Phosphatic Peritidal Limestone of the Neoproterozoic Salitre Formation, Brazil, and Precambrian Economic Phosphorite

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PHOSPHATIC PERITIDAL LIMESTONE OF THE NEOPROTEROZOIC SALITRE FORMATION, BRAZIL, AND PRECAMBRIAN ECONOMIC PHOSPHORITE

RENEE ALAYNE DELISLE

Thesis submitted in partial fulfillment of the requirements for the Degree of Master of Science, Geology

Acadia University Spring Convocation 2015

TABLE OF CONTENTS

TITLE PAGE ...... i

ACKNOWLEDGEMENT OF DEFENSE ...... ii

PERMISSION FOR REPRODUCTION ...... iii

TABLE OF FIGURES ...... vi

ABSTRACT ...... viii

ACKNOWLEDGEMENTS ...... ix

CHAPTER 1: INTRODUCTION ...... 1

CHAPTER 2: BACKGROUND ...... 3 2.1 Neoproterozoic climate and P-cycle ...... 3 2.2 Phosphorites and phosphogenesis ...... 4 2.3 Geochemistry of phosphorites and associated sedimentary rocks ...... 9

CHAPTER 3: GENERAL GEOLOGY ...... 13 3.1 Tectonic Setting ...... 13 3.2 Una Group ...... 18

CHAPTER 4: METHODS ...... 20 4.1 Field Methods ...... 20 4.2 Laboratory Methods ...... 21

CHAPTER 5: LITHOFACIES ...... 24 5.1 Facies 1: Tabular bedded intraclast-rich grainstone ...... 24 5.2 Facies 2: Cross-stratified intraclastic grainstone ...... 28 5.3 Facies 3: Flaser to lenticular-bedded packstone ...... 30 5.4 Facies 4: Hemispheroidal columnar stromatolites ...... 33 5.5 Facies 5: Tabular phosphatic columnar stromatolites ...... 36 5.6 Facies 6: Hummocky cross-stratified grainstone ...... 39

CHAPTER 6: PERITIDAL CYCLES AND SEQUENCE STRATIGRAPHY ...... 42 6.1 Peritidal cycles and cycle development ...... 42 6.2 Sequence Stratigraphy ...... 50 6.2.1 LST ...... 50 6.2.2 TST ...... 51

CHAPTER 7: PARAGENESIS OF THE NOVA AMERICA MEMBER ...... 54 7.1 Stage I: Carbonate deposition ...... 54 7.2 Stage II: Seafloor diagenesis and CFA precipitation ...... 58 7.3 Stage III: Meteoric and shallow-burial diagenesis ...... 67

7.4 Stage IV: Hydrothermal alteration and secondary enrichment of phosphate ..71 7.5 Stage V: Deep-burial diagenesis ...... 78 7.6 Economic phosphorite ...... 84

CHAPTER 8: STABLE ISOTOPE GEOCHEMISTRY ...... 85 8.1 Results ...... 85

CHAPTER 9: DISCUSSION ...... 92 9.1 Depositional model and phosphogenesis ...... 92 9.1.1 Stromatolites, hydrothermal alteration, and economic phosphorite ...... 93 9.2 The Neoproterozoic P-cycle ...... 96

CHAPTER 10: CONCLUSIONS ...... 100

REFERENCES ...... 103

APPENDIX I: X-RAY DIFFRACTION (XRD) DATA ...... 123

APPENDIX II: STABLE ISOTOPE GEOCHEMISTRY ...... 138

TABLE OF FIGURES

LIST OF TABLES Table 5.1 Summary of facies descriptions and environmental interpretations ...... 25

LIST OF FIGURES Figure 2.1 P-cycle and phosphogenesis ...... 5 Figure 2.2 Temporal distribution of phosphorites ...... 7 Figure 2.3 Microbial reactions associated with degradation of organic matter ...... 10 Figure 2.4 Chemical composition of francolite ...... 11

Figure 3.1 Map of São Francisco Craton ...... 14 Figure 3.2 Map of study area, Irecê Basin ...... 15 Figure 3.3 General stratigraphic section of units filling the Irecê Basin ...... 16

Figure 5.1 Facies F1carbonate deposition ...... 27 Figure 5.2 Facies F2 tepee structure and evaporites ...... 29 Figure 5.3 Facies F3 field photographs ...... 31 Figure 5.4 Facies F3 thin section photomicrographs ...... 32 Figure 5.5 Facies F4 Arrecife Ranch stromatolitic reefs ...... 34 Figure 5.6 Facies F4 evaporite pseudomorphs...... 35 Figure 5.7 Facies F5 intertidal columnar stromatolites ...... 37 Figure 5.8 Facies F5 pristine phosphatic laminae ...... 38 Figure 5.9 Facies F6 Gabriel deposits ...... 40

Figure 6.1 Fence diagram showing local lateral facies changes of drill cores ...... 43 Figure 6.2 Regional fence diagram showing regional stratigraphic relationships ...... 44 Figure 6.3 Legend for stratigraphic sections ...... 45 Figure 6.4 Relative sea level curve of the Nova America and Gabriel members ...... 46 Figure 6.5 Composite section of the Nova America and Gabriel members ...... 47 Figure 6.6 Stromatolitic reef of Arrecife Ranch ...... 49 Figure 6.7 Drill core FNC05 stratigraphic column ...... 52 Figure 6.8 Composite stratigraphic section through peritidal limestones ...... 53

Figure 7.1 Paragenetic pathways and mineralization...... 55 Figure 7.2 Simplified paragenetic history ...... 56 Figure 7.3 Stage 1 carbonate deposition, ooids and detrital input ...... 57 Figure 7.4 Stage 2 evaporite pseudomorphs ...... 59 Figure 7.5 Stage 2 pristine phosphorite ...... 61 Figure 7.6 Stage 2 coated phosphate grains ...... 62 Figure 7.7 Stage 2 phosphatic grains ...... 63 Figure 7.8 Stage 2 phosphatic grains ...... 64 Figure 7.9 BSE & EDS images of altered phosphate ...... 65 Figure 7.10 Stage 3 blocky zoned dolomite ...... 68 Figure 7.11 Stage 3 sucrosic and subhedral dolomite ...... 69 Figure 7.12 Stage 4 hydrothermal saddle dolomite ...... 73

Figure 7.13 Stage 4 saddle dolomite and hydrothermal mineralization ...... 74 Figure 7.14 Stage 4 hydrothermal fluorite ...... 75 Figure 7.15 Stage 4 fracture occluded by hydroxylapatite and saddle dolomite ...... 76 Figure 7.16 Stage 4 fractures infilled saddle dolomite ...... 79 Figure 7.17 Stage 5 calcite vein ...... 80 Figure 7.18 Stage 5 calcite vein ...... 81 Figure 7.19 Stage 5 cross-cutting calcite, quartz, and dolomite vein ...... 82 Figure 7.20 Stage 5 stylolites ...... 83

Figure 8.1 δ18O and δ13C results from the Nova America member ...... 86 Figure 8.2 δ18O and δ13C results of hydrothermal veins and surficial calcrete ...... 87 Figure 8.3 Composite cross-plot of all stable C and O from the Nova America member 90

Figure 9.1 Depositional model of the Nova America member ...... 94 Figure 9.2 Phosphogenesis in the Nova America member ...... 97

vii

ABSTRACT

The Neoproterozoic Salitre Formation (ca. 610 Ma) of the Irecê Basin, Brazil, is a

1.2-km-thick succession of peritidal limestone and economic phosphorite that accumulated on an unrimmed epeiric ramp. Lithofacies stacking patterns indicate deposition occurred during a marine transgression punctuated by higher order fluctuations in relative sea level that produced meter-scale, shallowing-upward peritidal cycles. Cycle bases are formed of subtidal, cross-stratified grainstones and hemispheroidal columnar stromatolite reefs. Phosphorite is restricted to the paleocoast where columnar stromatolitic biostromes colonized expansive intertidal flats that prograded over subtidal facies as accommodation filled. Aggradational and then retrogradational stratal stacking of cycles indicates deposition occurred during the transition from lowstand to transgressive conditions.

Phosphogenesis was restricted to the shore because stromatolitic biostromes created the necessary solution and surface chemistries for phosphogenesis. Microbial processes associated with stromatolitic growth actively store, release, and concentrate phosphate to promote authigenic francolite precipitation. Also important were the sealing effects of interbedded, fine-grained storm layers. Energetic subtidal environments where stromatolitic patch reefs developed promoted recycling of P back to seawater, preventing phosphogenesis. Petrographic and stable isotopic data (C, O) indicate that subsequent hydrothermal alteration resulted in pervasive dolomitization and remobilization of P to precipitate secondary phosphatic minerals and produce economic phosphorite. These results suggest that the benthic P-cycle in the Neoproterozoic was more complex than previously surmised and emphasize the multifaceted significance of microbial processes.

viii

ACKNOWLEDGEMENTS

I want to start by thanking my supervisor, Dr. Peir K. Pufahl, for the opportunity to work on an outstanding project at Acadia University and in Brazil. Not only does

Peir's research bring his students to some of the most exotic corners of the globe, his expertise in sedimentology, research, and technical writing has provided me with the tools to become a well-rounded person. My confidence in both my research and presentation skills have been greatly improved thanks to Peir's methods. I am also grateful for the openness and support of his family prior to and during my stay in

Wolfville.

The success of our field work in Bahia, Brazil, would not have been possible without the support and knowledge of CPRM's geologists, especially Maisa Abram and

Antonio Rocha Dourado. Dourado assisted with the field logistics and it was a pleasure to work with a passionate geologist with extensive knowledge of the area as well as a historical enthusiast who filled us in on most of the regions rich history in-between outcrops. Thank you to Domingo, who was more than our driver and always willing to lend a working hand.

Acadia's Earth and Environmental Science faculty were lovely and built on this enjoyable experience. In particular, I would like to thank Pam Frail for preparing thin sections and Haixin Xu for assisting me on the scanning electron microscope.

My family and friends deserve many thanks for their encouragement and support, especially Erika and Jeremy. My parents, Bob and Elaine Delisle, instilled within me a desire to work hard and follow my passion in life and I would not be where I am today without them. Thank you!

CHAPTER 1: INTRODUCTION

Phosphorite is a marine, chemical sedimentary rock rich in phosphorus (P), a bioessential element required for all life processes (Bentor, 1980; Jarvis et al., 1994;

Föllmi, 1996; Pufahl, 2010). By definition, phosphorite contains >18wt% of P2O5 making it an important fertilizer ore (Jarvis et al., 1994; Filippelli, 2011). In addition to its economic significance, phosphorite holds important scientific clues to understanding oceanic, atmospheric and biologic evolution. Changes to the P-cycle that result in phosphorite deposition also limit biological productivity over geologic time scales (Glenn et al., 1994; Föllmi, 1996; Compton et al., 2000) and thus drive biologic evolution

(Brasier and Callow, 2007; Pufahl, 2010; Drummond et al., in press).

Four main episodes of phosphorite accumulation are recognized in the geologic record (Pufahl, 2010; Pufahl and Hiatt, 2012). These phosphogenic events record the evolving chemical composition of Earth's atmosphere and oceans related to episodes of tectonic, climatic, and oceanographic change. The focus of research presented herein is on the second phosphogenic episode, which is late Neoproterozoic to early Paleozoic

(740-410 Ma) and includes the first true phosphorite giants. Phosphorite accumulation during this episode records myriad interrelated events that profoundly influenced the biogeochemical cycling of P. These include the snowball glaciations, breakup of Rodinia and formation of Gondwana, the ventilation of the deep ocean, known as the

Neoproterozoic Oxygenation Event, and the Ediacaran and Cambrian radiations

(Hoffman, 1999; Narbonne and Gehling, 2003; Meert and Lieberman, 2008; Gaucher et al., 2010; Nelson et al., 2010; Papineau, 2010; Shields-Zhou and Och, 2011; Och and

Shields-Zhou, 2012; Pufahl and Hiatt, 2012).

The purpose of this thesis is to add to what is known about the early P-cycle during the beginning of Earth's second phosphogenic episode by investigating phosphorite from the Neoproterozoic Salitre Formation (ca. 610 Ma), Irecê Basin, eastern

Brazil. The Salitre Formation is one of only a few occurrences of early Ediacaran phosphorites (Cook and Shergold, 1986), providing perspective on the Neoproterozoic P- cycle. Phosphorite in the Salitre Formation is associated with stromatolitic peritidal limestones that accumulated during the interglacial between the Marinoan and Gaskiers glaciations. A primary objective is to illuminate the feedbacks that changed the biogeochemical cycling of P to produce phosphorite through this tumultuous time in

Earth history. Understanding secondary enrichment processes that produced economic phosphorite is also an important goal.

The depositional environments and oceanography of the Salitre Formation are interpreted by documenting the sedimentology and regional stacking patterns of lithofacies. The paragenesis is interpreted in the context of this stratigraphic framework to further refine paleoenvironmental interpretations. This information assists in determining seawater and pore water Eh at the time of deposition, diagenesis, and secondary enrichment processes, which when combined with the sedimentology, permit construction of a paleoenvironmental and phosphogenic model for the Salitre Formation.

CHAPTER 2: BACKGROUND

Phosphorus is an important bio-essential element that regulates energy transfer in cells and forms the backbone of DNA (Föllmi, 1996; Ruttenberg, 2003; Paytan and

McLaughlin, 2007). Because P is ultimately derived through the weathering of continental rocks, its availability regulates biological productivity on geological timescales (Jarvis et al., 1994; Föllmi, 1996). Variations in the biogeochemical cycling of P through geologic time is linked to changes in tectonics, climate, biologic productivity, and ocean chemistry.

2.1 Neoproterozoic climate and P-cycle

The Neoproterozoic was a time of profound environmental change. Important events include the breakup and amalgamation of Rodinia and Gondwana, respectively, the Snowball glaciations, the oxygenation of the deep ocean, changes in the biogeochemical cycling of P, and the evolution of multicellular life (Gaucher et al.,

2010). Of these, the Snowball glaciations and the Neoproterozoic Oxygenation Event

(NOE), which led to deep ocean ventilation at ca. 580 Ma, were probably the most important in regulating the late Precambrian P-cycle (this study). Evidence for the

Snowball glaciations occurs on every continent except Antarctica (Och and Shields-

Zhou, 2012). In central Brazil, diamictites of the Bebedouro Formation from the Irecê

Basin and the Jequitaí Formation from the São Francisco Basin are interpreted to be

Marinoan glacial deposits (ca. 635-600 Ma; Gaucher et al., 2010; Caxito et al., 2012).

