Microbe-Mineral Interactions During Exceptional Fossil Preservation, Stromatolite Formation, and Desert Varnish Growth

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Michael Strange

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This Dissertation has been accepted for inclusion in UNLV Theses, Dissertations, Professional Papers, and Capstones by an authorized administrator of Digital Scholarship@UNLV. For more information, please contact [email protected]. MICROBE-MINERAL INTERACTIONS DURING EXCEPTIONAL FOSSIL

PRESERVATION, STROMATOLITE FORMATION,

AND DESERT VARNISH GROWTH

Michael A. Strange

Bachelor of Arts — Geology Utah State University 2014

A dissertation submitted in partial fulfillment of the requirements for the

Doctor of Philosophy – Geoscience

Department of Geoscience College of Sciences The Graduate College

University of Nevada, Las Vegas December 2020

Dissertation Approval

The Graduate College The University of Nevada, Las Vegas

December 21, 2020

This dissertation prepared by

Michael A. Strange entitled

Microbe-Mineral Interactions During Exceptional Fossil Preservation, Stromatolite Formation, and Desert Varnish Growth is approved in partial fulfillment of the requirements for the degree of

Doctor of Philosophy–Geoscience Department of Geoscience

Stephen Rowland, Ph.D. Kathryn Hausbeck Korgan, Ph.D. Examination Committee Chair Graduate College Dean

Douglas Sims, Ph.D. Examination Committee Member

Ganqing Jiang, Ph.D. Examination Committee Member

Minghua Ren, Ph.D. Examination Committee Member

Dennis Bazylinski, Ph.D. Graduate College Faculty Representative

ABSTRACT MICROBE-MINERAL INTERACTIONS DURING EXCEPTIONAL FOSSIL

PRESERVATION, STROMATOLITE FORMATION,

AND DESERT VARNISH GROWTH

Michael A. Strange

University of Nevada, Las Vegas, 2020

Dissertation Chair: Stephen M. Rowland

The Ediacaran to Cambrian Deep Spring Formation consists of mixed carbonate- siliciclastic strata which contain an increasingly complex and biogeographically important biota. Past investigations of the Deep Spring Formation at Mt. Dunfee, Nevada, explored the highly diverse microbialite reefs consisting of a wide range of stromatolite morphologies which exerted significant control on local sedimentation and topography. Early investigations also documented the biomineralizing metazoan Cloudina (an Ediacaran index fossil). However, recent exploration of the area has resulted in the discovery of several new metazoan fossil communities consisting of a diverse assemblage of Ediacaran soft-tissue tubicolous vermiforms (tube fossils) similar to Cloudina. The focus of this dissertation is on the geomicrobiological processes at work in the Deep Spring Formation and elsewhere which result in the preservation of soft-tissues (Ch. 2), the dissolution/precipitation of feldspars within stromatolites (Ch. 3), and the early formation of modern desert varnish (Ch. 4).

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Chapter 1 of my dissertation examines the stratigraphic and paleontologic context of the Deep Spring Formation. In this chapter, I summarize the current understanding of the paleontological transition at the Ediacaran-Cambrian boundary in the western United States. The Death Valley and White-Inyo-Esmeralda regions of California and Nevada contain some of the most fossiliferous terminal Neoproterozoic to early Phanerozoic strata in the world. Within these strata, the Cambrian explosion is in full display, resplendent with diverse trace and body fossils. The Wood Canyon and Deep Spring Formations, within which the boundary is contained, are comprised of mixed carbonate and siliciclastic lithologies—allowing for a more in-depth understanding of this critical transition in the history of life. Here, I briefly review our current state of understanding of life contained within these rocks. This chapter has been published (Strange and Rowland, 2017) Chapter 2 consists of a manuscript I submitted to the journal Geobiology describing the microbial-related mineralization pathways responsible for preserving the upper fossil horizon in the Deep Spring Formation at Mt. Dunfee, Nevada. The manuscript was recently rejected, with an offer to resubmit after reviewers’ comments have been addressed. In this manuscript, I describe a new mineralization pathway to exceptional fossil preservation which involves the metabolic activities of an iron- oxidizing bacterial community. This model expands on previous microbial/redox zonation models of exceptional preservation and requires rapid emplacement of organic matter into the microaerophilic sedimentary environment. Within these redox conditions, iron-oxidation by bacteria produce abundant Fe (III) from porewater sourced Fe (II), resulting in the precipitation of early Fe (III) mineral phases with the quick stabilization in the form of goethite and iron-rich aluminosilicates. I have named this taphonomic pathway “ferrumation”. Ferrumation incorporates well into the microbial/redox zonation

model of Schiffbauer et al. (2014) due to the occurrence of pyritization within the fossil assemblage. Chapter 3 moves away from iron-oxidizing bacteria to the mineral producing capabilities of the consortia of microbes involved in stromatolite formation. Feldspar assemblages found within stromatolite laminae differ between stromatolite localities and from the surrounding interstromatolite zone between columnar stromatolites. This suggests that microbial activities may be influencing their dissolution and precipitation. Textural associations between feldspars, quartz, and calcite are similar to the micrographic textures seen in some igneous rocks which result from the co-stability of feldspar and quartz in a magma chamber. These highly unusual feldspar textures within Deep Spring stromatolites led to the hypothesis that microbial activity resulted in the dissolution of detrital feldspars (mostly orthoclase) captured by the sticky surface of the stromatolite. The capture of dissolved ions from detrital orthoclase by organic molecules become released into the closed system of a mineralizing microbialite and precipitate out as authigenic albite, quartz, Ca-plagioclase, and a K-rich aluminosilicate, or their poorly crystalline precursors. This pattern of feldspar-quartz-calcite associations has been found at the Ediacaran-Cambrian Mount Dunfee section and the Molly Gibson Mine section of the Deep Spring Formation, in Miocene stromatolites from the Duero Basin, Spain, and also in a stromatolite from the Flinders Range, Australia. This research is significant because it could lead to the ability to isolate and date authigenic albite grains using the

Ar/Ar method, potentially opening the utilization of stromatolites as abundant repositories of syndepositional authigenic minerals for geochronology. Chapter four departs from the Ediacaran-Cambrian boundary but continues with the theme of microbe-mineral interactions. This chapter details the earliest phases of desert varnish formation on ephemeral wash sediments from Nelson, Nevada. I show that

the early development of incipient desert varnish occurs through small “microdots”. These are likely associated with Fe- and Mn-oxidizing microbial communities on the surface of the sediment grains. The Fe-oxide composition of some microdots and the mixed-composition of more advanced desert varnish suggests that desert varnish may accumulate through the cyclic deposition of layers during periods of habitability of each microbial community. Desert varnish microdots are shown to have progressive mineralization from smooth surfaces to micronodules, while advanced development becomes botryoidal. The previously reported “microstromatolitic” growth habit of some desert varnish seems to correspond with the microdot morphology of the incipient varnish found on these grains.

ACKNOWLEDGMENTS

I would like to sincerely thank the amazing geoscience faculty at the University of Nevada, Las Vegas, for their invaluable guidance throughout this process. In particular, I would like to thank my advisor Dr. Steve Rowland for his mentorship, guidance, and encouragement, and Dr. Minghua Ren for his patient training and insightful conversations. This dissertation would not have been possible without their consistent professionalism, experience, and intellect. I thank both of them for indulging me all these years while I explore the world in scientific curiosity. I would also like to thank my wife May Strange and daughter Rose Strange for providing emotional support throughout these difficult years. They have given me purpose beyond imagination and nearly insurmountable distraction at times. Without them, I may have given up years ago. I would also like to thank my parents for encouraging me to pursue my dreams and for their years of support while I finish school.

I thank my dissertation committee members, Dr. Ganqing Jiang, Dr. Douglas Sims, and Dr. Dennis Bazylinski, for being consistent sources of encouragement and assistance. I thank all the people who have spent hours collecting fossils in the field with me, including Jeremy Strange, Nate Strange, Greg Parsegov, Gretchen O’Neil, Dr. Tera Selly, Dr. Jim Schiffbauer, and Dr. Emmy Smith. I very much appreciate the stromatolite samples which were provided to me by to Carol Dehler and I hope to repay her generosity in the future.

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TABLE OF CONTENTS

ABSTRACT ...... iii

ACKNOWLEDGMENTS ...... vii

TABLE OF CONTENTS ...... viii

LIST OF FIGURES ...... xi

CHAPTER 1: THE EDIACARAN-CAMBRIAN BIOTIC TRANSITION IN EASTERN CALIFORNIA AND SOUTHWESTERN NEVADA AND ITS GLOBAL CONTEXT ...... 1

INTRODUCTION ...... 1 REGIONAL STRATIGRAPHY ...... 2 A WORMWORLD FAUNA IN THE GARDEN OF EDIACARA ...... 5 BIOTIC CHANGE RECORDED IN THE TRACE FOSSIL RECORD ...... 8 CONCLUSIONS...... 12

CHAPTER 2: DISENTANGLING PYRITIZATION FROM DIRECT IRON OXYHYDROXIDE AND IRON-RICH CLAY TEMPLATING IN EXCEPTIONAL FOSSIL PRESERVATION OF EDIACARAN CLOUDINOMOPRHS ...... 20

ABSTRACT ...... 20 INTRODUCTION ...... 22 A Unified Taphonomic Model ...... 24 The Powerful Role of Ferruginous Conditions ...... 26 METHODS ...... 28 Specimen collection & selection ...... 28 Electron probe microanalysis ...... 30 Scanning Electron Microscopy (SEM) ...... 31 Raman spectroscopy ...... 32 Data analysis ...... 33 RESULTS ...... 34 Taphofacies Description & Fossil Identification ...... 34 Mineral Morphology ...... 36 Geochemical Characterization of Specimens ...... 37 DISCUSSION ...... 40 Taphonomic Pathways at Play in the Deep Spring Formation ...... 40 Taphonomic Model Description and Justification ...... 43 Mineral Diagenesis ...... 53 CONCLUSIONS...... 55

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CHAPTER 3: MICROBIALLY MEDIATED DISSOLUTION OF DETRITAL ORTHOCLASE AND ITS REPRECIPITATION AS AUTHIGENIC AND COMPOSITIONALLY DIVERSE PLAGIOCLASE—A POTENTIAL MICROBIALITE DATING TECHNIQUE? ...... 66

ABSTRACT ...... 66 INTRODUCTION ...... 68 General Background ...... 68 Microbial Dissolution of silicates and the resulting precipitates ...... 70 Microbialites ...... 72 Microbial Dissolution ...... 75 Microbially-induced precipitation of silicates ...... 77 Authigenic feldspars in carbonate rocks ...... 79 Dating of authigenic feldspars in sedimentary rocks ...... 81 METHODS ...... 82 Sample collection and preparation ...... 82 Scanning Electron Microscopy (SEM) ...... 83 Electron Probe Microanalysis ...... 84 Mineral Separation ...... 85 PHREEQC Modeling ...... 86 RESULTS ...... 87 Geochemistry ...... 87 Unusual Mineralogical Features ...... 89 Other Stromatolite Localities ...... 94 Grain separation ...... 97 DISCUSSION ...... 99 Alternative Hypothesis for the Origin of Micrographic-like Textures...... 99 Stromatolite-authigenesis Model ...... 104 Stromatolite-Authigenesis Model Justification...... 112 Ar/Ar Dating ...... 116 CONCLUSIONS...... 118

CHAPTER 4: THE EARLY STAGES OF STROMATOLITIC GROWTH OF ROCK VARNISH ...... 133

ABSTRACT ...... 133 INTRODUCTION ...... 134 Geological Setting ...... 136 Biogenic and Abiogenic Contributions in Varnish Formation ...... 137 METHODS ...... 141 RESULTS ...... 142 Early Stages of Rock Varnish Formation ...... 143 Advanced Stage Varnish ...... 145

DISCUSSION ...... 147 Incipient Desert Varnish ...... 147 Unusual Compositional Aspects ...... 149 Clay Portion of Incipient Varnish ...... 152 Stromatolitic Growth ...... 153 CONCLUSIONS...... 155

