr concentration that accompanies a rise in Zr, Al, and Th in covariance with REE+Y), indicating a direct relationship between
917 the siliciclastic input and the increases in alkalinity and, as a consequence, a change in
918 the dolomitic-aragonitic features.

919 Our geochemical and microfacies data support an abrupt transition in the carbonate
920 factory, as enhanced continental weathering could have lowered seawater Mg/Ca ratios
921 with substantial input of Ca2+ to platform environments (Hoffman et al.1998, xx),
922 triggering CaCO3 oversaturation (Nogueira et al. 2019). Marine chemistry changes are
923 commonly associated with secular modifications such as the Mg/Ca ratio, pH, redox state
924 of the ocean, climate (icehouse-greenhouse conditions), pCO2, and the extent of
925 continental weathering (Hardie, 1996, 2003; Hood and Wallace, 2018; Ries, 2010;
926 Sandberg, 1983; Stanley and Hardie, 1998; Wood et al., 2017; Wey et al., 2022). The cap
927 dolostone-cap limestone transition saw an increase in the concentration of Sr in
928 covariance with Zr and Th (Fix and Fix). The cap dolostone in this study displayed a low
929 Sr content, ranging from 30 to 70 ppm, in contrast to the cap limestones with Sr 300-1400
930 ppm. The interpretation of primary dolomite precipitation by Nogueira et al.2003; 2019,
931 Font et al. 2006; Romero et al. 2020; Soares et al. 2020, and Santos et al. 2021, 2023 is
932 consistent with our data because ancient dolomite commonly has significantly lower Sr
933 concentrations (20-70 ppm) (Baker and Burns, 1985; Vahrenkamp and Swart, 1990).
934 This abrupt transition in Sr concentrations follows a decrease in the Mg/Ca ratio (0.9 to
935 0.0) (see Fig. x). According to Ahm et al. (2019), dolomitization related to an aragonitic
936 primary fabric is consistent with fluid flow and seawater upwelling during global
937 deglaciation. Our data agree with an upwelling event and a syn-deglacial dolomitization
938 event, as previously discussed:

1. Mixing of meltwater and seawater (Y/Ho = 27 to 71)

2. Shallow seawater interacting with deep water (high Eu/Eu anomalies, suggesting
the upwelling of syndepositional T-high hydrothermal fluids)
942

3. Low correlation between Mg/Ca and δ¹³C and δ¹⁸O, suggesting that dolomitization occurred at least early diagenesis.

943

944

4. The excellent preservation of gypsum pseudomorphs indicates early dolomitization (see Santos et al. 2021, 2023).

945

946

Despite our evidence agreeing with upwelling events, our sedimentological data disagree with an abrupt increase in water depth as a controller for the last dolomitization event. Although limestone beds are associated with a moderately deep platform, aragonite crystal fans are truncated, indicating wave action in shallow waters (Soares et al. 2020), and these precipitated on the seafloor were possibly dispersed as micrite by oscillation (Viera et al. 2015). The hydrodynamic conditions were similar during upper cap dolostone and cap limestone precipitation in a shallow platform influenced by wave actions. Despite this evidence, the product of the interaction with the mats still needs to be clarified if we initially have a precipitated proto-dolomite according to the organogenic model (Vasconcelos et al. 1994, Dupraz 2009; Qiu et al. 2017; Chang et al. 20) or if it was initially aragonite as suggested by Ahm et al. 2019. The concentration of Mg remained constant during the restriction of paleoenvironmental conditions on dolomitic platforms. The rapid sea-level rise facilitated more excellent dispersion of Mg ions, leading to the definitive interruption of dolomicrite precipitation, abruptly replaced by CaCO₃ supersaturated seas with massive aragonite precipitation influenced by a long-lasting transgression, still occurring during the Lower Ediacaran.

947

948

949

950

951

952

953

954

955

956

957

958

959

960

961

## 7. CONCLUSION

962

963

964

The REE+Y patterns and δ¹³C values observed in the Puga cap carbonate provide valuable insights into post-snowball Earth paleoceanographic conditions. The low Y/Ho ratios in the cap dolostones suggest dilution of seawater composition due to meltwater