Oxic chemical weathering of the continents between Snowball glaciations is interpreted to have increased the delivery of P to the oceans, preconditioning seawater for

3 the accumulation of phosphorite (Pufahl and Hiatt, 2012; Papineau et al., 2013). The flux of P was likely orders of magnitude higher when the Snowball glaciations receded because chemical processes easily degraded the newly exposed, vast, mechanically weathered glacial landscapes. These P pulses are thought to have stimulated primary productivity in the global ocean, driving deep ocean ventilation, referred to as the NOE, at ca. 580 Ma (Och and Shields-Zhou, 2012; Papineau et al., 2013). Increasing 87Sr/86Sr ratios through the Neoproterozoic support this interpretation because they signal enhanced continental weathering and P delivery to the Cryogenian and Ediacaran oceans

(Kaufman and Knoll, 1995; Misi and Veizer, 1998; Misi et al., 2014, n.d.).

Chemical weathering is the main mechanism that liberates P from continental rocks. P is transported to oceans via riverine input and/or aeolian sources as particulate matter, typically igneous apatite, or as a solute (Fig. 2.1; Föllmi, 1996; Compton et al.,

2000; Filippelli, 2008, 2011; Pufahl, 2010). Bioavailable dissolved P is quickly depleted in shallow environments through its incorporation by photosynthetic organisms. P that enters the system adsorbed to Fe-(oxyhydr)oxides and clays may be released by changes in oceanic pH and Eh, and subsequently become bioavailable (Benitez-Nelson, 2000;

Compton et al., 2000).

2.2 Phosphorites and phosphogenesis

Phosphorites are P-rich marine, sedimentary deposits that by definition contain >

18 wt.% P2O5. In some cases they possess nearly 40 wt.% P2O5, making them an important fertilizer ore (Jarvis et al., 1994; Pufahl, 2010; Filippelli, 2011). True phosphorite giants are a Phanerozoic phenomenon; only one phosphogenic event is

Figure 2.1 Simplified P-cycle (a) and phosphogenesis (b). Upwelling is related to an increase in primary productivity, which leads to an increase in organic matter (C) burial. Organic matter degradation releases phosphate into pore waters. Phosphate is shown as 3- PO4 . Non-upwelling environments are dominated by Fe redox-pumping in which saturation of pore water phosphate is maintained through cyclic precipitation of Fe- (oxyhydr)oxide and its dissolution below the Fe-redox boundary. (Modified from Jarvis et al., 1994; Compton et al., 2000; Nelson et al., 2010; Pufahl, 2010)

5 known to occur entirely in the Precambrian (Fig. 2.2). Although not well understood,

Precambrian phosphorites seem to be linked to changes in the biogeochemical cycling of

P related to the weathering of postglacial landscapes and ocean oxygenation (Papineau,

2010; Pufahl, 2010; Pufahl and Hiatt, 2012; Papineau et al., 2013).

Four types of phosphatic sedimentary systems are recognized: insular phosphorite, seamount phosphorite, continental margin phosphorite, and epeiric sea phosphorite (Glenn et al., 1994). Insular phosphorite forms on carbonate atolls when guano from seabirds alters limestone during meteoric diagenesis. Seamount phosphorite is thought to simply be submerged insular phosphorite. Continental margin and epeiric sea phosphorites are generally associated with coastal upwelling, although the current regimes pumping and delivering phosphate across the shelf are different (Glenn et al.,

1994; Pufahl, 2010). Although phosphogenesis is restricted to the outer shelf of modern continental margins, phosphorites can form across the entire platform in epeiric sea environments due to evaporation-driven lagoonal circulation. Evaporation in nearshore settings drives shoreward flow of nutrient-rich surface water and the outflow of saline water at depth. This lagoonal circulation transports dissolved phosphate from zones of active coastal upwelling to nearshore environments, promoting phosphogenesis across the entire platform (Pufahl, 2010).

In highly productive settings such as upwelling environments, P is biologically fixed in the surface ocean through photosynthesis. Upon death, phytoplankton rain to the seafloor, exporting this stored P to the sediment. As sedimentary organic matter accumulates, an oxygen minimum zone forms and impinges on the seafloor. Bacterial degradation in this zone saturates pore water with phosphate to precipitate carbonate

Figure 2.2 Temporal distribution of phosphorites, purple; modified from Pufahl and Hiatt (2012). Major events shown: GOE - Great Oxidation Event; BB - Boring Billion; NOE - Neoproterozoic Oxidation Event; CE - Cambrian Explosion; Glaciations numbers 1 - 7 (4: Neoproterozoic Snowball glaciations).

7 fluorapatite (CFA), a highly substituted sedimentary apatite (Ca10-a-bNaaMgb(PO4)6- x(CO3)x-y-z(CO3•F)y(SO4)zF2; Jarvis and Jarvis, 1985; Jarvis et al., 1994; Föllmi, 1996;

Pufahl, 2010).

Sedimentary organic matter degrades through a series of microbially mediated redox reactions (Fig. 2.3), all of which release phosphate to saturate pore waters. CFA precipitation occurs just below the sediment-water interface in association with the microbial reduction of nitrate, Mn-oxides, Fe-oxides, and sulfate (Pufahl, 2010).

Phosphogenesis is restricted to within 5-20 cm of the sediment-water interface due to the diffusion of seawater-derived fluorine (F). Thus, phosphogenesis is not a redox- controlled authigenic process, but depends only on the concentration of pore water phosphate and the availability of F derived from overlying seawater (Fig. 2.1; Jarvis et al., 1994).

In settings without prominent upwelling, phosphate concentrations in pore water are regulated primarily by Fe-redox pumping (Fig. 2.1; Heggie et al., 1990; Jarvis et al.,

1994). Fe-redox pumping is a cyclic mechanism that saturates pore water when phosphate adsorbed onto Fe-(oxyhydr)oxides, precipitated in the water column, dissolve below the Fe redox interface during burial. Released phosphate is prevented from diffusing out of the sediment by re-adsorbing to freshly precipitated Fe-(oxyhydr)oxides at the oxygenated seafloor.

Phosphatic lithofacies resulting from phosphogenesis through either the microbial degradation of organic matter or Fe-redox pumping are generally laminated and unbioturbated. These pristine phosphorites commonly contain phosphatic laminae and abundant CFA nodules formed in-situ (Föllmi, 1996). Some nodules are coated with a

8 variety of redox-sensitive authigenic minerals (e.g. pyrite, chamosite, and/or barite) that record the fluctuation of redox boundaries through the sediment (Pufahl and Grimm,

2003). Such variations in redox potential have been shown to reflect changes in the delivery of sedimentary organic matter to the seafloor.

Granular, reworked phosphorite is formed by hydraulically concentrating nodules that originally formed in pristine facies by tides, fair-weather waves, and storm currents.

The resulting intraclastic deposit often reflects reduced or net-negative sedimentation rates during episodes of stratigraphic condensation. Such conditions produce beds with a complex paragenesis that can record multiple episodes of phosphogenesis, reworking, and reburial into the zone of phosphogenesis (Glenn et al., 1994; Pufahl et al., 2003).

2.3 Geochemistry of phosphorites and associated sedimentary rocks

CFA can easily incorporate trace elements and undergo isotopic fractionation in

3- 2- both the PO4 and CO3 structural sites (Fig. 2.4; Jarvis et al., 1994; Hiatt and Budd,

2001). Therefore, it is commonly analyzed to reconstruct paleoenvironmental conditions, as well as the degree of diagenesis and metamorphism (McArthur et al., 1986;

McArthur and Herczeg, 1990; Jarvis et al., 1994). Alteration is most easily assessed

2- using the stable isotopic composition of C and O from the CO3 , because it can be directly compared to and interpreted with similar data from associated limestones and dolostones. This relationship is especially useful when CFA is present only in a very few sedimentary facies and/or is difficult to extract for isotopic analysis, which is the case for

CFA in this study.

Figure 2.3 Microbial reactions involved in organic matter degradation in order of decreasing energy yield and corresponding isotopic signatures of pore fluid (Glenn et al., 2000; Albarède, 2009).

Figure 2.4 Chemical composition of francolite, a highly substituted carbonate fluorapatite showing isotopic substitution of carbon and oxygen stable isotopes in both the phosphate, carbonate, and sulfate sites (Jarvis et al., 1994).

The chemostratigraphy of the Salitre Formation has been used to correlate between sedimentary basins and infer processes of Neoproterozoic environmental change

(Caxito et al., 2012). Remarkably, it is unclear whether the Salitre Formation truly contains a primary depositional signal or has been significantly altered. Stable oxygen

(18O) and carbon (13O) isotopic data from all carbonate facies, hydrothermal calcite veins, and pedogenic calcite complement petrographic analysis of sedimentary facies in this study to assess alteration and understand secondary, fluid-related ore-forming processes (Bathurst, 1975; Choquette and James, 1990; James and Choquette, 1990a;

Jarvis et al., 1994; Shen et al., 2000).

CHAPTER 3: GENERAL GEOLOGY

Phosphorites accumulated in the Neoproterozoic Irecê Basin of the state of Bahia

(Figs. 3.1, 3.2). Phosphatic and associated limestones belong to the Una Group (Sial et al., 2010; Teixeira et al., 2010; Guimarães et al., 2011; Misi et al., 2011; Alkmim and

Martins-Neto, 2012; Caxito et al., 2012). The Una Group (Fig. 3.3) is composed of two megasequences representing glacial and interglacial sedimentation (Sial et al., 2010).

Interglacial carbonate deposits belong to the phosphatic Salitre Formation (Misi and

Kyle, 1994), the focus of this study. The Salitre Formation is economically significant because of the economic phosphorite it contains. These deposits are currently mined by

Galvani Mineração in the Irecê Basin near the town of Irecê (Misi and Kyle, 1994).

3.1 Tectonic Setting

The Irecê Basin is a sub-basin of the São Franciscan Basin (this study), which is an expansive Neoproterozoic epeiric basin that occupies much of central Brazil (Fig. 3.1).

Deposition of the Salitre Formation occurred directly on the São Francisco Craton

(Gaucher et al., 2010). The configuration of the São Francisco Craton and overlying strata record the rifting, drifting, and collision related to the dispersal of Rodinia and its later incorporation into Western Gondwana (Cruz and Alkmim, 2006; Teixeira et al.,

2007).

The São Francisco Craton is bounded on all sides by fold-thrust belts of the

Neoproterozoic Brasiliano Orogen (Fig. 3.1), which includes the Brasília, Rio Preto,

Riacho do Pontal, Sergipano, and Araçuaí orogenies (Cruz and Alkmim, 2006; Alkmim and Martins-Neto, 2012), which represent the final assembly of Gondwana by ca. 550 -

Figure 3.1 Map of the São Francisco Craton (SFC) and configuration of the Irecê Basin (Modified after Alkmim et al. 2001; Cruz and Alkmim 2006; Alkmim and Martins-Neto 2012). Orange: Archean (basement); Blue: Paleo/Mesoproterozoic cover (Espinhaço sequences and Chapada Diamantina Group); Yellow: Neoproterozoic cover (Macaúbas and Bambuí sequences, including Una Group deposits); White: Phanerozoic cover; Red lines: Paramirim aulacogen boundary; Green lines: São Francisco Basin outline; IB: Irecê Basin; CD: Chapada Diamantina; PC: Paramirim corridor; SB: Salitre Basin; UB: Utinga Basin; TB: Ituaçu Basin; Study area is outlined by purple box.

Figure 3.2 Study area, the Irecê Basin, with prominent field and drill core locations. The dotted red line refers to the regional correlation of stratigraphic sections Fig. 6.2. See Chapter 4 Methods for specific details about each location including coordinates and section thickness.

Figure 3.3 Modified stratigraphic section of the Salitre Formation, Una Group of the Irecê Basin and modified correlation with the Bambuí Group of the São Francisco Basin. Carbonate is colored according to lithological description, i.e. pink dolostones and black limestone, non-specific dolostone and limestones are colored green (Souza et al., 1993; Misi and Veizer, 1998; Drummond, 2014).

520 Ma. The Irecê Basin is interpreted as an aulocogen that first developed during

Mesoproterozoic rifting of the Paleoproterozoic supercontinent Columbia (Alkmim et al.,

2001; Cruz and Alkmim, 2006; Teixeira et al., 2007; Alkmim and Martins-Neto, 2012).

Initial sedimentary fill of the Irecê Basin consists of sandstone of the

Mesoproterozoic Chapada Diamantina Group (Figs. 3.1, 3.2, 3.3; Alkmim and Martins-

Neto 2012). Neoproterozoic siliciclastic, carbonate, and phosphorite rocks of the Una

Group rest unconformably over the Chapada Diamantina Group (Fig. 3.3). The Una

Group is subdivided into the Bebedouro Formation and the phosphatic Salitre Formation, which are interpreted to correlate to the Macaúbas and Bambuí groups, respectively, of the São Francisco Basin (Misi and Veizer, 1998; Guimarães et al., 2011; Misi et al.,

2011).

The Bebedouro Formation, which is correlated with the Jequitaí Formation, is interpreted to record the deposition of Marinoan glacial diamictite during rift-related reactivation of the aulocogen associated with the breakup of Rodinia at ca. 825 and 740

Ma (Misi et al., 2005; Cruz and Alkmim, 2006; Li et al., 2008; Rodrigues, 2008;

Guimarães et al., 2011). The Salitre Formation unconformably overlies the Bebedouro

Formation and Morro do Chapeu Formation of the Chapada Diamantina, and accumulated between ca. 670 and 600 Ma (Misi and Veizer, 1998; Cruz and Alkmim,

2006; Rino et al., 2008; Rodrigues, 2008; Alkmim and Martins-Neto, 2012). It marks the final transgressive pulse in the Irecê Basin during the onset of the Brasiliano Orogen

(Cruz and Alkmim, 2006; Rino et al., 2008; Rodrigues, 2008; Alkmim and Martins-Neto,

2012).

The Brasiliano Orogen records the suturing of the São Francisco Craton and to form Western Gondwana between ca. 750 and 500 Ma (Alkmim et al., 2001; Condie,

2002). The Southern Brasília belt resulted from the collision of the São Francisco and

Rio de la Plata cratons at ca. 750 Ma and the subsequent collision with the Amazonia

Craton at ca. 640 Ma (Alkmim et al., 2001). Continued collision with Amazonia produced the Northern Brasília mobile belt. The Araçuaí and West Congo belt formed penecontemporaneously to close the Adamastor Ocean. The closure of the Brazilide

Ocean between ca. 630 and 500 Ma created the Riacho do Pontal and Sergipano belts of the Borborema Province (Alkmim et al. 2001).