REFERENCES ...... 162

CURRICULUM VITAE ...... 194

LIST OF FIGURES Figure 1.1 Generalized stratigraphic columns of Ediacaran and Cambrian strata in eastern California. Adapted from Hagadorn et al. (2000) and Corsetti and Hagadorn (2003)...... 13 Figure 1.2 Cloudinomorph fossils from the upper fossil horizon at the Mount Dunfee section of the Deep Spring Formation. (A, B, C) Tubular vermiform fossils (Cloudinids and Cloudinid-like forms) from the Esmeralda (middle) Member of the Deep Spring Formation, Mount Dunfee, Nevada. (D) hypothesized reconstruction of recalcitrant tube wall and unpreserved, fan-shaped feeding structures of Conotubus ...... 14 Figure 1.3 Trace fossils typical of the Lower Assemblage in the Death Valley region, consisting of simple, shallow burrows. (A) Treptichnus pedum from the Wood Canyon Formation. (C-D) Treptichnus pedum from the Deep Spring Formation. (B, E) Simple, shallow traces from the Wood Canyon Formation in the Montgomery Mountains, Nye County, Nevada ...... 15 Figure 1.4 Trace fossils typical of the Middle Assemblage in the Death Valley region, consisting of complex, mostly shallow burrows. (A) Large undifferentiated burrows. (B) Rusophycus, a trilobite burrow. (C) Small undifferentiated burrows from slab in “5E.” (D) Bilobed Helminthoidichnites burrow from slab in “5E.” (E) slab showing diversity of burrows typical of the Middle Assemblage with arrows indicating where magnified images were taken for images “5D” and “5E”...... 16 Figure1.5 Traces and body fossils typical of the Upper Assemblage of the Death Valley region comprised of large, complex, and deep burrows. (A) Harlaniella. (B) dense cluster of Skolithos-like vertical burrows. (C) Magnified image of trilobite cephalon in “F.”(D-E,H) large, complex, vertical burrows. (F) slab containing abundant trilobite hash. (G) large slab containing highly diverse traces Rusophycus, Treptichnus pedum, and Skolithos...... 17 Figure 1.6 Biostratigraphic ranges of various generalized groups within the Death Valley region. Representative trace fossils in each assemblagae are illustrated in Figs. 1.3, 1.4, and 1.5. Schematic reconstructions of each assemblage are shown in Fig. 1.7. Note the trend toward deeper and more complex trace fossils through time ...... 18 Figure 1.7 Reconstructions of three fossil assemblages in the Death Valley region. In the middle and upper assemblages, emphasis has been placed on trace fossil diversity. See Fig. 1.6 for graphical representation of assemblages ...... 19 Figure 2.1 Stratigraphic column of the upper Reed Dolomite and Deep Spring formations at Mount Dunfee, Esmeralda County, Nevada. Modified from Smith et al., (2016)...... 57

Figure 2.2 Thin sections of two fossiliferous horizons containing tubicolous fossils from the upper fossil horizon of the Deep Spring Formation. Taphofacies 1 (a-f) contains less common, high fidelity fossils within organic-poor sediments, while taphofacies 2 (g-j) contains higher organic content and abundant, poorly preserved fossils; a) polished cross section of taphofacies 1, arrow indicates fossil photographed in (b); b) Fossil from taphofacies 1 [arrow in (a)] photographed submerged in alcohol; c-e) polished cross sections of tubicolous fossil in taphofacies 1 photographed in the thin section from (a); f) Cross section of tubicolous fossils from taphofacies 1 from (a) imaged using BSE; g) polished cross section of taphofacies 2 showing small zone of dense preservation (orange color); h) several examples of typical fossils from taphofacies 2; i) magnified image of zone of dense preservation for taphofacies 2; j) abundant balls of iron oxyhydroxides found throughout the sediment of taphofacies 2, seen as bright spots in Electron BSE imaging. Stars indicate plane of splitting during collection of specimens that are filled with epoxy...... 58 Figure 2.3 Microprobe elemental maps and point analysis of specimen DSM-200 (LVNHM BLM 2020.001). a) reflected light photograph showing real colors; b) Iron x-ray map; c) Iron phase map using x-ray map from (b), orange indicates goethite [identified by Raman spectroscopy (Fig. 8)] and green indicates iron-rich clay; d) Iron x-ray map; e) Iron phase map from “d” with same color classifications as (c), note minor amounts of goethite dispersed throughout clay; f) Microprobe data points across red line in b, c, d, and e...... 59 Figure 2.4 Microprobe elemental maps of specimen DSM-300 (LVNHM BLM 2020.001). a) BSE image after polishing of specimen; b-e) elemental maps constructed using electron microprobe showing relative intensities of Fe (b), Al (c), Si (d), and S (e); f) Fe phase map showing distribution of hematite (purple) and an unidentified Fe (II)-phase. Note conspicuous absence of iron-rich clays associated with the fossil...... 60 Figure 2.5 Progressive mineralization seen in cloudinomorph specimen DSM-210 (LVNHM BLM 2020.001). a) polished specimen submerged in alcohol and imaged under fluorescent light showing parts of the specimen that are polished and under epoxy; b) section of DSM-210 showing three zones of mineralization (1 being most pervasively mineralized and 3 being least); c) detailed observations of mineral templating of microbial cells within the tube walls of DSM-210, ranging from partial coating (zone 3) to pervasively-infilled (zone 1. Note the second layer forming within the enclosed microspherules (arrow)...... 61 Figure 2.6 Light and BSE images showing variation of Fe-minerals in tube fossil specimens DSM-200 (a-c) and DSM-300 (d-f). The surrounding matrix in all images was made up of aluminosilicates with highly variable amounts of iron; a) goethite micronodules conglomerating into massive buildup goethite at the tube wall; b) goethite micronodules and pervasively-infilled microspherules within the iron rich aluminosilicate clay matrix (see more of this morphology in Fig. 4); c)

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morphological difference between iron-rich clay (large, platy) and iron-poor clay (fibrous); d) massive hematite at the tube wall; e,f) hematite pseudomorphs of pyrite rhombs (arrows) within iron-poor clay matrix. Note rhomb pseudomorph amalgamation in (f)...... 62 Figure 2.7 Hyolithid (UNLV-LSS-001) from the Cambrian Spence Shale Member, Langston Fm, of Northern Utah. a) Reflected light photograph of Hyolithes cerops (no color alteration). Sample has been polished near center of the hyolithid and the remainder was under epoxy; b-d) BSE images of oxidized pyrite crystals. Brighter areas represent pyrite, medium areas represent iron oxides, and darker areas represent clays...... 63 Figure 2.8 Non-Metric Multidimensional Scaling (NMDS) plots of Microprobe data and Raman spectra of Deep Spring Formation specimens DSM-200 (red dots/lines), DSM-300 (green dots/lines), and a Spence Shale hyolithid (blue dots) with interpretation. a) NMDS plot made using FeO, SiO2, Al2O3, and SO3 weight percentages, with a table of four points to illustrate differences in chemistry; b) Raman spectra of DSM- 200 (red) and DSM-300 (green) showing important peaks used for mineral identification...... 64 Figure 2.9 Iron biogeochemical cycle in relation to phase precipitation based on the sediment/microbial zonation model showing sources of sedimentary iron phases and zones of preservation. Addition of the microaerophilic zone into the model provides a small horizon for dynamic cycling of iron phases within the sedimentary column. Ferrumation would dominate the mineralization pathways within the microaerophilic zone...... 65 Figure 3.1 Stromatolites of the Mount Dunfee section of the Deep Spring Formation. A) stromatolitic reef; B) Stromatolite reef horizons; C) stromatolite collected from float which contains both stromatolite laminae (left, gray) and the interstromatolite zone (right, orange); D) regional map showing the location of the Mount Dunfee Section and the Molly Gibson Mine section of the Deep Spring Formation; E) Stromatolite sample (red circle) collected from stromatolite horizon A. F) stromatolite sample (red circle) collected from stromatolite horizon D...... 119 Figure 3.2 Stromatolites of the Molly Gibson Mine section of the Deep Spring Formation. A) field of stromatolites; B) Detailed picture of the stromatolites; C) stromatolite sample (red circle) collected from the Molly Gibson Mine section; D) regional map showing location of the Mount Dunfee section and the Molly Gibson Mine section of the Deep Spring Formation...... 120 Figure 3.3 Representative mineral assemblage of the Tieling Formation stromatolite. A) Backscatter Electron image of the polished stromatolite sample; B) Potassium elemental map of section shown in “A”; C) Silicon elemental map of the section shown in “B”; D) Sodium elemental map of the section shown in “A”. Qz = quartz, C = Calcite, Kspar = orthoclase...... 121

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Figure 3.4 Representative mineral assemblage of the Baicalia burra stromatolite sample from the Flinders Range, Australia. A) Backscatter Electron image of the polished stromatolite sample; B) Potassium elemental map of section shown in “A”; C) Silicon elemental map of the section shown in “B”; D) Sodium elemental map of the section shown in “A”. Qz = quartz, C = Calcite, Kspar = orthoclase, Ab = albite...... 122 Figure 3.5 Mineral Textures within the Duero Basin stromatolite. A) Diagonally oriented grain shows the wormy opal textures (which are composed of pure SiO2 in EDS spectrum) and associated elemental maps. Note the higher amount of alteration at the bottom left corner of the grain; B) orthoclase grain with a partial rim of Mg-Si organic matrix with little to no dolomite mineralization; C) broader view of grain in “A” showing the distribution of altered quartz grains within the Mg-Si organic matrix with associated dolomite precipitates. Dol = dolomite, Kspar = orthoclase, Mg-Si = amorphous Mg-silicate within organic matrix, Qz = quartz...... 123 Figure 3.6 Stromatolite sample preparation and SEM analysis...... 124 Figure 3.7 Geochemistry of the Deep Spring mineral assemblage. A) ternary diagram showing feldspar chemistry; B) SEM/BSE image with interpreted mineral assemblage; C) Calculated mineral formulae for typical minerals of each variety; D) typical probe dataset for each mineral...... 125 Figure 3.8 Mineral textures found within the stromatolite found in float (Figure 3.1C, 3.6) at the Mount Dunfee Section of the Deep Spring Formation. A)-C Examples of albite-rimed orthoclase grains from polished thin section; D) albite- rimed orthoclase grain from maceration residuals showing etch pits; E-G) SEM/BSE images showing examples of calcite captured within albite (Ab) from polished thin section; H-I)SEM/BSE images showing examples of albite (Ab)- rimed Ca-plagioclase grains with carbonate capture (Cc); J-K) SEM/BSE images showing examples of micrograins of quartz (Qz) and albite (Ab) within the carbonate matrix (C) within the stromatolite...... 126 Figure 3.9 SEM/BSE images showing examples of the complex albite-Ca- plagioclase-quartz-calcite mineral assemblages within the Molly Gibson Mine section of the Deep Spring Formation. A-C) complex mineral assemblages with SEM/EDS data showing amounts of silicon (Si), sodium (Na), calcium (Ca), and aluminum (Al); D) SEM/BSE image showing mica grain with associated albite- quartz-calcite grain with micrographic textures; E) SEM/BSE image showing mica grain and associated quartz and albite grains; F) SEM/BSE image showing small quartz grain with internal Ca-plagioclase (light gray) and nearby albite...... 127 Figure 3.10 SEM/BSE images of micrographic-like textures and associated elemental maps with phase interpretation from the Mount Dunfee section of the Deep Spring Formation. A-B) albite-calcite micrographic-like textures within the stromatolite laminae of the sample found in float (Figure 3.1C, 3.6), C) albite-

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quartz-calcite micrographic-like textures from interstromatolite zone of the sample found in float (Figure 3.1C, 3.6); D-G) albite-quartz-calcite micrographic- like textures within the stromatolite laminae of the stromatolite from stromatolite horizon A)...... 128 Figure 3.11 Elemental maps of the three stromatolite samples collected from the Mount Dunfee section of the Deep Spring Formation. A-B) sample found in float (Figure 3.1C, 3.6); C-D) sample collected from stromatolite horizon A; sample collected from stromatolite horizon D...... 129 Figure 3.12 Cathodoluminescence images from the in-float stromatolite sample from the Mount Dunfee section of the Deep Spring Formation. A) cathodoluminescence image of an albite (Ab)-rimed orthoclase grain (SEM/BSE image in “D”) from the stromatolite laminae; B) SEM/BSE image of mineral assemblage within the interstromatolite zone; C) cathodoluminescence image of the section seen in “C”; D) SEM/BSE image of the grain seen in “A”...... 130 Figure 3.13 SEM/BSE images showing examples of the K-rich aluminosilicate from the in-float stromatolite sample (Figure 3.1C, 3.6) from the Mount Dunfee section of the Deep Spring Formation. A) linear concentration of the K-rich aluminosilicate within the stromatolite laminae; B) K-rich aluminosilicate within the interstromatolite zone; C) K-rich aluminosilicate crystals (red arrows) captured within albite of the stromatolite laminae...... 131 Figure 3.14 Stromatolite-authigenesis model illustrating geomicrobiological processes leading to the precipitation of syndepositional authigenic albite and K- rich aluminosilicates...... 132 Figure 4.1 Sample preparation and SEM imaging. A) aluminum billet holding iron-poor grains; B) SEM/BSE image of the grain seen in the square area in “A” showing mottled pattern of incipient varnish; C) SEM/BSE image of the square area in “D” showing the details of a chipped surface of incipient varnish; D) SEM/BSE image of the square area in “E” showing the detail of the smooth and nodular mineralization of incipient varnish; E) SEM/BSE image of the incipient varnish seen in the square area in “B” showing details of the chipped surface...... 156 Figure 4.2 Examples of Fe-rich (no Mn) incipient varnish. A) SEM/BSE image of incipient varnish microdot; B-D) Elemental maps showing Fe (B), Mn (C), and Ba (D) of the microdot in “A”; E) SEM/BSE image of overlaying microdots showing light mineralization (right microdot) and heavy mineralization (left microdot); F) SEM/BSE image of Fe-rich microdot showing large chip through its diameter; G) SEM/BSE image of overlap of several microdots; H) SEM/BSE image of high-resolution SEM/BSE image of microdot with surficial chip; I) SEM/BSE image of an Fe-rich microdot showing nodular mineralization and light mineralization...... 157 Figure 4.3 Incipient Fe-rich varnish gradational mineralization from light (A) to heavy nodular (E) mineralization. A) SEM/BSE image of light, smooth

mineralization with chipped surface; B) SEM/BSE image of medium-light mineralization with some nodular formation; C) SEM/BSE image of medium mineralization with medium nodular formation; D) SEM/BSE image of light to medium mineralization with medium nodular formation; E) SEM/BSE image of heavy mineralization with heavy nodular formation; F) SEM/BSE image of two Fe-rich microdots showing nature juxtaposition of mineralization extremes (light to heavy, non-nodular to nodular)...... 158 Figure 4.5 Chipped incipient varnish surface. A) SEM/BSE image of chipped incipient varnish showing the Fe-rich microdot (bright) and the underlying grain (medium to dark gray); B) SEM/EDS point scan qualitative analysis showing the difference in composition between the incipient varnish and the underlying grain...... 160 Figure 4.6 Examples of advanced stages of varnish development. A) SEM/BSE image of smooth, Mn-rich varnish and associated elemental maps (Si, Mn, Pb, and Fe); B) SEM/BSE image of smooth and botryoidal varnish; C) SEM/EDS point scan quantitative analysis of points in “A” and “B” showing the change in composition relative to accelerating voltage (15 kV vs 8 kV)...... 161