965

966
influx from continents (Y/Ho ~29). Conversely, superchondritic Y/Ho ratios up to 71 in the basal cap dolostones suggest upwelling of hypersaline seawater in coastal areas. In addition, the flattened REE+Y pattern with a positive Eu anomaly (up to 3.2) may suggest meltwater with influences from a region with volcanic rocks (e.g., Amazon Craton).
REE+Y mass fractions in such settings may not accurately reflect global ocean water composition; instead, they primarily fractionate in response to local factors related to alkalinity and freshwater mixing. These findings imply that ocean destratification is associated with ocean circulation and glacial isostatic adjustment, with reduced siliciclastic inflow supporting a siliciclastic starvation model or rapid precipitation. The recurrence of shallow-water conditions, attributed to early evidence of relative sealevel decline, is linked to local ice sheet recession (ice gravity). This led to the continuous uplift of the coastal zone, the creation of an isolated shelf protected by ocean dynamics, and deformation in diamictite, resulting in irregular substrate relief morphologies. This phase witnessed the flourishing of extensive microbial communities and the rapid recovery of primary productivity, which induced dolomite precipitation in the coastal setting. Mixing water masses and massive detrital input, as indicated by the increase of insoluble elements (e.g., REE, Zr, and Th), generated a geochemical change caused by CaCO3-supersaturated seas influenced by significant landward shoreline migration during the longterm transgression. These observations comprehensively explain the complex interplay between post-glaciation and paleoceanographic dynamics during the Cryogenian-Ediacaran boundary.

988
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1570
<figure><img src="image_3.png" /></figure>

<figure><img src="image_4.png" /></figure>

ALTO PARAGUAI
GROUP

<figure><img src="image_5.png" /></figure>

Ediacara-Cambrian
Unconformity
~80 Myr

<figure><img src="image_6.png" /></figure>

Skolithos

~528 Ma

Uplift of southern margin of Amazon Craton

<figure><img src="image_7.png" /></figure>

$^{87}$Sr/$^{86}$Sr

0,7088

0,7081

0,7088

0,7085

0,7081

0,7085

0,7085

0,7086

0,7082

0,7089

0,7079

0,7080

0,7076

0,7079

0,7077

0,7076

0,7074

ARARAS GROUP

Nobres
Formation

<figure><img src="image_8.png" /></figure>

Serra do
Quilombo
Formation

<figure><img src="image_9.png" /></figure>

Guia
Formation

<figure><img src="image_10.png" /></figure>

Mirassol
d'Oeste
Formation

CAP

<figure><img src="image_11.png" /></figure>

Puga Formation

Cuiabá Group (Paraguai belt)

Cryogenian
Eduardian
Intracratonic basin

Metamorphic basement
Collisional Event

~720 Ma

ARARAS-ALTO PARAGUAI BASIN

Cambrian

EDiacaran

Cryogenian
<figure><img src="image_12.png" /></figure>

## Southern Amazon Craton
### Puga cap carbonate

<figure><img src="image_13.png" /></figure>

### FACIES ASSOCIATIONS

1. Glacial marine

2. Shallow platform with microbial activity

3. Wave- and storm-dominated shallow platform

4. $CaCO_3$-Over-saturated moderately deep platform

5. Anoxic deep platform

<figure><img src="image_14.png" /></figure>

<figure><img src="image_15.png" /></figure>

<figure><img src="image_16.png" /></figure>

<figure><img src="image_17.png" /></figure>

Cryogenian-Ediacaran boundary

<figure><img src="image_18.png" /></figure>

Siltstone/shale

<figure><img src="image_19.png" /></figure>

Fine dolostone

<figure><img src="image_20.png" /></figure>

Fine sandstone

<figure><img src="image_21.png" /></figure>

Current ripple marks

<figure><img src="image_22.png" /></figure>

Load cast, slump fold and massive truncanted lamination

<figure><img src="image_23.png" /></figure>

Tube-like structure-

<figure><img src="image_24.png" /></figure>

Fracture

<figure><img src="image_25.png" /></figure>

Fenestral porosity

<figure><img src="image_26.png" /></figure>

Fine limestone

<figure><img src="image_27.png" /></figure>

Peloidal dolostone

<figure><img src="image_28.png" /></figure>

Jurassic basalt

<figure><img src="image_29.png" /></figure>

Neptunian dike

<figure><img src="image_30.png" /></figure>

Low-angle truncated lamination

<figure><img src="image_31.png" /></figure>

Stratiform stromatolite

<figure><img src="image_32.png" /></figure>

Slump fold

<figure><img src="image_33.png" /></figure>

Terrigenous grains

<figure><img src="image_34.png" /></figure>

Calcareous breccia

<figure><img src="image_35.png" /></figure>

Diamictite

<figure><img src="image_36.png" /></figure>

Siltstone/shale

<figure><img src="image_37.png" /></figure>

Fibrous calcite

<figure><img src="image_38.png" /></figure>

Even parallel lamination

<figure><img src="image_39.png" /></figure>

Massive bedding

<figure><img src="image_40.png" /></figure>

Megaripple bedding

<figure><img src="image_41.png" /></figure>

Gypsum Pseudomorph

<figure><img src="image_42.png" /></figure>

Calcite crystal fans

<figure><img src="image_43.png" /></figure>

Sinsedimentary fault

<figure><img src="image_44.png" /></figure>

Macropeloid

<figure><img src="image_45.png" /></figure>

Dropstone

<figure><img src="image_46.png" /></figure>

Striated clasts

<figure><img src="image_47.png" /></figure>

Bitumen

This study (2)