3.2 Una Group

Collectively, the Bebedouro and Salitre formations of the Una Group are ~1.2 km thick (Misi and Veizer 1998; Misi et al. 2005; Sial et al. 2010). The age of the Una

Group is not well constrained because of the absence of interbedded volcanic deposits.

Detrital zircons from Bebedouro diamictites yielded U-Pb SHRIMP ages between 950

Ma (Misi et al. 2005) and 880 Ma (Rodrigues, 2008), which indicate a maximum depositional age of ca. 900 Ma. Diamictites of the Bebedouro Formation are correlated to the diamictites of the Jequitaí Formation of the São Francisco Basin, a Marinoan glacial deposit, based on field relationship and 87Sr/86Sr (0.7075-0.7077) values of the overlying carbonate deposits of the Salitre Formation and the Sete Lagoas Formation, respectively (Guimarães et al., 2011; Caxito et al., 2012). A detrital zircon age of ca. 610

Ma from phosphatic siltstones (Rodrigues, 2008; Pedrosa-Soares et al., 2011; Caxito et

18 al., 2012) of the Sete Lagoas Formation suggests that these basal diamictites are related to the Marinoan Snowball glaciation.

Carbonates, phosphorites, and siliciclastics of the Salitre Formation and related facies of the Sete Lagoas Formation of the São Francisco Basin, are interpreted to have accumulated on an extensive Neoproterozoic ramp in an epeiric sea (Misi and Kyle,

1994; Misi and Veizer, 1998; Drummond et al., in press). The Salitre Formation has been divided into four mappable units by the Geological Survey of Brazil (Sampaio et al.,

2001). These informal stratigraphic members include the basal Nova America member, a phosphatic peritidal limestone succession and the focus of this study, the Gabriel member, a mid-ramp carbonate succession, the Jussara member, a highly deformed, black, thinly bedded limestone, and the Irecê member, a phosphatic siltstone and limestone unit that is correlative to the phosphatic siltstones of the Sete Lagoas

Formation in the western Bambuí Basin (this study; Drummond et al., in press). The Una

Group is unconformably overlain by Phanerozoic sediment.

The Nova America member is composed of variably phosphatic peritidal carbonate cycles. Stacking patterns indicate that their deposition occurred during a marine transgression punctuated by smaller-scale oscillations that produced aggradational parasequences (this study). The overlying Gabriel member is composed of deeper- marine facies and is interpreted to represent continued flooding from lowstand to transgressive conditions. Economic phosphorites are associated only with columnar, intertidal stromatolites, and are characterized by P2O5 concentrations of > 20 wt.% (Misi and Kyle, 1994; Kyle and Misi, 1997; this study).

CHAPTER 4: METHODS

4.1 Field Methods

Fieldwork was conducted in Bahia, Brazil with funding and logistical support from Companhia de Pesquisa de Recursos Minerais (CPRM; the Brazilian Geological

Survey) and Galvani Mineração (Galvani Minerals Ltd.). The study area was approximately 125 by 150 km. Outcrop and drill cores were described and sampled to identify lithofacies and construct a sequence stratigraphic framework. Samples (n = 144) were collected from all lithofacies for petrographic and geochemical analysis.

Drill cores were described bed-by-bed with emphasis on describing sedimentary structures, the nature of contacts, and alteration textures to understand the vertical stacking of lithofacies and secondary fluid-related processes. Nine drill-cores, from an area approximately 1.5 by 0.3 km, were logged at Galvani Mineração's Unidade de

Mineração de Irecê mine near the town of Irecê (Fig. 3.2): FNC-01 (~40 m; Lat. -

11.35313, Long. -41.77166), FNC-02 (~45 m; Lat. -11.35327, Long. -41.77075), FNC-05

(~60.75 m; Lat. -11.35356, Long. -41.76986), FNC-08 (~35 m; Lat. -11.34527, Long. -

41.76842), FNC-10 (~35 m; Lat. -11.34346, Long. -41.76874), FNC-11 (~35 m; Lat. -

11.34255, Long. -41.76885), FNC-13 (~35 m; Lat. -11.34075, Long. -41.76861), FNC-18

(~30 m; Lat. -11.34074, Long. -41.76951), and FNC-19 (~30 m; Lat. -11.34164, Long. -

41.76971). The longest core is FNC-05 and penetrates the Nova America member to a depth of 60.75 m.

Outcrops were described to augment drill core descriptions and understand lateral facies changes. Forty-five outcrops were visited and seven detailed stratigraphic sections

20 were logged based on these field sections: the Achado Section (~44 m; Lat. -11.32987,

Long. -41.77854), Santa Clara Mine Section (~12 m; Lat. -11.4457, Long. -41.41224),

Galvani Mine Section (~15.5 m; Lat. -11.31623, Long. -41.80119), Villa Brejão Section

(~9 m; Lat. -11.28253, Long. -41.07118), Arrecife Farm Section (~4 m; Lat. -11.10220,

Long. -41.02877), Salvador CPRM Section (~26 m; Lat. -10.95878, Long. -41.42479), and Gabriel Section (~16 m; Lat. -11.67850626, Long. -41.76301545). The Villa Brejão

Section contains the contact between the Bebedouro Formation and the Nova America member.

4.2 Laboratory Methods

Analysis of 41 uncovered thin sections using standard petrographic techniques, transmitted-light, cathodoluminescent (CL), and scanning electron microscopy (SEM) was employed to identify and interpret mineralogy and paragenetic relationships. An abundance index of rare (<5%), uncommon (5-25%), common (25-50%), and abundant

(>50%) was applied to minerals and sedimentary structures.

Of the 41 thin sections, two sets of 20 thin sections were polished; one for staining with alizarin-red S and potassium ferricyanide solution (Dickson, 1966), and one for CL analysis. CL and SEM imaging were performed at the Acadia Centre for

Microstructural Analysis (ACMA) lab using a JEOL JSM-5900 LV SEM with a

Princeton Gamm-Tech IMIX-PC EDS detector for the chemical analysis of minerals.

Stable isotopic analysis of 13C and 18O was conducted on 85 samples in the

Queen’s Facility for Isotopic Research (QFIR). Analyses of six Tertiary calcretes that developed on top of the Nova America member provide a baseline for meteoric

21 diagenesis. Five hydrothermal calcite veins were also analyzed as benchmarks for high- temperature alteration processes.

Samples of sediment, calcrete, and vein calcite were powdered to react 0.5 mg with 100% anhydrous phosphoric acid at 72ºC . The evolved CO2 was analyzed on a

Thermo-Finnigan Delta XP Plus continuous flow stable-isotope-ratio mass spectrometer.

The fractionation factors used are those of O’Neil et al. (1969) for oxygen and Deines et al. (1974) for carbon in the system of calcite and water.

Although it is challenging to determine precise fractionation factors for dolomite because it is difficult to synthesize, they differ only slightly from those of calcite (Tucker and Wright, 1990; Warren, 2000). Researchers have overcome this hurdle by analyzing coexisting, associated calcite and dolomite cements (Humphrey, 2000). The oxygen isotope fractionation of dolomite formation compared with calcite has been predicted be

2 to 6‰ heavier at 25°C (Humphrey, 2000; Schmidt et al., 2005) and an averaged equilibrium value for the difference in dolomite and calcite has been calculated by Land

(1980):

(at 25°C)

Calcite samples (FNC01-7.30, FNC02-48.9, FNC02-49.7, FNC02-49.71, and

FNC11-19.6) have been equilibrated with dolomite samples based on this average.

However, because of the mix of calcite and dolomite, it is possible that the value of comparison lies between the measured and corrected values.

Stable isotopic results are reported in delta notation relative to the reference standard of the Peedee belemnite (V-PDB; Craig, 1957):

= [(Rsample/Rstandard) – 1] x 1000

22 where R represents 13C/12C or 18O/16O. Replicate analyses indicate a reproducibility of

±0.2 ‰ for both.

CHAPTER 5: LITHOFACIES

Six pervasively dolomitized lithofacies have been identified from the Nova

America (F1, F2, F3, F4, and F5) and Gabriel (F6) units, which make up informal members of the Salitre Formation. Lithofacies associations suggest cyclic accumulation in peritidal to shallow marine environments (Table 5.1). Lithofacies were defined based on lithological character as well as physical and biologic sedimentary structures.

5.1 Facies 1: Tabular bedded intraclast-rich grainstone

Facies 1 (F1) is a tan to gray moderately-sorted, intraclast-rich dolomite grainstone that is interbedded with Facies F2 (Fig. 5.1). Thin to thickly bedded tabular beds have sharp, erosive bases accentuated by stylolites (Fig. 5.1A). Coarse silt- to fine sand-sized intraclasts have been recrystallized to dolomite, producing a relict texture.

Where intraclasts are not completely recrystallized, they are elongate, well-rounded grains of micrite. Intraclasts are in a matrix of xenotopic, finely-crystalline to medium- crystalline mosaic of dolomite crystals (Fig. 5.1C,D; Gregg and Sibley 1984; Tucker and

Wright 1990).

Interpretation: F1 is interpreted to have accumulated in a shallow, subtidal environment.

The abundance of intraclasts and the tabular, erosive nature of beds suggest storm reworking, transport, and redeposition in a shallow subtidal environment (Pratt and

James, 1986; Wright and Burchette, 1996; Boggs, 2006). Intraclasts were derived from reworking an indurated seafloor composed of lime mud (Hardie and Garrett, 1977;

Table 5.1 Summary of facies descriptions and environmental interpretations.

Facies Lithology/Sedimentary Structures Interpretation F1 Tabular bedded Moderate- to well-sorted, intraclastic Shallow, subtidal intraclast-rich grainstone; tabular beds, thin to thickly environment grainstone bedded, commonly interbedded with F2; influenced by frequent Tan to gray with dark gray, white, and storm reworking (Pratt black coarse silt- to fine sand-sized, and James, 1986; irregular, well-rounded intraclasts, of relict Bosence and Wilson, granular texture; Recrystallization to 2003; Jones, 2010; dolomite produced a xenotopic, finely- to Pratt, 2010) medium-crystalline mosaic cement.

F2 Cross-stratified Gray with white and dark gray allochems; Shallow, storm- and intraclastic thick bed sets of ripple and herringbone tidal-influenced grainstone cross-stratification, commonly associated subtidal environment with F4; Contains well-rounded, coarse (Hardie and Ginsburg, silt- to very fine sand-sized intraclasts, 1977; Reineck and subangular to subrounded silt-sized detrital Singh, 1980; Dill et quartz and orthoclase, and rare very fine al., 1986) sand-sized ooids; Intraclasts are composed of very fine crystalline dolomite, grumuleuse texture.

F3 Flaser- to Interbedded flaser- to lenticular-bedded Arid, storm- lenticular-bedded dolomite packstone, with a sharp influenced, intertidal packstone undulated base; Red-brown, tan, and gray environment thick, wavy laminae; uncommon to (Bathurst, 1975; common crinkled microbial laminae; Hardie and Garrett, Mudcracks, tepee structures, fenestral 1977; Hardie and porosity, and evaporite pseudomorphs are Ginsburg, 1977; common; Ooids and grumuleuse textured Tucker and Wright, intraclasts are the most common grain 1990) components with rare detrital quartz. Uncommonly aggregates of ooids and intraclasts form grapestones.

F4 Hemispheroidal Macro-bioherms of coalesced columnar Shallow, tide- columnar stromatolites; bioherms 1 to 2 m tall and 1 dominated subtidal stromatolites to 1.5 m wide, consist of juxtaposed environment, with steeply convex columnar stromatolites, frequent storm 0.3-m-wide, 0.5 m tall; Bioherms maintain reworking (Dill et al., consistent width for 0.5 m and widen at the 1986; Tucker and top, plume-like structures; Bioherms Wright, 1990; Walter elongated in plan-view, aligned west-east; et al., 1992; Microbial laminae consists of microbial Srivastava and Rocha, micrite and fine crystalline dolomite 1999; Reid et al., cement with rare to uncommon, fine sand- 2000; Altermann, sized ooids and intraclasts, and rare detrital 2004) quartz.

F5 Tabular phosphatic Thin tabular biostromes of parallel, Low energy, intertidal columnar bifurcate-branching and anastomosed, environment (Logan stromatolites columnar stromatolites; small columnar et al., 1964; Hofmann, stromatolites occur as tabular beds with 1973) dolomite filling intercolumnar spaces; Uncommon evaporite pseudomorphs now composed of quartz or calcite present; Only facies to contain appreciable fluorapatite, P2O5 up to 30.5 wt.%; Phosphatic minerals present include fluorapatite, strontian fluorapatite, hydroxylapatite, and apatite, present in pristine non-granular deposits associated with microbial laminae as well as secondary phosphatic well-rounded granules; Rare quartz and orthoclase present.

F6 Hummocky cross- Hummocky cross-stratified (HCS) Storm-dominated, stratified grainstone and intercalate thinly bedded mid-ramp marine grainstone lime mudstones; rare intraclastic lags mark environment between the bottom of HCS bedsets. Microbial fair-weather and storm laminae are common in intervals of wave base (Dott and laminated mudstone; microcrystalline Bourgeois, 1982; dolomite comprise samples of Gabriel with Duke, 1985; Dumas disseminated pyrite, organic matter, and and Arnott, 2006) hematite; rare intraclasts of grumuleuse texture are present within mudstone beds, that have organic matter-rich stylocumulate.

Figure 5.1 Facies 1; A: Santa Clara Mine, showing thick intraclastic grainstone deposits; B: Drill core of F1 showing massive grainstone with intraclasts; C, D: Photomicrographs of sample FNC05-46.60, characteristic of F1, showing complete recrystallization to dolomite, plane-polarized view and cross-polarized view, respectively.

Hardie and Ginsburg, 1977; Tucker and Wright, 1990; Bosence and Wilson, 2003; Jones,

2010).

5.2 Facies 2: Cross-stratified intraclastic grainstone

Facies 2 (F2) is a cross-stratified intraclastic grainstone (Fig. 5.2). It is gray with white and dark gray allochems. Thick bed sets up to 1 m thick are rippled and herringbone cross-stratified (Fig. 5.2A,B,C). F2 is commonly associated with Facies F4, accumulating between stromatolitic patch reefs (Fig. 5.2A).