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CHAPTER 1: THE EDIACARAN-CAMBRIAN BIOTIC TRANSITION IN EASTERN CALIFORNIA AND SOUTHWESTERN NEVADA AND ITS GLOBAL CONTEXT

INTRODUCTION

In his 1989 book Wonderful Life, Stephen Jay Gould presented the now iconic Burgess

Shale fauna as the epitome of what evolution was able to accomplish in diversifying animal life during the Cambrian Explosion. While it would be difficult to exaggerate the role the Burgess

Shale fauna has played in our understanding of the diversification of multicellular animals, it is important to remember that the Burgess Shale postdates the beginning of the Cambrian Period— the nominal ignition of the Cambrian Explosion—by more than thirty million years. An understanding of the dynamics of the Cambrian Explosion and the events leading up to it requires us to look even deeper into the past, into the murky paleontological mists of the

Proterozoic Eon. The Eastern California/southwestern Nevada region is one of the most important regions in the worldfor research on the fossiliferous strata that straddle the

Neoproterozoic-Cambrian boundary, and recent discoveries and analyses of these fossils have shed new light on events and processes that were occurring during this seminal episode in Earth history. Here weprovide a review of the current status of research on the paleontology of the

Neoproterozoic-Cambrian boundary in eastern California and southwestern Nevada, along with the global context in which these fossils occur.

REGIONAL STRATIGRAPHY Figure 1.1 summarizes the regional Ediacaran-Cambrian stratigraphy across eastern California and southwestern Nevada. During the Nineteenth and most of the Twentieth centuries, the base of the Cambrian System was defined by the first occurrence of trilobites, once thought to be the earliest fossilized animals. However, in the early Twentieth Century, German geologists working in Namibia discovered some pre-trilobite, non-mineralized, frond-like fossils (Xiao, 2008). In the 1940s, similar fossils were found in the Ediacara Hills of South Australia

(Glaessner, 1961). Pre-trilobite fossils, including some conical ones with phosphatic or calcium carbonate shells, were later found in many other places in the world. This led to an intense debate within the community of Cambrian paleontologists about how to define the base of the Cambrian Period. Because the base of the Cambrian also defines the base of the Phanerozoic Eon—literally “the eon of visible animals”—many researchers favored redefining the base of the Cambrian downward, to include at least some of these newly discovered, pre-trilobite fossils. This debate culminated in 1992 with the international adoption of a new definition of the base of the Cambrian Period. This expanded Cambrian System does not, however, extend low enough to include most of the pre-trilobite, animal-appearing fossils. Notably the famous molds and casts of soft-bodied organisms of Namibia, South Australia, and elsewhere—so-called Ediacara-type fossils—remain true pre-Cambrian fossils. These are diagnostic fossils of the Ediacaran (ee-dee- AK-a-run) Period, which was formally established by the International Union of Geological

Sciences in 2004 (Knoll et al., 2006). Small tubular and conical fossils—some of which have biomineralized skeletons while others do not—also occur in the Ediacaran Period.

In deciding how to define the Ediacaran-Cambrian boundary, biostratigraphers wrestled with the problem of a lack of suitable body fossils. They ultimately chose a morphologically distinctive type of burrow named Treptichnus pedum as their index fossil. By international agreement, an assemblage of trace fossils including T. pedum occurs in the Cambrian but not in

2 the Ediacaran. The boundary is formally defined as the base of the T. pedum Assemblage Zone in an exposure in Newfoundland, Canada (Geyer and Landing, 2016). This is the only biostratigraphic boundary in the geologic time scale that is defined by a trace fossil. In the strata of the Death Valley region and southern Mojave Desert, the lowest occurrence of T. pedum—and therefore the Ediacaran-Cambrian boundary—is in the Lower Member of the Wood Canyon Formation; farther northwest, in the White-Inyo Mountains of Inyo County, California and in adjacent Esmeralda County, Nevada, the boundary occurs in the

Esmeralda (middle) Member of the Deep Spring Formation (Fig. 1.1)(Corsetti and Hagadorn, 2000, 2003; Hagadorn and Waggoner, 2000; Hagadorn et al., 2000; Rowland and Rodriquez, 2014; Smith et al., 2016a). The redefinition of the base of the Cambrian System/ Period resulted in an increase in the thickness of Cambrian strata and an older date for its beginning (now determined to be 539 Ma; Gradstein et al., 2012). Ever since the Cambrian Period was originally defined in the 1830s, it had been divided into three epochs: Early Middle, and Late (Berry, 1968). Lowering the boundary to a position well below the first occurrence of trilobites added a considerable thickness of strata to the Cambrian System. In the Death Valley region, for example, approximately 400 meters of formerly pre-Cambrian strata were suddenly added to the Cambrian System. Cambrian stratigraphers were discomfited by this increased thickness in the Lower Cambrian Series, and an international commission ultimately decided to abandon the traditional tripartite subdivision of the Cambrian in favor of a quadripartite scheme (Gradstein et al., 2012). This new stratigraphic framework is reflected in Figure 1.1. The Terreneuvian Series is the newly added pre-trilobite interval. Terre Neuve is the French name for Newfoundland, where the type section of this series is located. Series 2 is the stratigraphic interval formerly known as the Lower Cambrian, and the Miaolingian Series is the interval formerly known as the Middle Cambrian. These two series/ epochs will eventually be given geographic locality- based names,

3 after type sections (global stratotype sections and points—GSSPs) are chosen. The stratigraphic interval formerly known as the Upper Cambrian (not shown in Figure 1.1), is now the Furongian Series, named after its stratotype section in China. The age of the Terreneuvian/Series 2 boundary is 521 Ma, and the age of the Series 2/ Series 3 boundary is 509 Ma (Gradstein et al., 2012). Thus the approximately 700 meters of the Wood Canyon Formation, Zabriskie Quartzite, and basal Carrara in the Death Valley column of Figure 1.1, between the earliest occurrence of T. pedum and the extinction of olenellid trilobites (marking the end of Series 2) represents 32 million years. As illustrated in Figure 1.1, the Ediacaran and Cambrian strata of eastern California thicken dramatically toward the northwest. The sedimentary facies also change. In the Kelso Mountains and Death Valley regions, thick intervals of fluvial sandstone and conglomerate are dominant lithologies, while carbonate intervals are relatively thin. In stark contrast, roughly one hundred kilometers to the northwest, fluvial facies are absent and carbonate intervals are thick (The proximity of these contrasting facies may be exaggerated somewhat due to Cenozoic tectonism.) The strata in these two adjacent regions are so different that―below the Cambrian- Series-3 Bonanza King Formation―none of the Ediacaran and Cambrian stratigraphic intervals in the Death Valley region are known by the same formation names in the White-Inyo- Esmeralda region. The rocks in the Death Valley region were deposited in proximal continental- shelf and fluvial braid-plain environments; braided-stream sandstones and conglomerates alternate with peritidal shallow-marine deposits. Northwestward into the White-Inyo-Esmeralda region, these facies transition into distal, shallow-shelf, storm-dominated, exclusively marine settings with shelf-margin reefs (Christie-Blick et al., 1989; Mount and Rowland, 1981; Rowland et al., 2008).

A WORMWORLD FAUNA IN THE GARDEN OF EDIACARA Within the past several years, to better understand the confluence of evolutionary events and environmental factors that enabled metazoan diversification to explode in the Cambrian, researchers have been aggressively focusing on the Ediacaran fossil record. This effort includes detailed studies of the famous, Ediacara-type fossils first reported from Namibia and South Australia (e.g., Narbonne, 2005; Vickers-Rich and Komarower, 2007; Xiao, 2007; Fedonkin et al., 2008; Tarhan et al., 2017). But it also includes the study of a less well-known, terminal

Ediacaran, miniature vermiform fauna—nicknamed the “Wormworld” biota by Schiffbauer et al., (2016a)— which includes the earliest biomineralizing taxa (see also Schiffbauer, 2016). Both types of Ediacaran fossils have been reported from eastern California and southwestern Nevada (Hagadorn and Waggoner, 2000; Hagadorn et al., 2000; Smith et al., 2016b), but it is the “Wormworld” fauna that has attracted the most recent attention. “Wormworld” taxa have actually been known to occur in this region for half a century; Taylor (1966) described some poorly preserved, calcareous, tubular fossils from the uppermost Reed

Dolomite. He named them Wyattia reedensis, in honor of U.C. Berkeley paleontology professor J. Wyatt Durham. In the early 1980s a fauna of tubular fossils was described from the Dunfee (lower) Member of the Deep Spring Formation at Mt. Dunfee, in Esmeralda County, Nevada (Mount et al., 1983; Signor et al., 1983). Several new skeletal taxa were described in this lower Deep Spring fauna, however Grant (1990) later synonymized them, together with Wyattia, concluding that they all were closely related to, or con-generic with, Cloudina (although Wyattia is an earlier named taxon and therefore has priority over Cloudina). Cloudina, as originally described by Germs (1972), consists of a series of nested, conical, calcite tubes, flared at their openings, often forming a curved or sinuous shell morphology. In recent years, morphologically similar, non-mineralized, or weakly mineralized, fossils named Saarina hagadorni and

Costatubus bibendi have been placed with Cloudina into the informal form-tax “cloudinomorphs” (Selly et al., 2020). Cloudinomorphs are restricted to the Ediacaran Period. Figure 1.2 illustrates the diversity of non-biomineralized cloudinomorphs from the Esmeralda (middle) Member of the Deep Spring Formation at Mt. Dunfee. These fossils were first reported by Oliver (1990) but remained unstudied until recently. They were rediscovered, along with an additional fossiliferous horizon in the Dunfee (lower) Member of the Deep Spring Formation, and reported by Smith et al., (2016a). These two horizons contain disparate assemblages of the “Wormworld” fauna. The stratigraphically lower of the two horizons contains abundant casts and molds of a tubular fossil that Smith et al. (2016) identified as Gaojiashania, a genus previously known only from the Gaojiashan Lagerstätte of South China. They tentatively assigned other specimens to Wutubus, another tubular genus previously known only from South China, but was later reassigned to Saarina hagadorni (Selly et al., 2020). Along with the tubular organisms, the newly studied Deep Spring horizons contain several ichnotaxa of trace fossils. Smith et al., (2016a) reported the occurrence of relatively abundant, poorly preserved Cloudina specimens throughout the Dunfee (lower) Member of the Deep Spring Formation, and as far down in section as the Reed Dolomite, the stratigraphic level where Taylor (1966) had recovered his Wyattia specimens. Biostratigraphically, the “Wormworld” fauna of the Deep Spring Formation displays an enigmatic pattern of biomineralization. Biomineralized Cloudina appears lower in the section than the non-biomineralized cloudinomorphs Saarina Hagadorni and Costatubus bibendi. This is contrary to the anticipated pattern in which non-biomineralizing forms would logically be expected to evolve first; biomineralization would be expected to come later, culminating with the dawn of the Cambrian and the explosively evolving menagerie of biomineralizing taxa. In the Denying Formation of South China, in fact, biomineralization follows the anticipated script: the soft-tissue cloudinomorph Conotubus, together with a diverse assemblage of other annulated

6 vermiforms, disappears from the record well before the first appearance of biomineralized Cloudina (Cai et al., 2010). This enigmatic disparity between the relative positions of biomineralized and non-biomineralized cloudinomorphs in South China and southwestern North America may be resolved with future discoveries. However, the occurrence of soft-tissue, non- biomineralized vermiforms in the Deep Spring Formation stratigraphically above Cloudina indicates that these non-mineralizing members of the “Wormworld” fauna survived well after the development of biomineralization in Cloudina and any selective pressures which may have caused it, such as an increased level of predation (Hua et al., 2003). The emerging picture suggests that the evolution of biomineralization in the Ediacaran Period is not a simple story of the gradual, progressive acquisition of this ability history by vermiform animals. Whatever abiotic and ecological factors contributed to the evolution of biomineralized skeletons, their influence evidently fluctuated through time. Diverse assemblages of metazoan trace and body fossils, as well as many other putative metazoans and classic Ediacaran’s, have been found in the Wood Canyon Formation. Hagadorn and Waggoner (2000) reported fourteen taxa in a variety of lithologies in the Wood Canyon and adjacent strata. These assemblages contain some taxa that are members of the “Wormworld” fauna. Cloudina, for example, occurs in the middle carbonate units of the Sterling Quartzite (Hagadorn et al., 2000). However, definitive taxonomic assignments of many of these fossils are impossible; the style of preservation often does not preserve the detail needed to identify important morphological features.