(4)

(6)

(3)

(7)

<figure><img src="image_48.png" /></figure>

<figure><img src="image_49.png" /></figure>

<figure><img src="image_50.png" /></figure>
<figure><img src="image_51.png" /></figure>

<figure><img src="image_52.png" /></figure>

<figure><img src="image_53.png" /></figure>

<figure><img src="image_54.png" /></figure>

<figure><img src="image_55.png" /></figure>

<figure><img src="image_56.png" /></figure>

<figure><img src="image_57.png" /></figure>

<figure><img src="image_58.png" /></figure>
Tangará section

<figure><img src="image_59.png" /></figure>

<figure><img src="image_60.png" /></figure>

Cap limestone

<figure><img src="image_61.png" /></figure>

Cap dolostone

<figure><img src="image_62.png" /></figure>

Puga diamicitite

1-Diamicitite

2- Microbialites

3-Wave dominated dolostone

4- limestone

<figure><img src="image_63.png" /></figure>

Crystal fans
(Aragonite Pseudomorphs)

<figure><img src="image_64.png" /></figure>

Tubestone

<figure><img src="image_65.png" /></figure>

Gypsum cements
(Pseudomorphs)

<figure><img src="image_66.png" /></figure>

Y/Ho
>36 (Seawater)

<figure><img src="image_67.png" /></figure>

27-36 = (Mixwater)

<figure><img src="image_68.png" /></figure>

26 = Siliciclastic signal

<figure><img src="image_69.png" /></figure>

< 4 ppm (PC)

<figure><img src="image_70.png" /></figure>

> 4 ppm (IC)

<figure><img src="image_71.png" /></figure>

IC = Impure carbonate

<figure><img src="image_72.png" /></figure>

<0.6 ppm (PC)

<figure><img src="image_73.png" /></figure>

>1 ppm (IC)

PC = Pure carbonate

● Partial dissolution

● Bulk rock

Increase detritic
components

Low siliciclastic content
Terconi section

<figure><img src="image_74.png" /></figure>

1-Diamiticite

2- Microbialites

3-Wave dominated dolostone

4- limestone

<figure><img src="image_75.png" /></figure>

Crystal fans

(Aragonite Pseudomorphs)

<figure><img src="image_76.png" /></figure>

Tabestone

<figure><img src="image_77.png" /></figure>

Cypsum cements

(Pseudomorphs)

<figure><img src="image_78.png" /></figure>

>36 (Seawater)

<figure><img src="image_79.png" /></figure>

26-36(Mixwater)

<figure><img src="image_80.png" /></figure>

25 = Siliciclastic signal

IC = Impure carbonate

<figure><img src="image_81.png" /></figure>

<0.6 ppm (PC)

<figure><img src="image_82.png" /></figure>

>1 ppm (IC)

<figure><img src="image_83.png" /></figure>

PC = Pure carbonate

Bulk rock

Low siliciclastic content

Increase diagenetic
components
Samples/PAAS

<figure><img src="image_84.png" /></figure>

B

<figure><img src="image_85.png" /></figure>

Samples/PAAS

<figure><img src="image_86.png" /></figure>

C

<figure><img src="image_87.png" /></figure>

D
<figure><img src="image_88.png" /></figure>

<figure><img src="image_89.png" /></figure>

<figure><img src="image_90.png" /></figure>
A

<figure><img src="image_91.png" /></figure>

B

<figure><img src="image_92.png" /></figure>

C

<figure><img src="image_93.png" /></figure>

D

<figure><img src="image_94.png" /></figure>

E

<figure><img src="image_95.png" /></figure>

F

<figure><img src="image_96.png" /></figure>
δ¹³C

<figure><img src="image_97.png" /></figure>

Mg/Ca

<figure><img src="image_98.png" /></figure>

Mn/Sr

<figure><img src="image_99.png" /></figure>

Sr/Ca

<figure><img src="image_100.png" /></figure>
<figure><img src="image_101.png" /></figure>

<figure><img src="image_102.png" /></figure>

<figure><img src="image_103.png" /></figure>

<figure><img src="image_104.png" /></figure>
~635 Ma
Paleogeographic Reconstruction

<figure><img src="image_105.png" /></figure>

<figure><img src="image_106.png" /></figure>

<figure><img src="image_107.png" /></figure>
Amazon craton margin

<figure><img src="image_108.png" /></figure>