F2 consists of rare to common, well-rounded, coarse silt- to very fine sand-sized intraclasts, subangular to subrounded silt-sized detrital quartz and orthoclase grains, and rare, very fine sand-sized ooids (Fig. 5.2D). Abundant intraclasts are sub-spherical to spherical and are now composed of microcrystalline dolomite with differentially preserved grumuleuse texture. Rare recrystallized ooids are defined by a well-developed micrite envelope. Grains are cemented by subhedral dolomite cement.

Interpretation: F2 is interpreted to have accumulated in a tide and storm-wave swept shallow subtidal environment. As in F1 intraclasts are interpreted to have been created by winnowing and reworking of an indurated seafloor (Hardie and Garrett, 1977; Jones,

2010). Intraclasts with grumuleuse texture may represent weakly calcified microbial mats, composed of sub-globular cell aggregates, which were reworked by tides and storms (Bathurst, 1975; Lemon, 2000; Kazmierczak et al., 2004). The presence of ooids and herringbone cross-stratification suggests the precipitation of ooids was facilitated by

Figure 5.2 Facies 2 A, B: Arrecife Farm outcrop displaying F2 and its association with Facies 4 and herringbone cross-stratification, respectively. Hammer is 30 cm; C: Drill core showing truncated surface indicative of cross-stratification (white arrow). Drill core is upright, with the base at the bottom; and D: photomicrograph of sample FNC05-12.75 showing intraclasts and detrital quartz grains.

29 turbulence created through the bimodal oscillation of tidal currents (Reineck and Singh,

1980; Pratt and James, 1986; Boggs, 2006; Dalrymple, 2010).

5.3 Facies 3: Flaser to lenticular-bedded packstone

Facies 3 (F3) is a reddish brown and tan flaser- to lenticular-bedded packstone

(Fig. 5.3A, B). Mudcracks (Fig. 5.3B), tepee structures (Fig. 5.3C) evaporite pseudomorphs, and fossil pustular microbial laminae are common. Rare flat-pebble conglomerate beds have sharp, erosive bases that scour into underlying sediment (Fig.

5.3D). Ripple-cross-laminated packstones are composed of intraclasts similar to those in

F1 and F2. Fenestral pores are completely occluded with saddle dolomite and fluorite

(Fig. 5.4A). Ooids are present in thin, sharp-based tabular beds that are intercalated with these rippled layers. Ooids are tangential and completely recrystallized and altered to dolomite (Fig. 5.4B,C,D). Cortical layers are defined only by their micrite envelopes

(Fig. 5.4B). Rarely, ooids contain a collapsed nucleus, which is composed of grumuleuse texture surrounded by dolomite cement (Fig. 5.4C). Some ooids and intraclasts form grapestone aggregates (Fig. 5.4D). Calcite pseudomorphs after acicular gypsum are common.

Interpretation: F3 is interpreted to have accumulated on an intertidal flat along an arid, storm-influenced coast. Desiccation cracks, tepee structures, and evaporites indicate episodes of extreme aridity and subaerial exposure (Reineck and Singh, 1980; Tucker and

Wright, 1990; Dyer, 1998; Boggs, 2006; Pratt, 2010) similar to the Trucial Coast microbial flats in the Persian Gulf (Bathurst, 1975). The presence of flaser and

Figure 5.3 Facies 3 macro-fabrics; A: Flaser bedding with crinkled microbial laminae (white arrow) from the Santa Clara Mine outcrop; B: Flaser bedding with mudcracks on the surface from the Santa Clara Mine outcrop; C: Tepee structure, from the Achado section, filled by an evaporite precipitate originally gypsum or anhydrite, which has been replaced by calcite. Note the fractured and buckled sediment surround the cream-colored precipitate (Assereto and Kendall, 1977); D: Gravel-sized, flat, imbricated intraclasts associated with F3 from the Arrecife Ranch section.

Figure 5.4 Facies 3 micro-fabrics; A: Photograph of drill core FNC05, F3 displaying fenestral porosity (white arrow), infilled by saddle dolomite; B: Photomicrograph of sample FNC05-39.30 displaying tangential ooids (white arrow) and ooids that have undergone internal dissolution and re-precipitation of dolomite, with rare fabric retention, under plane-polarized light; C: Photomicrograph of sample FNC02-48.88 showing recrystallized ooids and an ooid with a collapsed nucleus (white arrow); D: Photomicrograph of sample FNC05-39.30 showing ooids and intraclasts that have been cemented by microbial communities, forming a grapestone.

32 lenticular bedding and fenestral porosity are also key indicators of mudflat deposition

(Tucker and Wright, 1990; Dyer, 1998; Pratt, 2010). The flat-pebble conglomerates and thin tabular oolitic beds are interpreted as storm layers. Conglomeratic beds reflect the reworking of tidal flat deposits during large storms, whereas the thinner grainy beds record less severe storms that blanketed the tidal flat with shallow subtidal sediment.

Microbial laminae and grumuleuse texture in some intraclasts likely reflect the colonization of tidal pools by cyanobacterial mats (Bathurst, 1975; Knoll et al., 1991;

Reid et al., 2003). These communities helped to bind the sediment to produce a well- cemented seafloor that when swept by tides and storms formed intraclasts (Bathurst,

1975; Lemon, 2000).

5.4 Facies 4: Hemispheroidal columnar stromatolites

Columnar stromatolitic reefs characterize Facies 4 (F4; Fig.5.5). Individual hemispheroidal heads coalesced to form reefs that are 2 to 3 m tall and 5 to 8 m wide.

Juxtaposed coalesced and steeply convex columnar stromatolites (Walter et al., 1992;

Altermann, 2004) gradually widen from their base to produce reefs that expand upwards

(Figs. 5.5A,C). Evaporite pseudomorphs composed of blocky calcite are common (Fig.

5.6). Microbial laminae have been completely replaced by microcrystalline dolomite.

Some laminae are accentuated by the presence of rare ooids, intraclasts, and coarse silt- sized detrital quartz grains. In plan view, bioherms are elongated (Fig. 5.5B) in an east- west orientation and in vertical section, intercalated with cross-stratified grainstone of

Facies F2.

Figure 5.5 Facies 4, columnar stromatolite patch reefs from Arrecife Farm; A, C: Cross- section of patch reefs, showing coalesced columnar stromatolites; B: Plan view showing elongation of the patch reefs.

Figure 5.6 Calcite-replaced evaporite, probably originally gypsum or anhydrite. Evaporite nodules (white arrows) are associated with columnar stromatolite patch reefs (F4) from Arrecife Ranch.

Interpretation: Facies 4 is interpreted to represent the development of photosynthetic cyanobacterial stromatolitic patch reefs in a subtidal environment that was well within the photic zone. The reefs' plan morphology is similar to that of other Paleoproterozoic

(Hofmann, 1973; Walter et al., 1992; Lemon, 2000; Altermann, 2004) and modern

Bahamian subtidal stromatolitic reefs (Dill et al., 1986; Reid et al., 1995, 2003;

Srivastava and Rocha, 1999; Bowlin et al., 2012). Around the Exuma Islands, Bahamas, similar subtidal stromatolitic reefs are present in 7 to 8 m deep tidal channels and consist of coalesced heads that produce elongate bioherms parallel to tidal flow (Dill et al., 1986;

Tucker and Wright, 1990). As in the Nova America member, stromatolites are enveloped by carbonate sand dunes, indicating high-energy, tidal-current-shaped patch reefs (Dill et al., 1986; Pratt and James, 1986; Tucker and Wright, 1990; Walter et al., 1992; Sami and

James, 1994). The evaporites associated with these stromatolites indicate that conditions were conducive to evaporite formation even in shallow subtidal environments (Kendall,

2010).

5.5 Facies 5: Tabular phosphatic columnar stromatolites

Facies 5 (F5) consists of tabular biostromes composed of parallel, bifurcate- branching and anastomosed, columnar stromatolites (Fig. 5.7; Walter et al. 1992;

Altermann 2004). Small columnar stromatolites (~ 8 cm tall by 2 cm wide) form tabular beds with microcrystalline dolomite filling intercolumnar spaces.

This is the only facies that contains CFA (Fig. 5.8). Numerous other phosphatic minerals such as strontian fluorapatite, hydroxylapatite, and apatite contribute to the P2O5 content, which can be as high as 30.5 wt.% (Srivastava and Rocha, 1999). CFA is

Figure 5.7 Facies 5 A-B: Tabular beds of F5, faulted against F3 (faults identified by yellow dotted line) showing sharp contact at base of F5. C: Drill core image of phosphatic, branching columnar stromatolites (S) and intercolumnar dolomite (D), showing branching nature (branching from arrow "a") and rare coalesced branches (arrow "b"); D: Photograph from Galvani mine showing high degree of alteration discussed further in Chapter 7 Paragenesis.

Figure 5.8 Facies 5, photomicrographs of sample FNC05-16.95, PPL and XPL, respectively, of stromatolitic laminae of F5, showing fluorapatite ("P"; pseudoisotropic) and xenotopic dolomite ("D") laminae.

38 present as pristine phosphorite within microbial laminae and strontian fluorapatite occurs as phosphatic coated grains. All other phosphatic phases are secondary and fill intercolumnar pores, fractures, and veins along with calcite and saddle dolomite. Some stromatolite laminae also contain rare reworked fluorapatite intraclasts as well as subangular, silt-sized quartz, muscovite, and orthoclase grains.

Interpretation: Biostomes composing F5 are interpreted to have formed on the lower intertidal flat. Parallel, bifurcate-branching, columnar stromatolites may indicate growth in association with tides and little accommodation for growth (Logan et al., 1964;

Hofmann, 1973). Microbial mat growth may have been prevented by a combination of fluctuating tidal waters, scouring due to storms, tidal water run-off, or desiccation during low tide. Such processes result in bifurcate-branching and anastomosing columnar stromatolites indicating stromatolites that characterize low-energy tidal flat environments

(Logan et al., 1964). The absence of thick storm layers in biostromes suggests that stromatolites thrived in protected bays removed from the influence of storm-generated currents.

5.6 Facies 6: Hummocky cross-stratified grainstone

Facies 6 (F6) consists of hummocky cross-stratified (HCS) grainstone and intercalated thinly bedded lime mudstones (Fig. 5.9A,B). Rare intraclastic lags and stylolites mark the bottom of HCS bedsets. Low angle cross-laminae are formed of fine- grained intraclasts cemented with microcrystalline dolomite. Disseminated pyrite, organic matter, and hematite are common (Fig. 5.9C,D). Microbial layers, 5 to 10 mm thick, are present in more mudstone-rich intervals. 39

Figure 5.9 Gabriel member, facies 6 A: HCS, highlighting the low-angle truncations (yellow dotted line) of the mid-ramp marine deposit; B: Flaser bedding associated with inner ramp microbial-rich lithofacies; C: Sample GAB-31 microcrystalline dolomite with disseminate pyrite; D: Sample GAB-2 recrystallized, microcrystalline dolomite and rare intraclasts associated with organic rich stylolites.

Interpretation: F6 is interpreted to have been deposited in a storm-dominated mid-ramp environment, between fair-weather and storm wave base (Dott and Bourgeois, 1982;

Arnott and Southard, 1990). Preservation of hummocks, the low angle of bedding truncations, repetitive nature of grainstone beds separated by a thin mudstone deposit are all indicative of HCS (Dott and Bourgeois, 1982; Dumas and Arnott, 2006). HCS forms under conditions of storm-generated combined flow, involving both fallout of suspended sediment and wave oscillation (Hamblin et al., 1979; Dott and Bourgeois, 1982; Duke,

1985). Intraclastic lags are interpreted to reflect the transport and deposition of clasts from shallower environments. Muddy microbially laminated intervals probably record the fair-weather deposition of lime mud and bacterial colonization of the seafloor.

CHAPTER 6: PERITIDAL CYCLES AND SEQUENCE STRATIGRAPHY

The vertical and lateral stacking pattern of lithofacies in the Nova America and

Gabriel members of the lower Salitre Formation indicate that peritidal cycles accumulated during an overall marine transgression that was punctuated by higher-order fluctuations in relative sea level (Figs. 6.1, 6.2, 6.3, 6.4). These superimposed fluctuations produced peritidal carbonate cycles (Pratt, 2010) that provide paleoenvironmental context for the accumulation of economic phosphorite. Stacked cycles record constant and then accelerated deepening, suggesting that they define the late phase of the lowstand systems tract (LST) and the early transgressive systems tract

(TST; Figs. 6.4, 6.5; Coe et al. 2003; Eriksson et al. 2005, 2013; Catuneanu and Eriksson

2007; Catuneanu et al. 2011, 2012).

6.1 Peritidal cycles and cycle development

Peritidal cycles are meter-scale aggradational successions that record carbonate deposition in shallow, tide-dominated settings (James, 1984; Goodwin and Anderson,

1985; Grotzinger, 1986a, 1986b; Jimenez de Cisneros and Vera, 1993; Pratt, 2010). In the context of sequence stratigraphy, peritidal cycles are considered parasequences and are the building-blocks of the various systems tracts that define a sea-level cycle

(Catuneanu et al., 2011, 2012). Like all parasequences, peritidal cycles of the lower

Salitre Formation are bounded by minor flooding surfaces reflecting periodic increases in accommodation followed by aggradation and in many cases eventual exposure

(Catuneanu, 2006; Catuneanu et al., 2012). Because drilling did not intersect the base of

and

cores cores - 6 drill res shown in Fig. res co - Correlation of drill cores from the Nova America member. The transgressive surface (TS) The in member. Nova the America identified transgressive from of Correlation drill cores

6.1 Figure features prominent (PS) 4 parasequence the datum correlations.are of facies, used as for remaining Thick intertidal was the between drill Distances identifiedline. blue the TS is a green by a and line (FS) are 5. Flooding surfaces 3.2.

tions. between distance The core correla core -

Regional correlation of the lower Salitre Formation. Paleoshore is to the NE. Regional correlations are based on based are Salitre correlations lower Formation. Paleoshore is toRegional of the the NE. correlation Regional

6. Figure locallydrill observed from specific and the patterns of relationships identification field uncertainty in the Formation upper Formation present contactsSalitre and Bebedouro and lower unobserved of the and locations correlations. regional

Figure 6.3 Legend for stratigraphic sections; FA: facies association; FWB: fair-weather wave base.