BIOTIC CHANGE RECORDED IN THE TRACE FOSSIL RECORD A major challenge in studying the Ediacaran-Cambrian transition globally is the dearth of stratigraphic sections containing multiple sedimentary facies. Trace fossils are best preserved in siliciclastic-dominated facies (sandstone-siltstone-shale), while carbonate-dominated facies record secular isotopic changes (especially those of carbon isotopes) in seawater that are not recorded in siliciclastic rocks. In Ediacaran and Cambrian research— more than in any other interval of geologic time—the combination of trace fossils in siliciclastic rocks and chemostratigraphic signals in carbonate rocks is critical for inter-regional and inter-continental correlation, and also for teasing the details of biotic change out of the stratigraphic record. A major reason the eastern California and southwestern Nevada stratigraphic sections have played such an important role in the history of research on the Ediacaran-Cambrian transition is that these are mixed-facies sections, with both siliciclastic and carbonate intervals. This allows the record of carbon isotopic excursions to be integrated with the trace fossil record (Corsetti and Hagadorn, 2000, 2003; Smith et al., 2016a). Here we briefly review the dynamic pattern of biotic change recorded in the trace fossil record across the Ediacaran-Cambrian boundary in this region. The Wood Canyon Formation contains abundant sandstone, siltstone, and shale, which provide ideal substrates for infaunal (living within the sediment) habitats and the preservation of the organism behaviors as trace fossils. Therefore, in sections such as those in the Death Valley region (Fig. 1.1), the story of the Cambrian explosion is told through the diversification and increasing complexity of trace fossils (Figs. 1.3-1.7). The trace fossil record thus serves as a proxy for bilaterian benthic activity (Jensen et al., 2005).

Generally speaking, the trace fossils become morphologically more complex, and they record increasing behavioral complexity, upward through the section. The quality of preservation and ambiguity of many of the trace fossils often makes taxonomic assignments difficult. For this reason, we will focus more on the general trends across the boundary and less on specific

8 ichnotaxa. As illustrated in Figures 1.3-1.7, the trace fossil record across the Ediacaran-Cambrian transition in the Death Valley regional strata changes episodically; it can be divided into three ichnofaunal assemblages: lower, middle, and upper. In the White-Inyo-Esmeralda region the different sedimentary facies of the Deep Spring Formation, compared to the Wood Canyon Formation, presents a different manifestation of the incipient Cambrian explosion than that of the Wood Canyon Formation. Trace fossils are relatively rare in the Deep Spring Formation, while in the Mt. Dunfee section, at least, tubular fossils seem quite common and widespread. So the dominant themes of the Ediacaran- Cambrian transition there are one of the advent of biomineralization, increased diversification of the epifaunal (living on the sediment surface) community, and the expansion of the Ediacaran food web (Smith et al., 2016a, Schiffbauer et al., 2016). The Mount Dunfee section provides the best opportunity for assemblage comparisons of “Wormworld” faunas throughout the world. Past studies in the Wood Canyon Formation, along with recent discoveries yet to be explored in detail (Smith et al., 2016b), suggest that a similar suite of tubular taxa exist there as well. When discussing the preservation of the behavior of early animals as recorded in their trace fossils, it is important to consider evolving sedimentary systems during the latest Neoproterozoic. The earliest burrowing animals lived in a pre- substrate-revolution world, meaning that they lived in and on non-bioturbated sediments with extensive microbial mats. During this time, finely laminated sediments could be deposited with little or no sediment mixing between layers, allowing for the preservation of shallow burrowing behaviors. The generalized view of trace fossil evolution during the early stages of the Cambrian explosion is one of increasing size, diversity, and depth of burrow penetration (Fig. 1.7). Trace fossils in the Lower Assemblage follow this generalized view perfectly, with small, shallow burrowers being the first to appear below the Ediacaran-Cambrian boundary (Figs. 1.3, 1.6, 1.7). Deeper burrowing by larger animals follows in progressively younger strata (Figs. 1.4, 1.5, 1.6, 1.7). Biostratigraphically, simple traversing burrows with an occasional loop appear in the Lower

Member of the Wood Canyon (Fig. 1.3B, E); these burrows have traditionally been referred to the ichnogenera Planolites, Helminthoidichnites, and Palaeophycus (Fig. 1.7). Within the lower member, Hagadorn and Waggoner (2000) and Waggoner and Hagadorn (2002) identified several other fossils including multiple tubular taxa preserved as casts and molds—Archaeichnium, Cloudina, Corumbella, and Onuphionella—and three members of the older Ediacara-type biota—Ernietta, Swartpuntia, and Tirasiana (Fig. 1.7). Collectively, these organisms comprise the “Lower Assemblage” of Ediacaran fossils and trace fossils in the Wood Canyon Formation, occurring within the biostratigraphic interval that extends from the first appearance of Ediacaran organisms up to the first occurrence of Treptichnus pedum. The global distribution of Ediacaran body fossils, such as those represented in the lower panel of Figure 1.8, has led to a wide variety of interpretations concerning their origin and extinction (Narbonne, 2005; Vickers-Rich and Komarower, 2007; Xiao, 2007; Fedonkin et al., 2008; Laflamme et al., 2013), specifically in connection to the later radiation of animal life in the Cambrian. A sparse assemblage of Ediacaran organisms has been found throughout the Death

Valley Region. Most of these come from the Wood Canyon Formation and include Swartpuntia, Eldonia, and Tirisiania (Hagadorn et al., 2000; Hagadorn and Waggoner, 2000; Waggoner and Hagadorn, 2002). Unconfirmed reports of the holdfast-fossil Aspidella (represented as a Charnia- like organism in the lower panel of Fig. 1.7) may extend the geographic range of Ediacara-type organisms into the Deep Spring Formation. The Ediacara-type taxa seem to have gone extinct globally prior to the beginning of the Cambrian (Xiao, 2008; Laflamme et al., 2013), and this pattern holds true in the Death Valley region. The Lower Assemblage contains the highest mixture of classic Ediacara-type organisms and members of the “Wormworld” fauna. The “Middle Assemblage” (Figs. 1.6, 1.7) consists of organisms found between the first occurrence of Treptichnus pedum and the first occurrence of trilobite body fossils. This assemblage represents the beginning of minor bioturbation and an increased complexity of trace fossils.

Because the trilobite trace fossil Rusophycus first appears well below the first trilobite body fossils, future biostratigraphic studies should be able to constrain the beginning of deep burrowing behaviors at the end of the Middle Assemblage instead of at the appearance of trilobites. The most prominent ichnotaxa within the Middle Assemblage include the large trilobite resting trace Rusophycus (Fig. 1.4B), the enigmatic, bilobed trace fossil Helminthoidichnites (Fig. 1.4A, D, E), and abundant Treptichnus pedum (Fig. 1.4E). However, a high diversity of complex burrow systems can be found with a variety of shapes and sizes (Figs.

1.4, 1.7). Importantly, the majority of trace makers in this assemblage created traces which penetrated only shallowly into the sediment. The exception is the occasional vertical burrow Skolithos. Trace fossils of the “Upper Assemblage” were made by larger animals with the ability to burrow deeply into the sediment (Fig. 1.7). Massive amounts of bioturbation begin with the introduction of this assemblage, with burrows that pass though many layers of sediment (Fig. 1.5D, E, H). While many of these large burrows are difficult to differentiate, many are similar to the simple, vertical Skolithos burrows in the Middle Assemblage, however in the Upper

Assemblage they occur in dense clusters (Figure 1.5B).

CONCLUSIONS The beautifully exposed, mixed-carbonate-and siliciclastic, Ediacaran-and-Cambrian strata of eastern California and southwestern Nevada preserve the story of the biotic transition from the Proterozoic to the Phanerozoic eons as well as it is preserved anywhere in the world, from the monotonous tranquility of the early Ediacaran seafloor to the riot of autotrophs, heterotrophs, and mixotrophs interacting in complex marine ecosystems in the Cambrian. The story of the Cambrian Explosion is recorded in great detail in these rocks, resplendent as they are with diverse assemblages of trace and body fossils, and with the recorded signals of global environmental events. As researchers become more skilled at deciphering these signals―at reading the story written in code in the strata―our understanding of this episode in Earth history and the history of life will continue to sharpen.

Figure 1.1 Generalized stratigraphic columns of Ediacaran and Cambrian strata in eastern California. Adapted from Hagadorn et al. (2000) and Corsetti and Hagadorn (2003).

Figure 1.2 Cloudinomorph fossils from the upper fossil horizon at the Mount Dunfee section of the Deep Spring Formation. (A, B, C) Tubular vermiform fossils (Cloudinids and Cloudinid-like forms) from the Esmeralda (middle) Member of the Deep Spring Formation, Mount Dunfee, Nevada. (D) hypothesized reconstruction of recalcitrant tube wall and unpreserved, fan-shaped feeding structures of Conotubus

Figure 1.3 Trace fossils typical of the Lower Assemblage in the Death Valley region, consisting of simple, shallow burrows. (A) Treptichnus pedum from the Wood Canyon Formation. (C-D) Treptichnus pedum from the Deep Spring Formation. (B, E) Simple, shallow traces from the Wood Canyon Formation in the Montgomery Mountains, Nye County, Nevada

Figure 1.4 Trace fossils typical of the Middle Assemblage in the Death Valley region, consisting of complex, mostly shallow burrows. (A) Large undifferentiated burrows. (B) Rusophycus, a trilobite burrow. (C) Small undifferentiated burrows from slab in “5E.” (D) Bilobed Helminthoidichnites burrow from slab in “5E.” (E) slab showing diversity of burrows typical of the Middle Assemblage with arrows indicating where magnified images were taken for images “5D” and “5E”.

Figure1.5 Traces and body fossils typical of the Upper Assemblage of the Death Valley region comprised of large, complex, and deep burrows. (A) Harlaniella. (B) dense cluster of Skolithos-like vertical burrows. (C) Magnified image of trilobite cephalon in “F.”(D-E,H) large, complex, vertical burrows. (F) slab containing abundant trilobite hash. (G) large slab containing highly diverse traces Rusophycus, Treptichnus pedum, and Skolithos.

Figure 1.6 Biostratigraphic ranges of various generalized groups within the Death Valley region. Representative trace fossils in each assemblagae are illustrated in Figs. 1.3, 1.4, and 1.5. Schematic reconstructions of each assemblage are shown in Fig. 1.7. Note the trend toward deeper and more complex trace fossils through time

Figure 1.7 Reconstructions of three fossil assemblages in the Death Valley region. In the middle and upper assemblages, emphasis has been placed on trace fossil diversity. See Fig. 1.6 for graphical representation of assemblages

CHAPTER 2: DISENTANGLING PYRITIZATION FROM DIRECT IRON OXYHYDROXIDE AND IRON- RICH CLAY TEMPLATING IN EXCEPTIONAL FOSSIL PRESERVATION OF EDIACARAN CLOUDINOMOPRHS

ABSTRACT A diverse assemblage of Ediacaran tubicolous vermiform fossils exists throughout the late Ediacaran–early Cambrian Deep Spring Formation, Esmeralda County, Nevada.

Taphonomic analysis of the Deep Spring fossils shows a prevalence of iron oxides, iron oxyhydroxides, and iron-rich aluminosilicate clays. The standard interpretation of these persevering minerals assumes original pyritization followed by oxidation, however, comparative analysis suggested multiple taphonomic pathways existed. To test this, electron probe microanalysis and Raman spectroscopy were conducted on two Deep Spring cloudinomorph specimens with divergent preservational characteristics and compared to the original pyrite and diagenetic iron oxides of a partially oxidized, pyritized hyolithid from the Cambrian Spence Shale in Utah. The hyolithid show a gradient in chemistry from original pyrite to the oxidized iron oxide products. Non-metric multidimensional scaling of microprobe compositional data show that the products of pyrite oxidation in the hyolithid are identical to the chemistry of one Deep Spring specimen but reveal obvious differences from the other. Raman spectroscopy confirms the presence of hematite in the former and goethite in the latter. These data, along with the lack of morphologic features diagnostic of original pyritization in the goethite-preserved specimen and their presence in the hematite preserved specimen, suggest that a taphonomic pathway in addition to pyritization (with later oxidation) was at play here. We expand and unify taphonomic modes broadly fitting under the Burgess Shale-type (BST) preservational umbrella and propose a new taphonomic pathway for soft-tissues, here termed ‘ferrumation’. It involves the direct

20 precipitation of iron oxyhydroxides and iron-rich clays on soft-tissues through abiotic and microbial iron-redox cycling within the microaerobic zone of the sedimentary column. Deeper investigations are needed to fully disentangle these disparate taphonomic pathways, however, the assumption of original pyritization for all fossils containing iron-rich minerals may be erroneous.

INTRODUCTION The emergence timing, form, and function of the earliest metazoan-grade eukaryotes have been long-standing questions in paleobiological research (Butterfield, 2015). As opposed to the vast majority of the fossil record, which is expressed through conservation of abundant skeletal hard parts, observing the earliest evolution of animals in the fossil record represents a particularly difficult challenge (Muscente et al., 2017). Much of the knowledge on this important stage in evolutionary history relies on modes of exceptional soft-tissue preservation (e.g.