Figure 6.4 Relative sea level curve indicating increasing accommodation during the lowstand systems tract (LST), and progressing through the transgressive systems tract (TST), which is followed by decreasing accommodation, not preserved in the Nova America or Gabriel members. The transgressive surface (TS) of the Nova America member indicates an increase in relative sea-level. Progradation of microbialite over deeper marine deposits may indicate deposition that continued into the highstand systems tract (HST). Parasequence development in the Nova America and Gabriel members resulted from minor fluctuations in relative sea level superimposed on a major transgressive event.

Figure 6.5 Composite section of the Nova America and Gabriel members of the Salitre Formation, showing meter-scale, shallowing-upward peritidal cycles.

47 the lower Salitre Formation, only seven complete parasequences were identified (Figs.

6.1, 6.2).

Parasequences of the Nova America member are 3 to 25 m thick and consists of subtidal (F1, F2, and F4) and overlying intertidal (F3 and F5) lithofacies bounded by flooding surfaces (Catuneanu, 2006; Catuneanu et al., 2012). Flooding surfaces are erosional contacts produced as the zone of wave abrasion migrates landward in step with relative sea level rise. Cycle bases are characterized by tidally deposited grainstones (F1,

F2) that locally contain interstratified stromatolitic patch reefs (F4).

The upward-expanding morphology of patch reefs (Fig. 6.6) may be a direct indicator of sea level dynamics during parasequence accumulation (Grotzinger, 1986b;

Walter et al., 1992). The steep lower sides of reefs record aggradation during rapid sea level rise that accompanied the beginning of cycle deposition. As accommodation filled, reefs began to accrete laterally, producing the flared morphology that typifies subtidal reefs in the Salitre Formation. This decrease in accommodation also caused the progradation of evaporitic intertidal mudstones and microbialites over subtidal deposits

(F3, F5). Economic phosphorite is present only in cycle tops where tidal flat sediment is associated with stromatolitic biostromes.

Parasequence development in the overlying Gabriel member records aggradation and progradation of lithofacies in deeper-marine environments than the peritidal sediment of the Nova America member. Intraclastic lags and sharp contacts define the bases of

HCS bedsets, which are 5 to 10 m thick. Interbedded microbial mudstone layers (F6) increase in thickness and abundance gradually over 10 m.

Figure 6.6 Arrecife Ranch subtidal (F4) stromatolitic patch reef. The base maintains a consistent width, growing aggradationally (A) for approximately 0.5 m before widening at the top, growing out and displaying a progradational nature (P). F2 fills in accommodation between patch reefs, F4.

6.2 Sequence Stratigraphy

Stratal stacking patterns of the peritidal parasequences of the Nova America and

Gabriel members indicate that deposition occurred during the late LST into the TST.

LST conditions are observed in parasequences PS1, PS2, PS3, PS4, and PS5, which are characterized by aggradational packages of peritidal sediment with thick intertidal packages (F3 and F5). A significant increase in accommodation is present at the base of

PS6, identified as the transgressive surface (TS), which represents rapid flooding and initiation of TST deposition. Subsequent parasequences (PS6 and above; Fig. 6.2) record increasing depth, displaying a retrogradational stratal stacking pattern indicative of transgressive conditions.

6.2.1 LST

The basal contact of the Salitre Formation rests unconformably on either glacial diamictite of the Bebedouro Formation or sandstone of Mesoproterozoic Chapada

Diamantina Formation. Thin subtidal deposits (F1) are <1 m thick, and rest directly on the unconformity. A thicker package of intertidal sediments (F3) conformably overlies the subtidal layer. Intertidal strata are 3 to 4 m thick and contain evaporite pseudomorphs, common tepee structures, and mud cracks (Fig. 6.2).

PS2 and PS3 are only documented in the longest drill core, FNC05 (Fig. 6.7).

PS2 is 5 m thick and consists of subtidal lithofacies that grade upwards into intertidal deposits that are 2 m in thickness. PS3 consists of 15 m of subtidal sediment (F1, F2, and

F4) that rest sharply on intertidal sediment (F3) of PS2. As in PS2, PS3 changes upwards

50 from subtidal facies (F2 and F4) to a thin package of intertidal sediment (F3 and F5) that is 1 to 2 m thick.

PS4 and PS5 contain thicker intertidal deposits and thus reflect more aggradational accumulation than underlying parasequences Basal subtidal deposits (F1,

F2, and F4) are 10 to 15 m thick overlain by 5 to 10 m of intertidal sediment (F3 and F5).

Intertidal facies in these parasequences also contain the thickest and best-developed phosphatic stromatolites in the Nova America member (PS5; Fig. 6.1).

6.2.2 TST

The basal contact of PS6 is a well-developed scoured and winnowed surface that is intersected in all but two drill cores. Basal subtidal deposits (F1 and F2) are 20 to 30 m thick and interpreted to represent the onset of purely aggradational accumulation. The sharp lower contact and the aggradational nature of PS6 suggests that this contact is the transgressive surface of erosion (Figs. 6.1, 6.2). Intertidal deposits (F3) that overlie this thick subtidal unit are 5 m thick.

Field relationships indicate that the Gabriel member is above PS6 and other eroded parasequences. The Gabriel member also consists of parasequences, but reflects deposition in a deeper-water setting than the Nova America member. Cycles are ~ 10 m thick and composed of basal lime mudstones overlain by HCS grainstone (F6).

Parasequences shallow slightly up-section and contain progressively more microbialite, which is interpreted to reflect the development of prolific benthic, microbial mats. This shallowing may reflect more pronounced progradation associated with onset of highstand conditions (Fig. 6.8).

Figure 6.7 Logged section of the longest drill-core FNC05, with depositional setting interpretation and parasequence development indicated on the left. See Fig. 6.3 for legend.

Figure 6.8 Idealized stratal stacking patterns of peritidal carbonates in response to relative sea level (RSL) change. Peritidal environment consists of subtidal, below low water mean (LWM), intertidal, between high water mean (HWM) and LWM, and supratidal, above HWM; A: Aggradational stacking packages result from equal carbonate production rates and increasing accommodation producing consistent stacking of parasequences. Aggradational sequences typically occur at inflection points on a RSL curve, the TS and the MFS; B: Retrogradational stacking patterns or back-stepping of depositional systems occur when increasing accommodation is greater than rate of carbonate sedimentation, resulting in incremental deepening of parasequences, and is typical during the TST; C: Progradational packages result from high sedimentation rates relative to accommodation. Progradation occurs when accommodation begins to decrease in the HST and can occur through the FSST and LST. See Fig. 6.3 for legend.

CHAPTER 7: PARAGENESIS OF THE NOVA AMERICA MEMBER

Peritidal carbonates of the Nova America member have experienced a complex paragenesis, including multiple dolomitization and phosphogenic events. Stages include

1) carbonate deposition; 2) seafloor diagenesis and CFA precipitation; 3) meteoric and shallow-burial diagenesis; 4) hydrothermal alteration and secondary phosphate enrichment; and, 5) deep-burial diagenesis (Fig. 7.1, 7.2).

7.1 Stage I: Carbonate deposition

Stage I encompasses the accumulation of all carbonate facies in peritidal environments (Fig. 7.3). Ooids in energetic grainy facies (F2, F3, F4) are identified by micrite envelopes despite complete recrystallization. Fine-grained carbonate sediments accumulated along less energetic segments of the paleocoast (F3). Stromatolites formed phosphatic biostromes on intertidal flats (F5) as well as reefs in tide-swept subtidal environments (F4).

Interpretation: Ooids precipitate in current-swept tropical environments supersaturated with respect to calcium carbonate (Bathurst, 1975; Tucker and Wright, 1990; Hardie,

2003; Duguid et al., 2010). The complete recrystallization and/or dissolution followed by cement precipitation of cortical layers suggests the original mineralogy of ooids was aragonite (Stages II and III).

Figure 7.1 Paragenetic sequence for the Nova America unit. Microcrystalline dolomite (D1), recrystallized granular, planar-s dolomite (D2), coarse crystalline, euhedral zoned dolomite lining vugs and rimming grains (D3), and saddle/baroque dolomite (D4).

mitization, hydrothermal -

dded packstone) as an example. Stage I and II an and dded represent I example. as packstone) Stage be - filled fractures and stylolites. and filled fractures - ugs. Stage IV, hydrothermal alteration, is best represented by the alteration, by the IV, represented is best hydrothermal ugs. Stage to lenticular

calcite - cutting naturecutting of the - Simplified paragenesis using Facies 3 using Facies (flaser Simplified paragenesis

7.2 Figure mats, diagenesis, includes dolo of microbial and ooids,input. III, meteoric extensive intraclasts, Stage formation detrital porosity, dolomite of and of zoned creation liningformation v diagenesis, burial post most dolomite, occurred of saddle V, deep Stage occludes of the porosity. formation which by indicated the cross as alteration, 56

Figure 7.3 Stage I, carbonate deposition, A: FNC02-38.25, XPL polished thin section photomicrograph, contains detrital grains, quartz and orthoclase, as well as authigenic and diagenetic minerals associated with microbial laminae; B: FNC05-39.30 (PPL polished thin section photomicrograph) contains ooids with extensive micritization of cortical layers, indicated with a white arrow.

Carbonate precipitation in inter- and supratidal settings was probably enhanced by evaporative concentration (Grotzinger and Kasting, 1993; Schröder et al., 2005). In these calm settings newly precipitated calcium carbonate, aragonite and high-Mg calcite, accumulated via suspension rain and later became indurated on the seafloor through subsequent seafloor diagenesis (Stage II). Cyanobacterial mats and stromatolites are interpreted to have aided this processes by binding sediment (Reid et al., 2000).

7.2 Stage II: Seafloor diagenesis and CFA precipitation

Stage II represents all seafloor diagenesis and authigenic mineral precipitation, including gypsum, anhydrite, and pristine phosphorite. Tepee structures (cm-scale) and replaced evaporite nodules (mm- to cm-scale) are present in intertidal facies (F3, Fig.

7.4), and rarely in subtidal deposits (F4). Displacive growth of tepees caused buckling of the sediment layers as evaporite growth pushed layers aside (Fig. 7.4A). Low-Mg calcite and quartz gypsum, and/or anhydrite, pseudomorphs are identified in thin section as acicular crystals that are 100 to 400-μm in length (Fig. 7.4B; Scholle and Ulmer-Scholle,

2003). Gypsum crystals grew displacively, pushing aside lime mud that is now dolomicrite.

Seafloor cementation produced hardgrounds and firmgrounds that were later reworked into subrounded to rounded intraclasts (F1, F2, F3). Internally, many intraclasts have grumuleuse-like textures, suggesting that they were derived from microbial mats.

Synsedimentary cementation also created grapestones (F3) of cemented ooids and intraclasts (Fig. 5.5).

Figure 7.4 Stage II seafloor diagenesis and authigenic precipiation, A: Outcrop exposure showing a tepee structure with evaporite precipitation (cream to tan coloured fill) from the Achado section; B: thin section photomicrograph of sample FNC01-7.30, polished and stained, contains gypsum and possible other evaporite minerals (yellow arrow), which have been replaced by chalcedony, in micritic dolomite matrix, with calcite veins (stained red by Alizarin red-S).

Pristine phosphorite is associated only with columnar stromatolitic biostromes in intertidal environments (F5; Fig. 7.5). Fluorapatite (Ca5.061(P2.87O11.46)F0.89) laminae and in situ peloids were originally organic-rich microbial layers in stromatolites. Phosphatic and microbial laminae are cemented by subhedral dolomite and dolomicrite, respectively.

Intercolumnar zones are filled by fine-grained carbonate sediment that is now microcrystalline dolomite (Stage II and III). Rare coated phosphate grains are now composed of strontian fluorapatite (Ca9.37Sr0.63(PO4)6F2; Fig. 7.6). Reworked phosphate grains are also present in some laminae as thin lags enveloped by drusy quartz (Figs. 7.7,

7.8, 7.9). Pyrite framboids and muscovite are also present (Fig. 7.9).

Interpretation: Diagenesis affects carbonate sediment as soon as it forms (James and

Choquette, 1990b). Grapestones and intraclasts (F3) are interpreted to represent tide and storm reworking of hardgrouds and firmgrounds (Bathurst, 1975, 1987; James and

Choquette, 1990b; Kazmierczak et al., 2004). Such early seafloor lithification was probably facilitated by microbial processes, as indicated by the presence of microbial mats prominent throughout 80% of the peritidal facies and the grumuleuse texture of intraclasts. Microbial processes aid seafloor lithification by trapping and binding sediment as well as promoting carbonate precipitation (Tucker and Wright, 1990; Pierson et al., 1992; Reid et al., 2003; Kazmierczak et al., 2004). Photosynthetic cyanobacterial communities, common in the Nova America member, may have drew down CO2, increased carbonate saturation and preconditioned the water for carbonate precipitation.

Aragonite, and to a lesser extent high-Mg calcite, are interpreted to have been the dominant carbonate mineralogies of the marine precipitates and cements because of the

Figure 7.5 Authigenic precipitation of francolite, pristine phosphorite facies, A, B: photomicrograph of polished and stained thin section FNC02-36.25 PPL and CL, respectively, cross-section through a phosphatic stromatolite. Francolite (honey-brown under PPL and purple under CL) laminae alternate with subhedral dolomite laminae (orange and red zoned crystals under CL) within small columnar stromatolites of F5. Pristine phosphorite precipitated as in-situ peloids and laminae and have been altered during Stage IV hydrothermal alteration resulting in fluorapatite mineralogy.

Figure 7.6 Coated phosphate grain, an authigenic precipitate A, B: photomicrograph of polished and stained thin section FNC01-28.10 under PPL and XPL, respectively. Each cortical layer is a continuous phosphatic (pseudoisotropic) coating, with a winkled texture.

Figure 7.7 Hydraulically reworked phosphorite deposit, A, B: Polished thin section photomicrograph, FNC02-25.35 under PPL and XPL, respectively showing rounded, reworked phosphatic peloids (honey-brown in PPL and pseudoisotropic under XPL) with rounded dolomite intraclasts (microcrystalline dolomite clasts). Hydrothermal alteration during Stage IV changed mineralogy from CFA to fluorapatite.

Figure 7.8 A: Photomicrograph of sample FNC02-25.35 under CL; fluorapatite displays purple luminescence whereas dolomite has a orange-yellow zoned luminescence. B: FNC10-21.60 photomicrograph under cross-polarized light showing needles of quartz cement and clay growing along fractures of reworked phosphatic laminae.

Figure 7.9 BSE image of a reworked phosphatic deposit, with EDS analysis with clays and pyrite occurring within fractures (sample FNC10-21.60).

65 lack of preserved internal fabric (Bathurst, 1975; James and Choquette, 1990a; Tucker and Wright, 1990).

Carbonate fluorapatite (CFA) is interpreted to form in situ in stromatolitic

3- biostromes when pore fluid becomes supersaturated with respect to PO4 (Glenn et al.,

1994; Föllmi, 1996; Pufahl, 2010). The intimate association of CFA with stromatolites indicates that microbial processes were probably critical for phosphogenesis.

Stromatolites are composed of photosynthetic cyanobacterial communities, which trap and bind sediment as well as promote the precipitation of carbonate minerals (Pierson et al., 1992). Microbial degradation of cyanobacterial mats releases phosphate, which concentrates in the inter-carbonate laminae of columnar stromatolites to facilitate phosphogenesis (Banerjee, 1971; Banerjee et al., 1980; Föllmi, 1996; Crosby and Bailey,

2012). Microbial degradation follows a series of redox-controlled processes, and although all release phosphate, bacterial sulfate reduction (BSR) and methanogenesis are thought to be the most efficient at increasing concentrations of phosphate in pore fluid

(Glenn et al., 1994; Pufahl, 2010).

Redox-controlled microbial processes are also interpreted to be responsible for the rare coated phosphate grains present. Unlike ooids, in which cortical layers are precipitated within the water column, the coated phosphatic grains precipitate authigenically beneath the sediment-water interface (Fig. 7.6; Garrison and Kastner,

1990; Pufahl and Grimm, 2003). Changes in the redox potential of pore fluid within the zone of phosphogenesis regulate the precipitation of francolite-rich (strontian fluorapatite) cortical layers (Pufahl and Grimm, 2003).

Lime mud was dolomitized to form dolomicrite. Early authigenic dolomitization probably accompanied evaporite precipitation and microbial degradation. Gypsum

2- precipitation increases the Mg/Ca ratio, increases pH, and removes sulfate (SO4 ), a kinetic inhibitor of dolomitization (Warren, 2000; Zentmyer et al., 2011). The presence of evaporite pseudomorphs, originally gypsum and anhydrite, are indicators that evaporation influenced early dolomitization, especially in intertidal and supratidal settings. Microbial processes are associated with phosphogenesis, and BSR and methanogenesis are also two mechanisms that facilitate dolomite formation (Wright,

1997; Mazzullo, 2000).

7.3 Stage III: Meteoric and shallow-burial diagenesis

Stage III represents carbonate dissolution, dolomitization, and silicification that occurred during meteoric diagenesis. Evaporite pseudomorphs were replaced by quartz or low-Mg calcite. Tepee structures filled with acicular crystals exhibit evidence of evaporite mineral formation (Fig. 7.4).

Non-fabric-selective dissolution produced of vugs in intertidal facies (F3). Vugs and grain boundaries are coated with isopachous euhedral to subhedral dolomite crystals

(Fig. 7.10, 7.11A). Isopachous dolomite is characterized by zoned dolomite crystals, visible under transmitted light as cloudy and limpid bands, μm-scale layers. Under CL, rimming dolomite consists of a minimum of 3 zones of dull orange, dull red, and bright orange luminescent bands (Fig. 7.10B). Sucrosic dolomite is also rarely present (Fig.

7.11B). Pore spaces are most commonly occluded with tightly packed planar-S to anhedral, xenotopic dolomite (Sibley and Gregg, 1987).

Figure 7.10 Polished and stained thin section photomicrographs of sample FNC02-48.88 under PPL and CL, respectively, displaying meteoric diagenesis; A: isopachous grain- rimming euhedral dolomite with cloudy cores ("C") and limpid rims ("L"), separated by a yellow-dashed line; B: pore-lining zoned dolomite cement, displaying alternating bands of high (yellow) and low (dark, pink-purple) luminescence. Pore is occluded by low- luminescent saddle dolomite (pink).

Figure 7.11 Evidence of shallow burial, phreatic dolomitization, polished and stained thin sections photomicrographs; A: Recrystallized ooid with a micritic envelope, planar-s dolomite infilling the completely dissolved ooid as well as a rimming, isopachous dolomite surrounding the grain, FNC02-48.88; B: Sucrosic dolomite replacing calcite matrix, fabric-destructive dolomite rhombs showing some evidence of cloudy centers and clear rims, FNC05-53.45.

In facies F3, ooids and intraclasts have undergone fabric-selective dissolution.

Grain boundaries are defined by an isopachous coating of finely crystalline dolomite

(Fig. 7.10). These dolomicrite walls are 5 to 20 µm thick and appear dark in thin section and bright yellow under CL. Ooid interiors consists of subhedral to anhedral dolomite cement (Figs. 7.10, 7.11B; Gregg and Sibley, 1984; Sibley and Gregg, 1987; Gaswirth,

2004).

Interpretation: Sucrosic and isopachous dolomite suggest pervasive phreatic dolomitization during meteoric diagenesis. During meteoric diagenesis, slight dilution of marine fluid with meteoric water promotes dissolution of metastable mineralogies

(aragonite and high-Mg calcite) of sediment and grain interiors (James and Choquette,

1990a). The sucrosic, zoned dolomite with euhedral to subhedral crystals are characteristic of dolomitization in the “mixing-zone” or the Dorag dolomitization model

(Figs. 7.11, 7.12; Tucker and Wright, 1990; Warren, 2000; Choquette and Hiatt, 2008).

Vugs and pores, which display non-fabric-selective dissolution, are interpreted to have been generated at the same time as planar-S dolomite under meteoric diagenetic conditions in the shallow burial environment. The isopachous nature of the dolomite cement lining vugs supports this shallow burial, phreatic diagenetic environment, as dolomite displays zoned luminescence (Choquette and James, 1990; James and

Choquette, 1990a; Tucker and Wright, 1990; Gaswirth et al., 2007).

The alternation of bright and dull zoned luminescence, displayed under CL (Fig.

7.10B, 7.13B), is attributed to varying trace elements, especially Mn2+ and Fe2+ ratios, which are redox-sensitive and fluctuate in the shallow burial realm (James and

Choquette, 1990a; Choquette and Hiatt, 2008). Mn2+ is known as an activator and results in bright luminescence, whereas Fe2+ is a quencher and produces a duller luminescence.

The ratio of Mn2+ and Fe2+ controls the luminescence rather than the absolute concentrations (Choquette and James, 1990).

Dolomitization of micrite envelopes is interpreted to have occurred contemporaneously with dissolution of ooid and intraclast interiors. The presence of collapsed nuclei of ooids (Fig. 7.10B, 7.11B) helps identify ooids and may indicate that some nuclei were composed of high-Mg calcite which underwent preferential dolomitization (Scholle and Ulmer-Scholle, 2003). It is possible ooids initially recrystallized to low-Mg calcite prior to undergoing dolomitization, a common process that occurs in the meteoric diagenetic environment (James and Choquette, 1990a).

Along with extensive dolomitization, silicification and calicitization of evaporite minerals also occurred during Stage III. Evaporite pseudomorphs have been replaced by quartz and calcite. It is interpreted that under the influence of meteoric fluid the evaporite minerals dissolved and fluid enriched in silica from weathering of silica rich rocks and percolation into pore water lead to the precipitation of quartz (Folk and

Pittman, 1971; Ulmer-Scholle and Scholle, 1994).

7.4 Stage IV: Hydrothermal alteration and secondary enrichment of phosphate

Stage IV represents hydrothermal alteration and secondary enrichment of phosphate, producing economic phosphorite. Planar-S dolomite recrystallized to tightly packed, xenotopic dolomite, particularly in subtidal facies (F1). Vugs produced during meteoric dolomitization were further occluded by baroque (saddle) dolomite (Fig. 7.12),

71 identified by curved faces and undulose extinction in thin section. Saddle dolomite crystals are 200 μm to >1 mm and commonly contain opaque inclusions <10μm (Figs.

7.12, 7.13).

Under CL, saddle dolomite crystals are dull red (Fig. 7.10). While saddle dolomite is present in most facies, vugs and fractures associated with F3 are commonly filled by fluorite (Fig. 7.14). Together with saddle dolomite and fluorite, other minerals in vugs and veins include pyrite (<10 to 25 μm), molybdenite (~100 μm), barite (~50

μm), quartz, muscovite 1M and 2M1, and bitumen (Fig. 7.13).

Hydrothermally produced veins are occluded by saddle dolomite (Figs. 7.15,

7.16). Secondary phosphate is present in fractures and veins as fluorapatite and hydroxylapatite (Ca9.42Sr0.18(PO4)6(OH)1.2(H2O)0.4; Fig. 7.15). Such secondary phosphate enrichment is closely associated with ferroan saddle dolomite, euhedral and framboidal pyrite, quartz, muscovite, and orthoclase, which precipitated penecontemporaneously in hydrothermal veins.

Interpretation: Hydrothermal alteration occurs when fluid temperatures are greater than the surrounding country rock, commonly under shallow burial conditions when temperatures exceed ~80°C (Choquette and James, 1990; Lapponi et al., 2014). In the

Nova America member, compressional tectonics associated with the Brasiliano orogeny are interpreted to have created ideal conditions for the generation of hydrothermal fluids and their migration (Misi et al., 2005). Brecciated dolomitic intraclasts commonly observed in calcite veins (Stage V) indicate hydraulic fracturing of the rock (Bathurst,

1975; Choquette and James, 1990; Lonnee and Machel, 2006; Smith, 2006).

Figure 7.12 Polished and stained thin section photomicrograph of sample FNC05-30.20 A, B: PPL and XPL, respectively, showing hydrothermal saddle dolomite occluding a fracture. Note the curved crystal structure (yellow arrow) and sweeping extinction (white arrow).

Figure 7.13 Polished and stained thin section photomicrographs; A, B: FNC02-48.88 provides evidence of hydrothermal indicators including saddle dolomite ("SD"), pyrite and molybdenite (white arrow).

Figure 7.14 Polished and stained thin section of sample FNC01-28.10; A: Plane polarized photomicrograph displaying non-fabric-selective dissolution. Pore-lining zoned, euhedral dolomite is present and porosity is occluded by fluorite (purple); B: CL photomicrograph highlighting zoned luminescence of dolomite crystals (orange and yellow rhombs), fluorite (blue), and strontian fluorapatite (purple).

Figure 7.15 A, B: Photomicrographs of sample FNC02-31.20 under PPL and CL, respectively, with phosphate minerals precipitated along fractures (purple-blue luminescence) associated with hydrothermal alteration. Vugs and fractures are occluded by saddle dolomite ("SD"), further evidence of hydrothermal origin (Lonnee and Machel, 2006).

Extensive dolomitization of the Nova America member occurred under hydrothermal conditions, as indicated by the presence of xenotopic and saddle dolomite

(Choquette and James, 1990). Kinetically, dolomitization is favored by warm fluids above 60C (Gregg and Sibley, 1984; Warren, 2000; Luczaj, 2006). Planar dolomite develops when temperatures are below 50 – 60C, known as the critical roughening temperature (CRT), above which non-planar (xenotopic and saddle) dolomite forms

(Gregg and Sibley, 1984; Sibley and Gregg, 1987; Lonnee and Machel, 2006).

Homogenization temperatures in the Irecê Basin have been documented between 150 and

200°C (Misi et al., 2005). The dull luminescence of saddle dolomite indicates a low

Mn/Fe ratio and is typical of burial diagenetic cements (Choquette and James, 1990).

Saddle dolomite is also commonly associated with hydrocarbons, indicating that its formation occurs within the oil window, which typically is marked by temperatures between 60 – 150C (Tucker and Wright, 1990). Bitumen present along fractures in intertidal facies (F3) was mobilized during hydrothermal alteration as the organic-matter rich intertidal flats were heated above 60°C, where it underwent subsequent biodegradation (Smith, 2006; Smith and Davies, 2006). The hematite found with the bitumen probably resulted from oxidation of pyrite that formed with the petroleum- bearing hydrothermal waters (Challis, 1975). This association with hydrocarbons further supports the hydrothermal interpretation and is consistent with the presence of bitumen.

Hydrothermal carbonate alteration is also associated with Mississippi Valley-type

(MVT) Pb-Zn deposits (Gregg and Shelton, 1989; Lonnee and Machel, 2006; Luczaj,

2006), which are present throughout the Irecê Basin (Kyle and Misi, 1997; Misi et al.,

2005). MVT deposits are characterized by the presence of metal sulfides such as pyrite,

77 galena, and sphalerite, as well as the non-metallic minerals fluorite, dolomite, and calcite

(Gregg and Shelton, 1989). Pyrite and fluorite are the most common of these minerals in the study area. The source of F for fluorite precipitation was probably the dissolution of primary fluorapatite (McArthur, 1980; Spirakis and Heyl, 1988; McArthur and Herczeg,

1990; Zeeh, 1995).

7.5 Stage V: Deep-burial diagenesis

Stage V encompasses burial diagenesis that occurred after hydrothermal alteration. The hydrothermal minerals are cross-cut by veins (Figs. 7.17, 7.18, 7.19) and stylolites (Fig. 7.20). Veins are ~50 μm to >1 cm wide, composed of poikilotopic calcite and rare quartz, and cross-cut all previous fabrics including saddle dolomite. Planar-S dolomite is also observed overgrowing calcite veins as euhedral to subhedral rhombs <10

µm. Vertical and horizontal fractures average 100 μm thick. Under CL, calcite and dolomite occluding these fractures generally display a dull red luminescence. Hematite is concentrated along stylolites (Fig. 7.19).

Interpretation: The cross-cutting nature of veins and stylolites indicates further compaction after all major dolomitization events and hydrothermal mineralization

(Bathurst, 1975; Choquette and James, 1990). Late-stage poikilotopic calcite probably precipitated from Ca2+ that was derived from the dedolomitization of saddle dolomite

(Merino and Canals, 2011). Hydraulic fracturing of the rock during hydrothermal alteration is occluded by this late-stage calcite precipitate.

Figure 7.16 A, B: Core photograph and thin section photomicrograph, respectively, of vein FNC18-30.20, a pink-cream-colored vein cross-cut stromatolites. The vein is composed of saddle dolomite.

Fig. 7.17 A, B: FNC02-49.7 drill core image and thin section photomicrograph, respectively, of dominantly low-Mg calcite and rare dolomite.