Muscente et al., 2017; and reference therein). For soft tissues to enter the fossil record, a unique set of conditions is required in order to replace or replicate labile organic materials by minerals or organic molecules capable of surviving geologic time. These circumstances are ultimately rare and exceptional in the history of life (e.g., Allison & Briggs, 1993; Schiffbauer & Laflamme, 2012; Muscente et al., 2017). For studies of the Ediacaran-Cambrian transition, such exceptional fossil biotas provide valuable glimpses into the deep history of metazoan evolution. The Deep Spring Formation is exposed across the White-Inyo Mountains of Eastern

California and in Esmeralda County, Nevada. It is considered more proximal stratigraphic succession to the well-known Death Valley Ediacaran-Cambrian successions as the western Laurentia margin transitioned from an active rift during the terminal Neoproterozoic (Stewart, 1970; Armin & Mayer 1983) to a passive pargin during the early-middle Cambrian (Levy and Christie-Blick, 1989, 1991). Recently, a diverse soft-bodied, tubicolous, vermiform fauna within the Ediacaran–Cambrian Deep Spring Formation of Esmeralda County, Nevada, has been published and placed into broader chemo- and broader lithostratigraphic context (Fig. 2.1)

(Smith et al., 2016). Some fossils discovered in one stratigraphic interval of the Esmeralda Member of the Deep Spring Formation were recently placed into the recently termed ‘cloudinomorph’ morphoclade (Selly et al., 2019). Within the Deep Spring Formation, these animals appear just below the Ediacaran–Cambrian boundary, as marked by the first appearance

22 of Treptichnus pedum as well as the occurrence of the basal Cambrian negative carbon isotope (δ13C) excursion (Smith et al., 2016). Regional correlation with the Wood Canyon Formation in the Death Valley region (CA and NV), indicates that a similar assemblage of tubiform organisms, including Corumbella, Gaojiashania, Saarina, Costatubus, and the terminal Ediacaran index fossil Cloudina occur stratigraphically near the Ediacaran-Cambrian boundary (Corsetti & Hagadorn, 2000; Hagadorn & Waggoner, 2000; Smith et al., 2017). Abundant and diverse tubular fossils are also well documented in the terminal Ediacaran Gaojiashan Member of the

Dengying Formation, Shaanxi Province, China, representing a remarkably taxonomically and taphonomically similar and contemporaneous assemblage of soft-bodied tubiculous fossils, including Conotubus and Gaojiashania (Cai, Schiffbauer, Hua, & Xiao, 2012).

A Unified Taphonomic Model Constructive taphonomic forces often require microbial degradation to concentrate mineral constituents around buried organisms (Sagemann, Bale, Briggs, & Parks, 1999). To achieve this, a delicate balance must be struck between decay-related destructive forces and constructive authigenic mineral formation. Geomicrobiological-taphonomic experimentation and modeling aims to highlight the role of microbial-mediated constructive forces in creating the delicate features typified by the diverse organisms of the Burgess Shale (Allison, 1988; Kenrick

& Edwards, 1988; Grimmes et al., 2001; Briggs, 2003; Rickard et al., 2007; Mutel, Waugh, Feldmann, & Parsons-Hubbard, 2008; Schiffbauer et al., 2014; Muscente et al., 2017). The common association of authigenic minerals such as pyrite, aluminosilicate clays, phosphates with kerogenized organic matter in Burgess Shale-type (BST) fossils has resulted in much research and speculation about the nature of these enigmatic fossils (Muscente et al., 2017 and references within). Pyritization—a taphonomic mode which proceeds via microbial sulfate reduction during organic degradation within anoxic conditions—has received particular focus for its role in preserving delicate soft-tissues as well as its common association with BST kerogenized fossils (Cai et al., 2012; Schiffbauer et al., 2014). Kerogenization proceeds through the maturation of organic matter to form geologically stable organic molecules (Briggs et al., 1999; Petrovich, 1991; Stankiewicz et al., 1998), utilizing both detrital and authigenic clays as a form of organic tissue stabilizer. As such, the timing and development of authigenic clays have received considerable attention due to the uncertainty surrounding their preservational role. Some research has suggested a later stage diagenetic (metamorphic) origin for the association of clays with BST fossils (Butterfield, 1990; Butterfield, Balthasar, & Wilson, 2007; Gains, Briggs, & Yuanlong, 2008; Gaines et al., 2012; Page, Gabbot, Wilby, & Zalasiewicz, 2008), while others have suggested clays play an earlier and more constructive role in preservation (Orr, Briggs, & Kearns, 1998; Orr, Kearns, & Briggs, 2009; Petrovich, 2001; Zhu, Babcock, & Steiner, 2005), as

24 typified by the Soom Shale Lagerstätte (Gabbott, 1998; Gabbot, Norry, Aldridge, & Theron, 2009). Together, the taphonomic modes within the BST preservational window may be intimately related, requiring an inclusive model to account for differing microbial zones in the sediment column, resulting from sediment redox and microbial zonation (Canfield & Thamdrup, 2009) within which the decaying organism likely interacts (Schiffbauer et al., 2014). Per this redox zonation model, BST preservation (kerogenization) and Beechers Trilobite-type preservation (pyritization) occur as a continuum of preservational modes for soft-tissue fossil in the Gaojiashan Lagerstätte during early diagenesis (Schiffbauer et al., 2014). Taphonomic modes in the redox zonation model (Schiffbauer et al., 2014) result from the interaction of the buried organism with the surrounding microbial/redox (bioredox) environment. As sedimentation proceeds, the taphonomic, redox, and geomicrobiological environment of the sediments hosting the decaying organism shift over time, allowing for evolution of the taphonomic environment and resulting mineralization. The rate of sediment deposition was therefore a key control for the duration in which organic tissues reside in certain bioredox zones. Inclusion of several other taphonomic modes into the sediment zonation model—silicification, phosphatization, siderite mineralization, and calcification (Muscente et al., 2017)—has broadened our understanding of the geomicrobiological and geochemical conditions necessary for exceptional preservation to occur.

The Powerful Role of Ferruginous Conditions Microbial Fe (II) oxidation has been linked to massive deposition of Fe (III) oxyhydroxides, such as in banded iron formations (e.g. Posth et al., 2013; Chan et al., 2016). The resulting biogenic Fe (III) minerals play a potentially important role in the preservation of microbial cells and organic remnants (Posth, Canfield, & Kappler, 2014; Chan et al., 2016; Picard et al., 2015). However, the microaerobic zone of the sedimentary column has not been explored for its role in taphonomic processes. This was likely due to taphonomic overprinting by later stage sulfate and ferric iron reduction and the resulting pyritization within the sulfate reduction zone of the sedimentary column. The Gunflint biota is the notable exception, as early speculation led to the conclusion that it represents Fe (II)-oxidizing bacteria which, through their own metabolic activity, encased themselves in ferric iron precipitates (Barghoorn & Tyler, 1965; Planavsky et al., 2009). The taphonomic power of ferruginous microaerophilic environments is likely very high. Organic degradation is slow due to low oxygen concentrations and absence of metabolic diversity beyond the chemolithoautotrophy of iron-oxidizing bacteria (FeOB), resulting in conditions in which mineralization is rampant and organic degradation is slow. Coupled with the cycling of iron through the microaerobic zone of sediment due to reduction occurring below and oxidation occurring above—as well as the demonstrated complementary microbial consortia of Fe (II)-oxidizers and Fe (III)-reducers within the same sedimentary horizon (Gault et al., 2011)—such conditions seem highly conducive to Fe (III)-mineral templating of organic tissues. We present data here that indicate that these hidden taphonomic conditions can be misinterpreted as the mineralogical remnants of pyrite oxidation. We show that diagnostic signs of pyrite oxidation (compositional and morphological) are lacking in some Deep Spring cloudinomorph specimens, and we propose a novel taphonomic pathway to explain these differences. The purpose of this paper is to expand the sediment zonation model of Schiffbauer et

26 al., (2014) and Muscente et al., (2017) to explore the role FeOB played within the microaerobic zone to create the taphonomic conditions within the Deep Spring Formation.

METHODS Specimen collection & selection More than 100 specimens were collected from the upper fossil horizon within the Esmeralda Member of the Deep Spring Formation at Mount Dunfee, Esmeralda County, Nevada. Deep Spring specimens come from a siliciclastic interval between the lower two microbialite horizons within the section (Fig. 2.1) (Smith et al., 2016). Two Deep Spring cloudinomorph specimens—DSM-200 (LVNHM BLM 2020.001) and DSM-300 (LVNHM BLM 2020.002)— representative of taphonomic varieties which displayed markedly different preservational characteristics such as mineral morphology, mineral chemistry, and overall fossil preservational quality, were selected for further analysis using electron microprobe and Raman spectroscopy. Several other specimens collected from the same fossiliferous horizon were polished and imaged, one of which (LVNHM BLM 2020.003) was included here to highlight the gradational features of the taphonomic processes at work in the Deep Spring Formation. The taxonomic classification of these specimens has been established before (Selly et al., 2019). To better understand the mineralogical products of pyrite oxidation within similar aged fossils, a hyolithid specimen (Hyolithes cerops) (UNLV-LSS-001) collected from float in the Spence Shale Member of the Langston Formation (early middle Cambrian) at Minors Hollow in the Wellsville Mountains, Box Elder County, Utah. The apparent oxidation state of the pyritized hyolithid fossil provides the opportunity for direct geochemical comparison of original preserving pyrite and its oxidized products. Fossils of the Deep Spring Formation, as well as those of the Spence Shale Member, are hosted in comparable shaley lithologies (Albers & Stewart, 1972; Garson et al.,

2011; Liddell et al., 1997). Specimens were photographed using a Nikon D3100 while submerged in alcohol to accentuate color and contrast. For petrologic analysis, a pair of unoriented thin sections from fossil-bearing strata of the two different taphonomic varieties were prepared and polished. Deep Spring specimens DSM-200 and DSM-300, as well as the Spence

Shale hyolithid specimen, were analyzed using the same analytical methodology discussed below, including scanning electron microscopy, electron microanalysis, Raman spectroscopy, and non-metric multidimensional scaling.

Electron probe microanalysis Data were collected through wavelength dispersive X-ray spectroscopy (WDS) using an electron probe microanalyzer (EPMA; JEOL Superprobe). Compared to the semi-quantitative geochemical method energy dispersive X-ray spectroscopy (EDS), WDS provides fully quantitative data on elemental composition of minerals with comparison to standards, thus providing more accurate measurements of mineral composition and more reliable statistical analysis. Point analyses and elemental maps were collected of the two selected Deep Spring specimens and the one Spence Shale Hyolithid. Point analyses provide the most accurate quantitative chemical data of the complicated phases, while elemental maps allow for interpretation of the distribution of elements (and interpreted phases) within the specimens. Point analysis measured the “major” element package, including Mg, Ti, Ca, Na, Si, Fe, K, Al, P, Mn, Cl, F, Cr, Ba, and S. The probe condition was using 15kV, 10 nA, and 1 μm beam spot.

Scanning Electron Microscopy (SEM) Samples previously subjected to EPMA analysis were polished with the 0.25 μm diamond paste finishing and coated with carbon for electrical conductivity before imaging with the scanning electron microscope (SEM). SEM imaging was done to assess the morphological characteristics of preserving minerals using JEOL JSM-5600 for lower magnification images and JEOL JSM-6700F for high-resolution images. Backscatter electron (BSE) imaging was used to image relative differences of the average atomic weight of the minerals, and to allow real time distinguishing of preserving minerals from background sediment. BSE imaging effectively provides a contrast differentiation based largely on atomic number (which is related to backscatter coefficient, η), thus providing a compositional image, rather than the secondary electron imaging based on surface topography.

Raman spectroscopy To determine mineralogy, Raman spectroscopy was performed using a Horiba T64000 triple-grating spectrometer, equipped with an Olympus MPlan N 50x microscope objective with numerical aperture 0.75. A Spectra-Physics Stabilite 2018-RM operating at 514.5 nm was used for excitation with ≤15 mW power in a 1 µm spot. Inelastic scattering (Raman scattering) of incident light results in the presence of optical modes from the interaction of light with phonons/molecular bonds within the material. Known frequency changes of Raman modes can be used to identify and assess chemical species in both organic and mineral phases. For this study, Raman spectroscopy was used to identify the iron phases within the two selected Deep Spring specimens when compared to data collected by Hanesch (2009) and Das & Hendry (2011). The use of Raman spectra in previous studies of BST fossils (Marshall et al., 2012) has expanded the available data on mineral composition in the taphonomic environment.

Data analysis Non-metric multidimensional scaling (NMDS), a multivariate ordination technique was used to organize groupings of chemical characteristics relative to each specimen from microprobe data. To produce ranked comparisons of chemical data sets, a Bray-Curtis dissimilarity coefficient was used to highlight the differences between points within the dataset. To assess the amount of phase mixing and the nature of pyrite oxidation, four elements are selected for NMDS analysis—Fe, S, Si, Al. Microprobe data points used for NMDS analysis included only those of iron oxides. Two types of data were collected using the microprobe: point scans and elemental maps. Phase maps using EPMA point analysis data are constructed of the two selected Deep Spring specimens using elemental maps in conjunction with point scans. Bins are made using point scans of iron oxyhydroxide and iron-rich clays and correlated to elemental map raw point counts to show the distribution of both phases within the fossil.

RESULTS Taphofacies Description & Fossil Identification Unoriented cross sections of two distinct, fossil-bearing horizons (taphofacies) within the roughly 30 cm thick upper fossil horizon were thin-sectioned and imaged for sedimentological information. Taphofacies 2 (Fig. 2.2g-j) was made of organic-rich clays, indicating consistent deposition of clay and organic matter throughout the polished section. Within taphofacies 2, the majority of preserved tubicolous specimens are found in a wedge-shaped layer of dense organic debris evident by the typical red/orange of iron oxyhydroxides and iron-rich clays (Fig. 2.2g). This thin layer of preservation was 4.5 mm thick at its thickest point. The geometry of this layer, pinching out laterally, suggests that deposition/preservation of dense organic matter was heavily spatially dispersed into discrete pockets roughly 10 cm-long and up to 4.5 mm-thick. Taphofacies 2 also contained abundant isolated spots of concentrated precipitation of iron oxyhydroxides dispersed throughout the sediment (bright spots in Fig. 2.2j). The consistency of sediment through taphofacies 2 are in stark contrast to that of taphofacies 1, which displayed abundant depositional changes (Fig. 2.2a). The roughly 1 cm thick section of taphofacies 1 contained siltstone as well as very-fine to fine grained sands and a ~ 3 mm thick fossil-bearing, organic-poor siltstone occurs adjacent to a thin layer of tapering, fine-grained sandstone and alternating fine and very-fine sandstone. Taphonomic quality was highly disparate between the two taphofacies. Taphofacies 2 contained abundant, poorly preserved cloudinomorph fossils (Fig. 2.2h), while taphofacies 1 contained rare, three dimensionally preserved specimens of high fidelity (Fig. 2.2b). Cross sections of cloudinomorph specimens from taphofacies 2 shows linear concentrations of iron oxyhydroxide with poorly defined edges surrounded by iron-rich clays (Fig. 2.2i). The occasional, well preserved specimen from taphofacies 2 may contain detailed morphological information, such as of the specimens of focus for this study—DSM-200 (Fig. 2.3)—as well as

DSM-210 (Fig. 2.5). Fossils found in taphofacies 1, however, represented the best-preserved specimens in the collection, such as the other Deep Spring specimen of focus for this study— DSM-300 (Fig. 2.4) and other specimens collected from the same fossil horizon (Fig. 2.2b). These specimens provided the most detailed, three-dimensional morphological information on cloudinomorph fossils of the Deep Spring Formation. Cross sections of cloudinomorph specimens from taphofacies 1 shows well-defined iron oxide tube walls (Fig. 2.2c-f, 4), which may or may not be pervasively filled (Fig. 2.2f shows this gradation), or be associated with an iron-rich clay halo (Fig. 2.2c-e). These iron-rich clays are not apparent in horizontally polished specimens under SEM analysis due to their low iron content and lack of drastic density differences in SEM/BSE imaging. However, they appeared clearly in polished section due to the bright red coloration from reflected light (Fig. 2.2c-e).

Mineral Morphology Iron oxyhydroxides display a wide variety of morphologies throughout specimen DSM- 200 (Fig. 2.3) and DSM-210 (Fig. 2.5), ranging from individual micronodules (Fig. 2.5b,c zone 3, 2.6a-b) to large, massive, amalgamations (Fig. 2.5b-c zone 1, 2.6a). Micronodules can be solitary or occurring in apparent clusters with varying morphologies and connectedness—many in the form of hollow microspherules (Fig. 2.5b,c zones 2&3) and sometimes with the beginnings of an inner sphere (Fig. 2.5c section 2)—while others form internally consistent spheres or misshapen masses (Fig. 2.5c section 1, 2.6a-b). Sub-µm-scale (~100 nm wide) nanofibers radiated outward from some individual micronodules, particularly those that form clustered hollow spheres. Within specimen DSM-200, a morphological change is seen in clays associated with high amounts of iron (identified using BSE imaging), with higher-iron clays forming larger plates (Fig. 2.6c). The majority of iron oxides in specimen DSM-300 (Fig. 2.4) shows more homogenous and massive (Fig. 2.4a, Fig. 2.6d). The few areas which displayed complex morphology (Fig.

2.6e-f) are dissimilar to those of specimen DSM-200 in that DSM-300 lacked the ever-present micronodules and microspherules seen in specimen DSM-200 (Fig. 2.3) and DSM-210 (Fig. 2.5). DSM-300 did, however, show similar mineral amalgamations, but those had cubic morphologies (Fig. 2.6e-f). The Spence Shale hyolithid shows typical signs of pyritization, including rhombs (Fig.

2.7c) and framboids (Fig. 2.7d). Most importantly, pyrite within the hyolithid was currently undergoing oxidation and transforming into iron oxides. SEM/BSE imaging shows the presence of pyrite cores surrounded by diagenetically oxidized rims of iron oxides (Fig. 2.7b) which, was pseudomorphing the original pyrite morphologies (Fig. 2.7c-d).

Geochemical Characterization of Specimens The detailed geochemical and mineralogical characterization of specimens DSM-200 and DSM-300 was undertaken using electron microprobe and Raman spectroscopy. Three different mineralogies with incorporated iron are defined within the Deep Spring specimens and their surrounding sedimentary environment—iron-rich aluminosilicate clays, goethite, and hematite. Aluminosilicate clays with little to no iron made up the majority of the sediment encompassing the fossils. Raman spectroscopy indicated that goethite was present only within specimen DSM-

200 as shown by characteristic peaks at~250, 300, 400, 460, 555, and 685 cm-1 (Fig. 2.8b), while hematite was present only within DSM-300, as indicated by peaks at ~225, 300, 410, 505, 670, 800, and 1325 cm-1 (Fig. 2.8b). Microprobe data shows highly varying elemental compositions throughout DSM-200 which indicated that goethite was mixed with clays on the scale of a few microns (the size of the electron beam) (Fig. 2.8a). Microprobe data from DSM- 300 was remarkably homogeneous (Fig. 2.8a), however, variation was also likely the result of phase mixture with clays but in smaller amounts.

To determine whether mineral phases shows a specific distribution within the fossils, Fe- phase maps were constructed for the two Deep Spring cloudinomorph specimens (Fig. 2.3c, e, and Fig. 2.4f). Views at different scales shows distributions within entire specimens (Fig. 2.3c and 2.4f) and distributions focused on one transverse cut of the tube wall for DSM-200 (Fig. 2.3e). When looking at DSM-200 overall, the distribution of high-iron clays near the tube walls was evident. The majority of DSM-200 was preserved through iron-rich clays, with only small amounts of iron oxyhydroxide dispersed throughout the clays and larger concentrations of iron oxyhydroxide occurring near the edges. In contrast, DSM-300 was preserved through the nearly ubiquitous presence of iron oxyhydroxides throughout the specimen (Fig. 2.4f) with the uncommon inclusion of an unidentified Fe-mineral with higher microprobe Fe total counts (“Distinct Fe-mineral” of Fig. 2.4f). The lack of any iron-rich clays within DSM-300 was the

37 most conspicuous difference between the two Deep Spring tubular specimens. Specimen DSM- 200 (Fig. 2.3) shows a relatively higher amount of iron-rich clays than DSM-300 but relatively lower mineralization overall. Clays containing iron within specimen DSM-300 are disseminated throughout the matrix, had relatively little iron, and seemingly had no connection to the fossil itself while iron mineralization of the fossil was more pervasive (Fig. 2.4f). In contrast, Fe maps of DSM-200 revealed high concentrations of iron-rich clays close to the tube walls (Fig. 2.3). Microprobe data across the Fe-rich clay-preserved tube wall of DSM-200 (red line in Figure 3b, c, d, e, g) shows a centrally increasing iron content (in weight percent oxides from WDS data) with iron counts highest at the center of the tube wall (17%) and decreasing outward with a perfect gradient toward 0% iron within the surrounding sediment (Fig. 2.3f). Fossils from the Spence Shale (Fig. 2.7) shows similar taphonomic characteristics to fossils from the Deep Spring Formation and from previously reported specimens from the Gaojiashan Formation (Cai et al., 2012). Massive pyrite nodules could be found lightly scattered throughout the Spence Shale, along with large nodules of seemingly oxidized pyrite shown as red staining of sediment (see the iron halo around the hyolithid specimen in figure 2.7a and compare to the iron halo around the specimen in figure 2.2c-d). However, while original pyritization was evident in the form of cubic pseudomorphs, framboids, and remnant pyrite cores, most Spence fossils have been partly or completely diagenetically oxidized. The Spence Shale hyolithid (Fig. 2.7) chosen for analysis was only partially oxidized, allowing for direct analytical comparison of original taphonomic pyrite and diagenetic iron-mineral weathering pseudomorphs. Microprobe analysis of the chemistry of the Spence Shale hyolithid shows the continued presence of sulfur in iron-mineral products of up to 13% and as little as 0.4% (Fig. 2.8a). Levels of sulfur and iron shows moderate negative regression within the iron oxide products of pyrite preserving the hyolithid (r2= -0.78), a similarly moderate negative regression was found in DSM-300 (r2= -0.78), and a low regression was found in DSM-200 (r2= 0.07).

NMDS analysis of microprobe data of Si, Fe, Al, and S, of specimens DSM-200 (Fig. 2.3), DSM-300 (Fig. 2.4), and the Cambrian hyolithid (Fig. 2.7) revealed several important details (Fig. 2.8a). A trend in the hyolithid dataset can be seen between the original hyolithid pyrite of point H-1 to the fully oxidized iron oxides of point H-2 (Fig. 2.8a). Data from DSM- 300 appear highly clustered at the oxidized end of this trend within the hyolithid (Fig. 8a). Data from DSM-200 appear to be highly scattered and separate from the hyolithid pyrite oxidation trend (Fig. 2.8a).

DISCUSSION Taphonomic Pathways at Play in the Deep Spring Formation Iron-rich minerals in the Deep Spring Formation cloudinomorphs contained a variety of morphologies (Fig. 2.6) and associated chemistries (Fig. 2.8). Geochemical analyses show iron- rich minerals with little or no sulfur within the Deep Spring Formation, leaving only iron oxides, iron oxyhydroxides, and iron-rich clays as preserving minerals. Similar chemistry has been reported in weathered specimens from the Gaojiashan Lagerstätte, which were preserved through pyritization, some of which shows later oxidation (Cai et al., 2012; Schiffbauer et al., 2014). Specimens from the Gaojiashan, however, shows clear evidence of original pyritization in the form of iron oxide framboidal and cubic pseudomorphs. Features diagnostic of original pyritization, best characterized by Grimes et al., (2002) as “microcrystalline, framboidal, massive polycrystalline, and subhedral or euhedral forms”, are present in some Deep Spring specimens in the form of cubic and octahedron pseudomorphs (Fig. 2.6e,f, see 2.7c,e for hyolithid comparison) within specimen DSM-300 (Figure 2.4). However, those features are completely absent in other specimens—specimen DSM-200 (Fig. 2.3, Fig. 2.6a-c) and DSM-210 (Fig. 2.5). Furthermore, the lack of iron oxides in large sections of the tube wall of DSM-200 (Fig. 2.3d-e) which are preserved with iron-rich clays, represented a mineralogy completely removed from pyrite and its oxidation. This suggests mixed or diverging taphonomic pathways resulting in different precipitate morphologies and chemistries.

Even though the resulting diagenesis resulted in similar chemistries (iron-rich, sulfur- poor minerals), mineralogical differences between specimens DSM-200 (goethite) and DSM-300

(hematite) indicated that these two specimens likely are the end products of different mineralization pathways, before diagenetic history had nearly homogenized the resulting chemistry. When the four main elemental constituents are analyzed using NMDS (Fig. 2.8), a pattern emerged showing oxidation of the pyritized hyolithid, starting with relatively high sulfur

(point “H-1”: 10.1% SO3, 48.9% FeO) and ending in a grouping of fully oxidized points with higher relative iron (point “H-2”: 0.4% SO3, 71.6% FeO, and point “3”: 0.1% SO3, 79.3% FeO). Adjacent to these oxidized pyrite points are the majority of points from the Deep Spring specimen DSM-300 (Fig. 2.8). This trend and correlation between specimens suggest that they had been subjected to similar taphonomic conditions—that of original pyritization and oxidative alteration. Specimen DSM-200, however, was situated outside of this oxidation pattern with its own trend, perpendicular to that of the hyolithid data points (Fig. 2.8). This suggests that it was not subjected to the same chemical pathway to mineralization. This was also suggested by grain morphology, with specimens DSM-300 and the hyolithid displaying mineral morphologies dissimilar to those seen in DSM-200. Microprobe data revealed another important similarity between the Spence hyolithid and DSM-300: both specimens show negative correlation between

2 2 abundance of FeO and SO3 (r = -0.78 and r = -0.72, respectively) while the same index for DSM-200 resulted in no correlation (r2= 0.07). This suggests that within the hyolithid specimen and DSM-300, the amount of remnant sulfur in diagenetic hematite, which had a pyrite precursor within the hyolithid (Fig. 2.7), was very specific, based on its original composition and chemical pathway to oxidation (Luther et al., 1982). In contrast, DSM-200, which shows no morphological signs of original pyritization, had no correlation between sulfur and iron which likely assimilated sulfur during the formation of authigenic iron oxyhydroxides. We therefore propose two possible hypotheses to account for this in the preservation of the Deep Spring cloudinomorph organisms: 1) pyritization was the original preserving authigenic mineralization pathway, proceeding as previously described with the aid of sulfate-reducing microbes (Schiffbauer et al., 2014), with later oxidation and recrystallization to its current form; or 2) authigenic mineralization involving the direct precipitation of non-sulfidic, iron-bearing minerals, such as iron oxyhydroxides or iron-rich clays. The former hypothesis, however, would be based on negative evidence as several of the analyzed Deep Spring specimens show no

41 evidence of original pyritization (Fig. 2.3, 2.5a-c) and they provide different microprobe compositional data and Raman analysis (Fig. 2.7) from those that show signs of original pyritization (Fig. 2.4, 2.5d-f).