Figure 7.18 A, B: FNC01-7.30 low-Mg calcite vein in drill-core and photomicrograph, respectively, showing brecciated F2. Thin section has been stained with Alizarin red-S to highlight the calcite (red) and dolomite (not stained) components, with minor quartz inclusions.

Figure 7.19 A, B: FNC18-14.57 PPL and XPL photomicrograph, respectively, showing the cross-cutting nature of late-stage burial fractures (yellow arrow), which horizontally cross-cut hydrothermal saddle dolomite ("SD") and Stage III & IV mineral phases.

Figure 7.20 A: FNC02-25.35 photomicrograph demonstrates deep burial diagenesis displaying stylolitic compaction with hematite concentrated along stylolites. Stylolites cross-cut all pre-existing minerals and grains, such as the ooid highlighted by a yellow arrow; B: Sample FNC05-48.88 also provides evidence of deep burial diagenesis post- hydrothermal alteration, because the stylolites cross-cut saddle dolomite ("SD").

7.6 Economic phosphorite

Peritidal carbonates of the Nova America unit exhibit a complex diagenetic history that resulted in both primary and secondary precipitation of phosphatic minerals.

Primary authigenic francolite formed on intertidal flats in association with stromatolitic biostromes. Later hydrothermal processes led to the dissolution of francolite and the remobilization of P, F, and Ca to produce fluorapatite and hydroxylapatite (McArthur,

1980; Yoshimura et al., 2004). This secondary precipitation of phosphatic minerals enriched intertidal deposits in the Nova America member to produce an economic phosphorite.

CHAPTER 8: STABLE ISOTOPE GEOCHEMISTRY

Seventy-one samples were analyzed for their stable C and O isotopic composition

(Figs. 8.1, 8.2; Table AIII.1). Analysis of the various lithofacies, veins, and calcretes provide the foundation to interpret these data in depositional, diagenetic, and hydrothermal context.

8.1 Results

Seven calcrete samples, composed of friable, white to pale yellow chalky, low- magnesium calcite were analyzed. Values for the calcrete samples plot between -9.2‰

13 18 and -8.0‰ δ CVPDB and between -3.3‰ and -1.3‰ δ OVPDB. Average values are -8.7‰ and -1.8‰ respectively, with the exception of one outlier. Values for sample CAL G2-18

13 18 are +0.6‰ δ CVPDB and -3.7‰ δ OVPDB.

Six vein samples analyzed are composed of low-Mg calcite with inclusions of

18 saddle dolomite. All of the veins display similar low δ OVPDB values, ranging between -

18 13 4.7‰ and -3.0‰ δ OVPDB, with an average of -4.2‰. δ CVPDB values possess a greater

13 degree of variation, plotting between -7.0‰ and +5.6‰ δ CVPDB, averaging -1.75‰.

Sample V18-30.20 is composed almost exclusively of saddle dolomite and has the

13 highest δ CVPDB value of +5.6‰.

Thirteen F1 samples, five F2 samples, twenty-eight F3 samples, twelve F4 samples, and twelve F5 samples were analyzed (Fig. 8.1). All samples have negative

δ18O values, ranging from -10.2‰ to -0.5‰, with an average of -3.9‰. Values for δ13C have a much larger range from -9.2‰ to +10.0‰, with an average of +2.8‰. Each facies sample set contains at least one outlier, and when these samples are excluded, the

Figure 8.1 δ18O and δ13C results from samples of the Nova America member plotted by facies.

Figure 8.2 δ18O and δ13C results from samples of the Nova America member plotted by facies, including hydrothermal veins and surficial samples, showing both δ18O values measured and values corrected to equilibrate calcite to dolomite. All samples are dominated by low-Mg calcite except V18-30.20 (highlighted by arrow), which is dominantly saddle dolomite, and therefore carries a hydrothermal signature. All calcite 18 18 veins plot between -8.0 and -6.0‰ for δ Omeasured and -5.0 to -3.0‰ for δ Ocorrected (Land, 1980), with a wide range in δ13C values.

18 13 values range between -4.0‰ and -1.75‰ δ OVPDB and +2.6‰ and +9.0‰ δ CVPDB (Fig.

8.2) with averages of -3.2‰ and +5.5‰, respectively.

Interpretation: The paragenesis of the Nova America member is complex and records multiple dolomitization, recrystallization and phosphogenic events. Stable isotopic data support this interpretation and highlight the importance of hydrothermal fluids in significantly altering the isotopic composition of original carbonate minerals and authigenic francolite (Land, 1980; Jarvis et al., 1994; Warren, 2000). The pervasive and widespread dolomitization that characterizes the Nova America member is common in hydrologically open systems, which can deliver Mg and allow dolomitizing fluid to thoroughly flush limestones (Warren, 2000; Machel and Lonnee, 2002).

Although the precise isotopic composition of Neoproterozoic seawater is not well- constrained, data from the calcretes and hydrothermal veins provide possible end members to properly assess such dolomitization and related alteration as documented petrographically. Estimates for Neoproterozoic seawater have broad ranges from ca. -

13 18 12‰ to +10‰ δ CVPDB and -18‰ and -1‰ δ OVPDB at ca. 600 Ma (Shields and Veizer,

2002), which probably reflects complex paragenetic histories involving diagenesis and possibly metamorphism.

Calcretes provide meteoric signatures resulting from Tertiary weathering of rocks in the Irecê Basin (cf. James and Choquette, 1990a). The low δ13C values are characteristic of meteoric environments, because microbial processes increase 12C in soil

13 gas and pore water. The values for the outlier, CAL G2-18, +0.6‰ δ CVPDB and -3.7‰

18 13 δ OVPDB, have a higher δ C value than expected for meteoric signatures (Fig. 8.2). CAL

G2-18 is well indurated and the coloring is consistent with hydrothermal alteration seen in phosphatic samples from the Galvani Mine site. Therefore, the isotopic signature of this hardpan is interpreted as the consequence of hydrothermal alteration, rather than meteoric diagenesis.

The low δ18O values of hydrothermal veins is interpreted to record fractionation associated with high precipitation temperatures >80°C (O’Neil and Epstein, 1966;

Choquette and James, 1990; Hoefs, 1997; Warren, 2000; Loyd and Corsetti, 2010). The high δ13C values probably reflect precipitation from hot pore waters where fermentation of organic-rich lithofacies produced methane, which can yield water values of up to

+15‰ (Warren, 2000). Fermentation or methanogenesis involves methanogenic bacteria stripping oxygen from organic matter and producing methane (CH4). CH4 is enriched in

12 13 C, whereas C dominates in resultant CO2 in pore water (Moore et al., 2004). This

2- process is linked to increased pH and alkalinity, by concentrating CO3 and decreasing the Mg and Ca hydration, creating conditions conducive for the precipitation of saddle dolomite (Mazzullo, 2000).

Similar low δ18O and high δ13C values from dolomites comprising the Nova

America member suggest that the late-stage hydrothermal fluids from which saddle dolomite precipitated reset the isotopic composition of earlier carbonate phases. The only hydrothermal vein, V18-30.20, carries this hydrothermal signature (Fig. 8.2). Nova

America dolomite values are comparable to Paleozoic and Precambrian burial and saddle dolomite values (Fig. 8.3; Warren, 2000). Such hydrothermal alteration coincides with the oil and gas window (Choquette and James, 1990) and thus is also interpreted to have produced the bitumen present in organic-rich lithofacies.

Figure 8.3 Stable C and O isotopic compositions from the Nova America member and from previous studies on modern and ancient dolomites. Modified from Warren (2000). This comparison indicates that dolomites of the Nova America have similar isotopic values as other ancient burial and hydrothermal dolomites.

The majority of outliers from all sample sets are from exposed surface localities with intense modern weathering. Outliers have low δ13C values and the majority have higher δ18O than burial dolomites. The low δ13C values are similar to those from calcretes and indicate the importance of Tertiary weathering on the isotopic signal of some samples.

Contrary to published chemostratigraphic syntheses (Misi and Veizer, 1998;

Azmy et al., 2006; Misi et al., 2007, 2011, 2014, n.d.; Caxito et al., 2012) these data indicate that limestones in the Irecê Basin should not be used for stratigraphic correlation.

Any primary signal has apparently been completely obliterated during burial diagenesis and subsequent hydrothermal alteration.

CHAPTER 9: DISCUSSION

The Nova America phosphorites accumulated during the onset of Earth's second phosphogenic episode, spanning the late Neoproterozoic into the early Phanerozoic.

These phosphorites record an important transition in the chemical state of Earth's oceans and atmosphere leading to the development of true phosphorite giants. The onset of this phosphogenic episode coincides with the NOE and the evolution of complex organisms

(Och and Shields-Zhou, 2012; Pufahl and Hiatt, 2012). Understanding phosphogenesis in the Nova America member provides new information on the nature of the Neoproterozoic

P-cycle and implications for the NOE.

9.1 Depositional model and phosphogenesis

Paleogeographic reconstructions of the São Francisco Craton (SFC) suggest that the Salitre Formation accumulated on a passive margin at low latitude (~30°S; Misi et al.,

2005; Teixeira et al., 2010; Li et al., 2013). Deposition occurred during the late stages of the break-up of supercontinent Rodinia and post-Marinoan glaciation, ca. 635 Ma and

600 Ma (Li et al., 2013). Post-depositional deformation of the Salitre Formation took place during the late Neoproterozoic Brasiliano orogeny, which is associated with the amalgamation of western Gondwana.

Accumulation of peritidal limestones of the Nova America member occurred on a late Cryogenian epeiric ramp (Fig. 9.1). The lack of a prominent shelf-slope break on seismic profiles through the southern SFC support this interpretation (Martins-Neto,

2009). Shallow, sunlit waters and tidal flats were conducive for carbonate accumulation and microbial colonization, however, only intertidal stromatolitic biostromes created the

92 necessary conditions for phosphogenesis. Intertidal deposits (F3, F5) record periods of subaerial exposure, evaporite precipitation, and the deposition of intermittent, thin storm deposits. Subtidal deposition was dominated by grainstones (F2) and stromatolitic reefs

(F4) that were shaped by strong tidal currents. The contrasting depositional conditions between intertidal and subtidal environments are interpreted to have played a critical role in the accumulation of phosphorite in the Nova America member.

9.1.1 Stromatolites, hydrothermal alteration, and economic phosphorite

Stromatolites form through microbial biomineralization processes and contain a complex consortium of photoautotrophs and heterotrophs (Reid et al., 2000; Baumgartner et al., 2006; Spadafora et al., 2010). Their metabolic processes are important in preconditioning pore water for phosphogenesis. Cyanobacteria colonizing the exposed layer of stromatolites (Walter et al., 1992; Grotzinger and Knoll, 1999; Aloisi, 2008;

Spadafora et al., 2010) have the ability to store excess inorganic phosphorus, more than required for their metabolic requirements, which can lead to accelerated accumulation of phosphate (Banerjee et al., 1980; Diaz et al., 2008). Precambrian stromatolites could have further promoted phosphogenesis through the production of oxygen, which pushed redox-sensitive phosphogenic processes into the sediment to concentrate P (Brasier and

Callow, 2007; Nelson et al., 2010; Pufahl and Hiatt, 2012).

FWB) marks FWB) marks section represents the represents section - weather base base ( weather - 2, and F4 with energetic tidal tidal F4 with energetic 2, and ramp consists of the Gabriel member, F6. High water mean mean member, ramp Gabriel consistsF6. of the High water - ramp environmental settings. Geometry in the cross in the environmental settings. Geometry ramp - Depositional model for the lower Salitre Formation. The inner consists Salitre inner of intertidal Nova Formation. The deposits the lower ramp of the for model Depositional

America member, F3 (green) and F5 (yellow and F1, member, stromatolites)F subtidal(tan), F3 (green) deposits and America Mid marks). (arrows) producing (ripple tidal ripples currents Fair high the marks, low tide (LWM) mark and respectively. tide low water mean and (HWM) themid ramp and the inner between the transition member. America unconformity and of the Nova relationship basal of parasequence average of the an architecture Figure 9.1 Figure

Although photoautotrophs dominate the active growth surface of a stromatolite, associated heterotrophic bacteria dominate internal biomineralization and degradation of the stromatolite through a series of tiered metabolic processes (Fig. 2.3; Baumgartner et al., 2006). These processes break down organic matter and release organically bound phosphate into pore fluid to promote phosphogenesis (Föllmi, 1996). Some heterotrophic bacteria can also incorporate large amounts of intracellular polyphosphate under oxic conditions, particularly sulfide-oxidizing bacteria Beggiatoa, Thioploca, and

Thiomargarita (Schulz and Schulz, 2005; Crosby and Bailey, 2012; Bailey et al., 2013).

Under anoxic conditions, these bacteria release orthophosphate, which acts as a nucleation site for the rapid precipitation of francolite (Bailey et al., 2013).

Another prerequisite for phosphogenesis was the ability of these microbial consortia to seal phosphate (Brasier and Callow, 2007). Microbes very effectively prevent the escape of phosphate from sediment by reincorporating liberated phosphate that would otherwise escape to seawater. Further sealing occurs when microbial layers are buried by fine-grained sediment, which also acts as an effective barrier to diffusion of phosphate released through the microbial degradation of organic matter.

Finally, stromatolite growth morphology was apparently an important factor controlling phosphogenesis. Intertidal stromatolites from the Paleoproterozoic Aravallian

Supergroup of Udaipur, Rajasthan, India, also contain pristine phosphorite (Banerjee,

1971; Banerjee et al., 1980). The nucleation rate of francolite precipitating in these stromatolites is interpreted to have been controlled by the amount of intercolumnar space, packing tightness of stromatolitic laminae, and stromatolite growth morphology

(Banerjee, 1971).

In the Nova America member, the occurrence of pristine phosphorite in intertidal stromatolitic biostromes is interpreted to reflect these processes (Fig. 9.2). Most important are the microbial metabolic processes that released and then sealed phosphate in the stromatolite. Thin, fine-grained tidal deposits interbedded in biostromes probably acted as a secondary barrier to phosphate diffusing out of the stromatolite (Fig. 9.2). The subtidal stromatolitic reefs are probably devoid of phosphorite because the internal structure of the stromatolite was not conducive to phosphogenesis. Unlike the intertidal stromatolites, fine-grained sediment that could act as a phosphate seal was continually swept away in this tide-dominated setting. Recycling of phosphate out of stromatolitic reefs in this “leaky” subtidal system may have promoted continued primary productivity and stromatolitic growth (Fig. 9.2).

Petrographic and geochemical evidence indicates that subsequent burial and hydrothermal alteration of intertidal stromatolites increased their phosphate concentrations from <15 wt.% P2O5 up to 25 wt.% (Misi and Kyle, 1994). Stromatolites underwent a complex paragenesis that includes pervasive post-depositional dolomitization and precipitation of the secondary phosphatic minerals fluorapatite and hydroxylapatite from hydrothermally remobilized P to produce economic phosphorite.

9.2 The Neoproterozoic P-cycle

The Neoproterozoic represents an important transition in Earth's history as the deep oceans became ventilated, creating conditions ideal for the evolution of multicellular animals during the Ediacaran radiation (Och and Shields-Zhou, 2012;

Pufahl and Hiatt, 2012; Drummond, 2014). The expansion of phosphogenic

sensitive releaseinto processes phosphate microbial -

d with fine grained muddy deposits. This sedimentation pattern seals stromatolites muddy seals into grained This deposits. pattern the d with fine sedimentation Phosphogenesis model for the Nova America member. The subtidal member. is America by continued The dominated environment Phosphogenesis the Nova model for 9.2 recycling of phosphate in subtidal stromatolite reefs. In contrast, the intertidal stromatolites, which are shaped by tidal the intertidal which are stromatolites, phosphatecontrast, of In in recycling subtidal reefs. stromatolite interbedde are processes, redox phosphate sediment where and traps in the anoxic sediment fluid, promoting phosphogenesis. pore Figure Figure

97 environments during the NOE is interpreted to have assisted in preconditioning benthic environments with bioessential P for this diversification (Drummond et al., in press).

Phosphogenesis, although not directly dependent on the Eh state of precipitating fluids, is mediated by redox-sensitive microbial processes that concentrate phosphate within pore fluid. These microbial reactions are generally considered the most important processes for concentrating P in organic-rich sediment accumulating beneath sites of active coastal upwelling (Glenn et al., 1994; Pufahl, 2010). In non-upwelling areas with an oxygenated seafloor, Fe-redox cycling concentrates pore water phosphate (Heggie et al., 1990; Glenn et al., 1994; Pufahl, 2010). This cyclic mechanism releases phosphate adsorbed onto Fe-(oxyhydr)oxides when burial occurs beneath the Fe-redox interface.

Throughout much of the Precambrian an anoxic water column precluded phosphorite from forming except where photosynthetic oxygen oases along the coast produced a suboxic seafloor that pushed these redox-sensitive phosphogenic processes into the sediment to promote the precipitation of authigenic francolite (Nelson et al., 2010; Pufahl and Hiatt, 2012; Drummond et al., in press). Phosphogenesis was precluded in deeper- water environments because these processes were suspended in the water column (Nelson et al., 2010; Pufahl and Hiatt, 2012). Only when the water column became fully oxygenated at ca. 580 Ma during the height of the NOE (Och and Shields-Zhou, 2012) did phosphogenic environments expand to create the first true phosphorite giants (Nelson et al., 2012; Drummond et al., in press). The Doushantuo phosphorites, ca. 570 Ma, of the Yangtze Platform in southern China represent this transition and expansion of phosphogenic environments (Shen et al., 2000)

Results presented here provide new information regarding the nature of the P- cycle during the onset of the NOE and Earth’s second major phosphogenic episode. The close association of phosphorite with intertidal stromatolites suggests that the presence of oxygen was not the only prerequisite for phosphogenesis in Precambrian coastal environments. Results imply that the benthic P-cycle was more complex than previously surmised and emphasize the multifaceted significance of microbial processes and their relationship to depositional environments. In addition to Eh, this microbe-sediment feedback concentrates bioavailable P at the seafloor and facilitates the precipitation of authigenic francolite.

In energetic subtidal environments, where stromatolitic patch reefs developed, P was recycled to seawater, preventing francolite precipitation. Such microbial recycling in deeper environments allows increased primary production (Goldberg and Shields, 2006;

Filippelli, 2011), which may have contributed to water-column oxygenation on a global scale during the NOE. As oxygen levels increased, the environments hosting eukaryotes also expanded (Rasmussen et al., 2008). Eventually a shift from a cyanobacteria- dominated to a eukaryote-dominated biological pump affected the rate of organic matter accumulation, and thus the pathway of organic carbon and P to the sediment (Lenton et al., 2014). This shift in oxygen demand is thought to have assisted in driving the oxygenation of the deep ocean that was critical for the eventual evolution of multicellular animals (Lenton and Watson, 2004; Lenton et al., 2014).

CHAPTER 10: CONCLUSIONS

1) Economic phosphorite in the Salitre Formation is associated only with intertidal stromatolitic biostromes that formed the tops of peritidal carbonate cycles in the

Neoproterozoic Salitre Formation. Peritidal cycles are formed of basal subtidal, herringbone cross-stratified grainstone (F1, F2) and hemispheroidal columnar stromatolitic patch reefs (F4) that grade upward into intertidal flaser to lenticular-bedded packstone (F3) and interbedded stromatolitic biostromes (F5).

2) The stratigraphic stacking of preserved peritidal cycles suggests that deposition occurred during the latter stages of the LST and into the early TST. Lithofacies relationships also indicate that deposition occurred on an arid epeiric ramp that was influenced by tides and storms. The abundance of evaporites suggests that the abiotic precipitation of peritidal carbonate was facilitated in part by the evaporitic concentration of seawater.

3) Phosphogenesis was restricted to the nearshore because stromatolitic biostromes that colonized intertidal flats created the necessary conditions and chemical gradients in the sediment for the authigenic precipitation of francolite. Cyanobacteria and associated heterotrophic bacteria in the stromatolite interior can actively store, and in some cases, release P. Other redox-sensitive microbial processes degraded this P-enriched organic matter in the interior of the stromatolite and concentrated phosphate and precipitated francolite. Also important for phosphogenesis were the sealing effects of interbedded, fine-grained tidal deposits and storm layers, and the nature of the intercolumnar space

100 and tight packing of the small columns forming biostromes. Energetic subtidal environments where stromatolitic patch reefs developed promoted recycling of P back to seawater, preventing francolite precipitation.

4) Petrographic and stable isotopic data (δ13C and δ18O) indicate that these peritidal carbonates underwent a complex paragenesis. Meteoric and burial diagenesis altered primary mineralogy and created secondary porosity. Subsequent hydrothermal alteration resulted in pervasive dolomitization, occlusion of secondary porosity, and the remobilization of P to precipitate secondary phosphatic minerals. Such enrichment of intertidal facies created economic phosphorite. Paragenetic relationships also indicate that stable isotopic signatures were continually reset during alteration, rendering them unusable for chemostratigraphic correlation with other Neoproterozoic successions.

5) Finally, the Salitre Formation provides important insight into the Neoproterozoic

P-cycle during the onset of Earth’s second major phosphogenic episode. As in the

Paleoproterozoic, presented sedimentologic evidence indicates phosphogenesis was restricted to the coast in association with photosynthetic oxygen oases. The creation of suboxic conditions allowed important redox sensitive phosphogenic processes to concentrate P in stromatolites. Results of the present study also suggest that the late

Neoproterozoic benthic P-cycle was more complex than previously surmised, and emphasize the multifaceted significance of microbial processes. In the Ediacaran, these redox-sensitive phosphogenic processes expanded into progressively deeper

101 environments as the oceans became fully oxygenated to produce the Earth’s first true phosphorite giants.

102

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APPENDIX I: X-RAY DIFFRACTION (XRD) DATA

Figure AI.1 XRD data showing deep burial calcite vein intruding ferroan dolomite.

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Figure AI.2 XRD data of an intertidal deposit with ferroan dolomite and coated strontian fluorapatite grains.

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Figure AI.3 XRD data of pristine phosphate of F5 composed of fluorapatite, which have been secondarily enriched in phosphate.

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Figure AI.4 XRD data of minerals associated with extensive hydrothermal alteration especially the presence of hydroxylapatite.

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Figure AI.5 XRD data of pristine phosphate of F5 composed of fluorapatite, which have been secondarily enriched in phosphate.

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Figure AI.6 XRD data of pristine phosphate of F5 composed of fluorapatite, which have been secondarily enriched in phosphate.

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Figure AI.7 XRD data of subtidal deposit subjected to dolomitization with the presence of burial calcite veins.

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Figure AI.8 XRD data of a dolomitized subtidal deposit with abundant intraclasts and the presence of detrital sediment.

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Figure AI.9 XRD data of a dolomitized subtidal deposit.

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Figure AI.10 XRD data of pristine phosphate of F5 composed of fluorapatite, which have been secondarily enriched in phosphate.

132

Figure AI.11 XRD data of a completely dolomitized subtidal lithofacies.

133

Figure AI.12 XRD data of a completely dolomitized subtidal lithofacies.

134

Figure AI.13 XRD data of pristine phosphate of F5 composed of fluorapatite, which have been secondarily enriched in phosphate from hydrothermal alteration.

135

Figure AI.14 XRD data of pristine phosphate of F5 composed of fluorapatite, which have been secondarily enriched in phosphate.

136

Figure AI.15 XRD data of pristine phosphate of F5 composed of fluorapatite, which have been secondarily enriched in phosphate.

137

APPENDIX II: STABLE ISOTOPE GEOCHEMISTRY

Table AII.1: Carbon and oxygen stable isotopic analysis.

18 13 Lithofacies Sample No. δ OVPDB (‰) δ CVPDB (‰)

F1 FNC 05 - 46.60 -2.7 2.7

F1 FNC 05 - 46.80 -2.9 3.1

F1 FNC 05 - 46.81 -2.8 2.9

F1 FNC 05 - 49.05 -3.8 7.7

F1 FNC 05 - 50.85 -2.8 2.6

F1 FNC 05 - 53.45 -2.1 1.3

F1 FNC 05 - 58.20 -2.8 0.9

F1 FNC 08 - 3.40 -4.2 7.5

F1 FNC 08 - 30.30 -2.6 3.0

F1 FNC 10 - 1.80 -3.8 8.4

F1 FNC 10 - 1.81 -3.8 8.4

F1 FNC 11 - 39.40 -2.6 4.6

F1 FRS 02 - 1.00 -7.4 -3.2

F2 FNC 02 - 46.00 -3.2 3.1

F2 FNC 05 - 3.65 -3.3 8.4

F2 FNC 05 - 12.75 -4.3 8.2

F2 FNC 05 - 13.30 -4.5 8.4

F2 FNC 19 - 16.70 -3.2 8.0

F3 FNC 01 - 15.90 -3.2 7.0

138

F3 FNC 01 - 28.10 -2.4 4.6

F3 FNC 01 - 28.11 -2.2 4.6

F3 FNC 01 - 28.20 -1.9 4.1

F3 FNC 02 - 4.50 -3.9 8.5

F3 FNC 02 - 19.87 -4.7 10.0

F3 FNC 02 - 44.00 -3.3 4.1

F3 FNC 02 - 48.88 -2.7 2.9

F3 FNC 18 - 30.20 -4.1 6.4

F3 SCM 01 - 30.00 -5.5 0.3

F3 SCM 01 - 31.00 -5.6 0.2

F3 FNC 01 - 24.60 -3.2 4.7

F3 FNC 01 - 32.60 -2.4 5.6

F3 FNC 02 - 31.60 -2.8 5.3

F3 FNC 02 - 33.20 -3.4 4.9

F3 FNC 02 - 34.25 -4.7 4.4

F3 FNC 02 - 38.25 -3.2 5.3

F3 FNC 02 - 40.15 -2.4 5.9

F3 FNC 05 - 7.05 -2.2 8.7

F3 FNC 05 - 21.10 -2.2 4.0

F3 FNC 18 - 5.50 -3.3 8.4

F3 FNC 18 - 5.55 -3.3 8.3

F3 FNC 05 - 24.30 -2.0 5.2

F3 FNC 05 - 30.20 -3.4 4.7

F3 FNC 05 - 42.90 -2.7 3.0

139

F3 FNC 05 - 42.91 -2.7 2.9

F3 SCM 01 - 2.00 -0.5 -0.1

F3 SCM 01 - 2.10 -0.6 0.0

F3 VB 01 - 1.00 -10.2 -5.7

F4 FNC 02 - 23.25 -3.0 7.3

F4 FNC 02 - 24.85 -3.6 6.4

F4 FNC 02 - 28.25 -3.8 4.8

F4 FNC 02 - 31.20 -3.0 4.7

F4 FNC 05 - 38.25 -2.9 3.5

F4 FNC 10 - 21.60 -6.5 -2.1

F4 FNC 18 - 3.00 -3.0 7.7

F4 FNC 18 - 3.01 -2.9 7.6

F4 FNC 18 - 26.36 -3.5 7.8

F4 FNC 18 - 28.96 -3.4 6.6

F4 FRS 05 - 0.00 -4.5 -3.3

F4 FRS 05 - 1.00 -4.6 -3.3

F5 FNC 02 - 35.35 -3.7 5.1

F5 FNC 02 - 36.25 -3.9 5.0

F5 FNC 05 16.95 -3.1 5.1

F5 FNC 02 - 25.35 -5.0 4.0

F5 FNC 02 - 35.25 -3.9 5.4

F5 FNC 05 - 15.30 -3.4 5.2

140

F5 FNC 05 - 15.31 -3.3 5.3

F5 FNC 05 - 16.30 -3.4 5.0

F5 FNC 11 - 26.80 -3.1 4.7

F5 GAL 02 - 14.00 -3.5 1.5

F5 GAL 02 - 16.00 -6.7 -6.5

F5 GAL 02 - 16.10 -6.8 -6.4

VIENS V 01 - 7.30 -7.6 -7.0

VIENS V 02 - 48.90 -7.0 -0.6

VIENS V 02 - 49.70 -5.9 -4.6

VIENS V 02 - 49.71 -6.8 -6.4

VIENS V 11 - 19.60 -6.7 2.5

VIENS V 18 - 30.20 -5.4 5.6

CALCRETE CAL 01 - 1.00 -4.3 -8.6

CALCRETE CAL 01 - 2.00 -4.2 -8.5

CALCRETE CAL 02 - 1.00 -5.3 -8.9

CALCRETE CAL G1 - 4.00 -4.2 -9.2

CALCRETE CAL G1 - 6.00 -4.3 -9.0

CALCRETE CAL G2 - 11.00 -6.2 -8.0

CALCRETE CAL G2 - 18.00 -3.7 0.6

141