Taphonomic Model Description and Justification Expanding from the microbial zonation model of Schiffbauer et al., (2014), we present a taphonomic model which emphasizes the metabolic actions of FeOB communities during early stages of sedimentation and organic matter decomposition (Fig. 2.9). Here termed ferrumation, this taphonomic pathway utilizes the abiotic and microbially-mediated cycling of iron within the microaerophilic zone of sediment to precipitate iron oxyhydroxides and iron-rich clay templates of cloudinomorph organisms. Focusing on pathways of tissue stabilization and mineralization, this model relies on changes in mineral phase precipitation due to changes in redox conditions and dominant microbial communities as sedimentation proceeds after initial burial of the organism. Comparable to the Schiffbauer et al., (2014) model, the rate of sedimentation (first) and the sediment microbial profile (second) are the key determining factors for the timing and duration in which taphonomic windows are open by constraining how long decaying organic matter remains in certain microbial zones. We suggest that, while pyritization and kerogenization predominantly occur in the sulfate reduction and methanogenesis zones, respectively

(Schiffbauer et al., 2014), the direct association of decaying organisms with minerals generated from microbial iron oxidation promoted the authigenic precipitation of iron oxyhydroxides and iron-rich clays within the microaerophilic zone of the sediment column. Fossil preservation through direct precipitation of iron oxyhydroxides is not a new concept (Douglas & Beveridge, 1998; Ferris, Fyfe, & Beveridge, 1988; Konhauser, Fyfe, Ferris,

& Beveridge, 1993; Orange et al., 2014; Urrutia & Beveridge, 1994), nor is biologically induced/mediated clay formation considered rare in the rock record (Bontognali et al., 2014;

Gehling, 1999; Konhauser & Uruttia, 1999; Sánchez-Navaz, Martin-Algarra, & Nieto, 1998; Tuck et al., 2006; Wacey et al., 2014). Fossils of the Gunflint biota provide an excellent example of such preservation at work, however on a much smaller scale—that of individual Fe-oxidizing bacteria (FeOB) (Barghorn &Tyler, 1965). With the notable exception of the recent Shapiro &

Konhauser (2015) work, most research on the Gunflint fossils has suggested that the biota consisted of FeOB, which, through their own metabolic activities, encased themselves in iron- rich minerals (ferrihydrite) during the oxidation of ferrous iron, with later stabilization by a microcrystalline quartz matrix. Expanding from this limited preservation by direct iron oxyhydroxide precipitation on individual cells to a taphonomic model in which entire, macroscopic animals are encased in a tomb of iron minerals during the metabolic activities of a microbial consortia is the next stage of developing this taphonomic scenario.

Redox Zonation and Microbial Influences on Ion Concentration

At least two distinct fossiliferous taphofacies occur within the upper fossiliferous layer at Mount Dunfee, Nevada, each of which contains specimens with specific preservational characteristics (Fig. 2.2). These differences provide important insights into the interconnectedness of mineralization pathways. Differences in ion concentration and mineralization between the two taphofacies are likely due to the amount of background organic matter within the sediment, a relationship discussed in Osés et al., (2017). The particular mineral phase was determined by the position within the sediment column where mineralization takes place. A significant difference in the amount of iron oxyhydroxide disseminated throughout the sediment surrounding the fossils was evident by the relatively bright spots from BSE detection [Fig. 2.2f (taphofacies 1) and 2.2j (taphofacies 2)]. The highly concentrated iron oxide mineralization in taphofacies 1 (Fig. 2.2f) at the site of tubicolous decay suggests that iron ions are highly localized, originating within the confines of the tube walls. This concentrated mineralization resulted in highly localized precipitation, which yielded high-quality fossils typically preserved in three dimensions (Fig. 2.2b-f). Alternatively, the higher number of mineralized patches within the surrounding sediment of taphofacies 2 (Fig. 2.2j) suggests a more diffuse iron ion concentration front, in which mineralization occurred more slowly and with a

44 lower degree of concentration. This resulted in more abundant fossils with lower preservational quality (Fig. 2.2g-j). Taphofacies 2 (an organic rich shale) represents a more diffuse mineralization, while taphofacies 1 (an organic poor shale) represents highly concentrated sites of mineralization. Per the microbial zonation model, the oxic-anoxic boundary results in drastically changed microbial metabolisms and resulting metabolites. Straddling this boundary was the microaerophilic redox zone, within which oxygen and iron gradients interact in the sedimentary column to produce a wide variety of iron oxyhydroxides and iron oxides. This zone contains oxygen concentrations between .2%-12% (Ferrara-Guerrero & Bianchi, 1990), ideal conditions for iron oxidation by microaerophilic FeOB. Within this relatively thin zone, microaerophilic FeOB retain a significant advantage over aerobic and anaerobic bacteria, as oxygen concentrations are too low for vigorous aerobic metabolisms and too high for anaerobic metabolisms. This relatively thin redox zone may have very high taphonomic potential due to these characteristics but has previously been overlooked in taphonomic studies (Schiffbauer et al., 2014), likely due to the similarity of the mineral products with those of pyrite oxidative alteration. Herein, I emphasize the potential importance of the iron cycling within the microaerophilic zone of the sedimentary column in taphonomic modeling due to (1) the reduced decomposition efficiency of microbes, (2) iron absorption onto structural biopolymers (Petrovich, 2001), and (3) the nucleation of authigenic minerals which template soft-tissues.

Applying our amended microbial zonation model (Fig. 2.9) to Deep Spring fossils, specimen DSM-200 (Fig. 2.3) experienced relatively rapid emplacement into the microaerophilic zone (taphofacies 2, Fig. 2.2g-j), below the influence of aggressive aerobic decay. Specimen DSM-300 (Fig. 2.4) appears to have been pyritized, which implies rapid emplacement of the decaying organism within the anoxic sulfate reduction zone (taphofacies 1, Fig. 2.2a-f). Kerogenized fossils, comparable to a minority of those in the Gaojiashan biota, are exceedingly

45 rare in the Deep Spring Lagerstätte (Rowland & Rodriguez, 2014). Kerogenization requires rapid burial with minimal decay through the sedimentary column, followed by emplacement for a relatively longer period in the methanogenesis zone (Fig. 2.9). These three preservational pathways represent a continuum of mineralization modes moving downward through microbial zones in the sediment column. They are not necessarily isolated or independent and may co- occur in a single specimen (Schiffbauer et al., 2014), implying that mineralization takes long enough for organic matter to transition between taphonomic zones in the sedimentary column.

The type and abundance of animal tissues which interact with the biogeochemical environment outlined by the microbial zonation model will determine mineralization pathways. Outwardly comparable to the cloudinomorph organisms of the Gaojiashan biota (e.g., Cai et al., 2011; Schiffbauer et al., 2016; Smith et al., 2016), the Deep Spring fossils appear to be constructed of relatively recalcitrant tube walls. These may or may not be lightly mineralized and contain within them relatively more labile inner tissues, which have only been found partially preserved (Schiffbauer et al., 2020). The refractory nature of the tube walls allowed them to resist decay, while the labile inner tissues provided an abundant, localized organic carbon source that promoted metabolic degradation by whichever microbial community dominated the specific bioredox environment in which the tube resided. As the labile inner tissues are degraded, metabolites and ions become concentrated in the surrounding and internal chemical microenvironment of the decaying organism due to the diffusion- and decay-resistant tube walls

(Schiffbauer et al., 2014), resulting in the rapid supersaturation and precipitation of preserving minerals.

Fe-Mineralization Pathways

A major question persists in models of fossil preservation through the replication of labile tissues by the precipitation of authigenic minerals: how do nucleation sites originate? Earlier

46 work on Gaojiashan pyritization (Schiffbauer et al., 2014), BST kerogenization (Petrovich, 2001), and Soom Shale aluminosilicification (Gabbot, 1998) suggests that the tube wall tissues of tubular organisms are the substrate for mineral precipitation by providing ample charged surfaces for nucleation sites. Within the Gaojiashan Lagerstätte, pyrite precipitation began on the inner tube wall and proceeded inward toward the center of the tube (Schiffbauer et al., 2014). A similar pattern can be seen in several Deep Spring cloudinomorphs which show progressive mineralization with nearly complete mineralization at the tube wall (Fig. 2.3a-c, and zone 1 of

Fig. 2.5b-c) and little to no mineralization near the center of the tube (Fig. 2.3a-c, and zone 3 of Fig. 2.5b-c). These biogenic nucleation sites originate from the deprotonation of some tissues at neutral pH giving them negatively charged surfaces mostly contained within carboxyl and phosphate functional groups—structural biopolymers (Hoyle & Beveridge, 1983, Hoyle & Beveridge, 1984). The iron adsorption model proposed by Petrovich (2001) for kerogenization should also be considered. In the Petrovich (2001) model, nucleation sites can be generated from small amounts of Fe2+ and Fe3+ which are adsorbed onto these negatively charged sites on structural biopolymers of an organism’s tissues. Adsorption of these ions may be controlled by a variety of factors, including pH (Warren & Ferris, 1998). Metal cation adsorption has been well documented to occur on microbial surfaces and on negatively charged extracellular polysaccharides (Beveridge, 1993; Eushima & Tazaki, 2001; Fein, Daughney, Yee, and Davis, 1997; Konhauser & Uruttia, 1999; Shultze-Lam, Harauz, & Beveridge, 1992; Schultze-Lam,

Fortis, Davis, & Beveridge, 1996; Tashiro & Tazaki, 1999; Tazaki, 2005; Theng & Orchard, 1995; Tuck et al., 2006).

The abundance of chemical pathways in which microbial communities produce iron minerals makes precise determination of the origin of preservational iron oxyhydroxides difficult (reviews in Emerson et al., 2010; Posth, Canfield, & Kappler, 2014). Additionally, many abiotic processes can result in the same mineral products. Relevant pathways include, (but are not

47 limited to): (1) the abiotic precipitation of iron hydroxides with later binding to negatively charged biopolymers of bacterial cells (both dead and alive), (2) the active precipitation of ferric iron phases due to the redox changes around a metabolically active cell (Konhauser, 1998), and (3) the release of Fe3+ ions into the surrounding environment where they then precipitate out as amorphous iron hydroxides (Park et al., 2014). Potentially, all of these pathways could contribute to the mineralization patterns seen in Deep Spring cloudinomorphs. Within the taphonomic environment of the Deep Spring Formation, precursor ferrihydrite precipitation likely occurred around the sheath of the metabolizing bacterium, on negatively charged tissues of partially decayed organic matter as organic matter-mineral macromolecules (Lalonde, Mucci, Ouellete, & Gelinas, 2012), and also within the surrounding sediment (Park et al., 2014)(Fig. 2.2j). Since ferrihydrite is relatively unstable, it quickly alters to FeOOH (goethite), FeOOH*H2O (limonite), through hydration and dehydration reactions (Konhauser, 1998) depending on the redox history of the strata. If sedimentation proceeds at a fast-enough pace and places organic matter into the sulfate reduction zone, then decay proceeds through microbial sulfate reduction and preservation occurs through tissue replication by pyrite (as described in detail by Schiffbauer et al., 2014). To gain energy through the oxidation of organic matter, sulfate-reducing microorganisms (SRM) use the reduction of sulfate (SO42-) as an electron acceptor producing abundant HS- ions in the surrounding chemical microenvironment of the decaying organism. These HS- ions can then combine with ambient porewater-sourced Fe2+, tissue-adsorbed Fe2+, or Fe-reducing bacteria- derived Fe2+ to produce pyrite. Since bioavailable iron in the form of sedimentary iron oxyhydroxide is a more thermodynamically available ion to reduce than that of sulfate—and it is reducible by at least some SRM (Coleman, Hedrick, Lovely, White, & Pye, 1993; Lovely & Phillips, 1987; Lovely, Roden, Phillips, & Woodward, 1993), iron oxyhydroxides must not have consumed all organic matter to allow for sulfate reduction to take place (Konhauser, 1998).

However, limited amounts of microbially-mediated iron oxyhydroxide production before the tissues have reached the anoxic zone of sediment may aid in pyritization and kerogenization by providing localized bioavailable iron for iron reducers (Love, 1967; Kenrick & Edwards, 1988), which plays an important role in iron adsorption onto structural biopolymers (Petrovich, 2001). In this sense, pyritization and kerogenization may be aided by earlier ferrumation of the fossil- containing strata by locally concentrating highly reactive iron oxyhydroxides and hydroxides to combine with biogenic bisulfide during pyritization (Li, Vali, Yang, Phelps, & Zhang, 2006;

Peiffer et al., 2015) or to further protect tissues from methanogenic decomposition during kerogenization. Fe (III) reduction by SRM and FeOB can occur through at least two pathways, either 1) direct reduction of Fe (III) by SRM or 2) the reduction of Fe (III) while reacting with biogenically produced sulfides (for experimental evidence of this see Peiffer et al., (2015) for abiotic conditions and Li et al., (2006) for biotic conditions). The presumed destruction of early ferrumation-derived iron oxyhydroxides during the pervasive pyritization of tubular fossils of the Gaojiashan Lagerstätte, may have eradicated all signs of early ferrumation. Conversely, the limited pyritization seen in Deep Spring specimens at Mount Dunfee may have allowed the mineralogical signs of ferrumation to slip through the ever-present destruction of fine-scale features in the geologic realm and expose a hidden taphonomic pathway for the preservation of soft-tissues. The microbial influence on soft-tissue preservation is well known (Briggs, 2003) and well-integrated into taphonomic models (Schiffbauer et al, 2014). However, the effect of microbial metabolites on mineral dissolution has been largely unexplored in taphonomic research. Enhanced dissolution of some silicate minerals by microbial metabolites is also a well- documented process and may provide the necessary Al and Si ions for the authigenic precipitation of clays (Douglas, 2005; Ehrlich, 1996; Fenter et al., 2003, Fenter, Zapol, He, and Sturchio, 2014; Lee & Fein, 2000;Lee, Brown, Hodson, MacKenzi, & Smith, 2008; Liermann,

Barnes, Kalinowski, Zhou, & Brantley, 2000; Li, Liu, Chen, & Teng, 2016; Liu et al., 2006; Ren et al., 2017; Sánchez-Navaz et al., 1998; Teng, Fenter, Cheng, & Sturchio, 2001; Ullman, Kirchman, Welch, Vandevivere, 1996; Vandevivere, Welch, Ullman, & Kirchman, 1994; Visscher & Beveridge, 2005; Vorhies & Gaines 2009; Welch & Ulmann, 1996; Welch, Barker, & Banfield, 1999). This has been documented previously when microbial cells and extracellular polysaccharides become incased within authigenically precipitated aluminosilicates, silicates, and Fe/Mn Oxides (Darroch et al., 2012; Douglas & Beveridge, 1998; Konhauser & Urrutia,

1999; Lee, Brown, Hodson, MacKenzie, & Smith, 2008; Orange et al., 2014; Pheonix & Konhauser, 2008; Sánchez-Navas et al., 1998; Tazaki, 1997, 2005; Tuck et al., 2006; reviewed by Newman et al., 2017). Incorporation of iron into clays can occur if ions are available, which seems to be evident within Deep Spring fossils by the centrally increasing iron concentration seen in the cross section of the iron-rich clay-preserved tube wall of specimen DSM-200 (Figs. 2.3c,e,f, also see next section for further discussion). The ubiquitous mineral spheres (hollow and filled) found throughout many of the Deep

Spring cloudinomorphs may represent coated microbial cells present during mineralization of the fossils (Fig. 2.5, 2.6b). The size of the hollow spheres was at the extreme end of microbial size (~10 microns diameter), however, similarly sized hollow spheres are preserved within an endolithic microbial biofilm through Mg-silica precipitation in the sand tufa of Mono Lake, California (Souza-Eqipsy, Wierzchos, Ascaso, & Nealson, 2005). Souza-Eqipsy et al., (2005) interpreted these hollow spheres as forming during the precipitation of Mg-silicates on microbial cell walls and extracellular polymeric substrate. A similar interpretation of smaller dolomitic hollow spheres found within Miocene microbialites (Sanz-Montero et al., 2009a, 2009b ) and selenite gypsum beds (Ayllón-Quevedo et al., 2007) has been proposed, in which hollow spheres are interpreted as having formed during the biomineralization of the cell walls. Within the Deep Spring specimens, a second internal sphere can be seen forming (Fig. 2.5c zone 2) which may

50 represent continued mineralization of the cell wall after post-death cellular dehydration and collapse. Other than general shape and size consistency, no direct evidence for microbial coating of the cell was evident. However, no other mechanisms are known which produce hollow spheres with such remarkably consistent shape and size.

Early Stage Authigenic Clays

Several studies have expressed the importance of early stage, authigenic clays in templating soft-tissue fossils as primary preserving minerals (Gabbott, 1998, 2001; Gabbot, Aldridge, & Theron, 2007; Orr et al., 1998; Orr et al., 2009; Petrovich, 2001) and as secondary stabilizing minerals (Zhu et al., 2005), while others have suggested they formed during diagenetic alteration (Butterfield et al., 1990; Butterfield et al., 2007; Gains et al., 2008, 2012; Page et al., 2008). The complete preservation of sections of the tube wall of DSM-200 through iron-rich clay precipitation suggested herein, while reminiscent of kerogenized fossils from other BST localities, represents a drastic divergence from the pyritization-kerogenization gradients found preserving cloudinomorph fossils from the Gaojiashan Lagerstätten (Schiffbauer et al., 2014) due to the lack of association with kerogenized organic matter. Specimens from the Soom Shale, preserved by early taphonomic illite which precipitated authigenically as a response to changes in pH during degradation of buried tissues (Gabbott, 1998; Gabbot, Norry, Aldridge, & Theron, 2001), are the closest comparison to the taphonomic clays of Deep Spring specimens. As such, authigenic precipitation of iron-rich clay in Deep Spring fossils has no bearing on the debate behind the origin of clays in association with carbonaceous films of karogenized fossils.

However, this study does confirm the powerful role authigenic clays play in preserving exceptional fossils. Significant portions of specimen DSM-200 are preserved primarily through iron-rich clay replication of the tube wall (Fig. 2.3). There are four potential origins for the iron-rich clays in

51 this taphonomic environment: 1) Iron was incorporated into the detrital clay surrounding the decaying cloudinomorph due to the concentration of iron ions within the zone of preservation during iron oxidation/reduction, creating an iron-rich clay outline of tube walls, 2) nano-scale particles of iron oxyhydroxide which are too small for Field Emission Electron Microscopy detection (~10 nm detection resolution) are coating the clay matrix, making it appear during microprobe data collection that the iron has been incorporated into the clay minerals, while in reality there are two separate phases contributing to the data, 3) authigenic precipitation of iron- rich clays, likely in the form of ferroan saponite (Petrovich, 2001), templates the outermost layers of the tube walls as a result of changes in pH during tissue degradation [similar to the authigenic illite templating in the Soom Shale (Gabbott, 1998; Gabbot et al., 2001; Whittle et al., 2007)] or 4) a thin layer of Fe-mineral, deposited during the earliest stages of mineralization, precipitated when ample adsorbed Fe nucleation sites and free Fe ions are available, is situated below iron-poor clay which precipitated later, when both are less abundant. All four of these scenarios theoretically can exist, depending on the preservational conditions relating to available ions and their concentrations. Based on imaging and microprobe data, we conclude that early, authigenic iron-rich clays formed and templated decaying tube walls (scenario 3). This interpretation is bolstered by the fact that SEM imaging of clays in this area shows that individual iron-rich grains display drastically different morphologies than nearby iron-poor grains (Fig. 2.6c). Also, no known mechanism for iron incorporation into clay structure is known

(scenario 1), SEM imaging shows no signs that nano-scale precipitates are coating clay grains (scenario 2), and no evidence was found that iron oxyhydroxides are situated beneath iron-rich clays (scenario 4).

Mineral Diagenesis The oxidation of pyrite has been shown to produce a variety of mineralogical byproducts (Luther et al., 1982; Nordstrom, 1982) including jarosite (KFe2(SO4)2(OH)6, ferrihydrite (Fe (OH)3), goethite (FeOOH), and hematite (Fe2O3). Mineralogical transformations occur through dehydration/hydration and oxidation/reduction reactions, often with the aid of microbial metabolites. The late stage diagenetic alteration of minerals is the result of large-scale conditions during burial and uplift, in contrast to small-scale biogeochemical conditions that are present during sedimentation. Such large-scale, late-stage diagenetic changes would be expected to homogenize the mineral products through a series of mineral transformations. Here, we suggest that two mineralization pathways are involved in the preservation of the Deep Spring cloudinomorphs because hematite and goethite (both mineralogical products of pyrite oxidation) exist within the fossils; if both are pyrite oxidation products, it would be expected that oxidation would have homogenized the products to either goethite or hematite, but not both. Therefore, two early diagenetic mineralization pathways, each with their own diagenetic alteration sequence, are needed to explain the presence of both goethite and hematite within the Deep Spring fossil horizon: Ferrihydrite→goethite (DSM-200 & 210) Pyrite precursors→pyrite→ferrihydrite→goethite→hematite (DSM-300) The geochemical conditions which result in the mineralogical transformation of pyrite-to hematite (pathway 2) without the concurrent dehydration of the presumed early diagenetic goethite of DSM-200 (Fig. 2.3) is unknown and enigmatic. However, it is likely due to the relative timing of goethite formation being early in DSM-200 preservation (~545 ma) and later stage, post-uplift diagenesis for DSM-300. The association of early goethite with organic matter during the DSM-200 preservation may have imparted mineralogical characteristics which allowed for more stable, dehydration-resistant goethite. Alternatively, conditions during late-

53 stage pyrite oxidation may never have reached a true goethite stage and went straight to hematite—as seems to be evident for the hyolithid, where no goethite stage was evident in NMDS analysis (Fig. 2.8a). Furthermore, the clay-rich taphonomic environment may have allowed for significant microbially-mediated silica dissolution during tissue mineralization, as evident by authigenic iron-rich clay associated with the goethite-replicated tissues (Fig. 2.3). Elevated preservation of microbially-associated goethite was shown to occur in the presence of Si during experimental diagenesis studies (Picard, Obst, Schmid, Zeitvogel, & Kappler, 2016). The coexistence of authigenic iron-rich clays within DSM-200 and authigenic goethite may have allowed for the reduced alteration of original goethite within this taphonomic environment beyond its typical geological stability. These unique conditions have revealed the early mineralogical transformations of iron-rich minerals associated with decaying organic matter in the sediments of late Ediacaran seas.

CONCLUSIONS Comparative taphonomic analysis of two Ediacaran, non-mineralized, cloudinomorph fossils from the Deep Spring Formation along with a partially oxidized pyritized hyolithid from the Cambrian Spence Shale reveals subtle differences in chemistry and morphology which we interpret to have resulted from variation in chemical pathways to preservation. The taphonomy of all three specimens are similar—consisting of iron oxides in all three, with remnant, original pyrite in the hyolithid—however, detailed analysis suggests that Deep Spring specimen DSM-

300 and the Spence Shale hyolithid are preserved through the same pathway of pyritization, while Deep Spring specimen DSM-200 was preserved through the direct precipitation of iron oxyhydroxides and authigenic iron-rich clays. This novel preservational mode is herein termed ferrumation—soft-tissue replication through the direct precipitation of iron oxyhydroxides and iron-rich aluminosilicate clays on tissues and microbial cells within microaerophilic sedimentary zones. This taphonomic pathway begins with metal cation adsorption on structural biopolymers

(ligand bonds) during early stages of decay, reducing decomposition efficiency and forming abundant nucleation sites for preserving minerals. Diverging taphonomic pathways then proceed through a branching series of iron-mineral phase transformations after the initial precipitation of amorphous ferrihydrite on cloudinomorph tube walls as determined by the positioning within the microaerophilic zone of the sedimentary column (Fig. 2.9); ferrihydrite either 1) matures to a more stable iron oxyhydroxide such as goethite after little or no interaction with the sulfate reduction zone, or 2) transforms to pyrite after Fe (III)-reduction in the presence of bisulfide within the sulfate reduction zone, later oxidizing to hematite during late-stage exhumation. The assumption of original pyritization for all fossil-associated iron oxides is likely due to the similarity of products of pyrite oxidation to those of direct iron oxyhydroxide precipitation. While signs of this taphonomic window are likely rare, a closer examination of pyritized fossils

55 and their oxidized mineral products may reveal additional examples of this biogeochemical interaction.

Figure 2.1 Stratigraphic column of the upper Reed Dolomite and Deep Spring formations at Mount Dunfee, Esmeralda County, Nevada. Modified from Smith et al., (2016).

Figure 2.3 Microprobe elemental maps and point analysis of specimen DSM-200 (LVNHM BLM 2020.001). a) reflected light photograph showing real colors; b) Iron x-ray map; c) Iron phase map using x-ray map from (b), orange indicates goethite [identified by Raman spectroscopy (Fig. 8)] and green indicates iron-rich clay; d) Iron x-ray map; e) Iron phase map from “d” with same color classifications as (c), note minor amounts of goethite dispersed throughout clay; f) Microprobe data points across red line in b, c, d, and e.

Figure 2.4 Microprobe elemental maps of specimen DSM-300 (LVNHM BLM 2020.001). a) BSE image after polishing of specimen; b-e) elemental maps constructed using electron microprobe showing relative intensities of Fe (b), Al (c), Si (d), and S (e); f) Fe phase map showing distribution of hematite (purple) and an unidentified Fe (II)-phase. Note conspicuous absence of iron-rich clays associated with the fossil.

Figure 2.5 Progressive mineralization seen in cloudinomorph specimen DSM-210 (LVNHM BLM 2020.001). a) polished specimen submerged in alcohol and imaged under fluorescent light showing parts of the specimen that are polished and under epoxy; b) section of DSM-210 showing three zones of mineralization (1 being most pervasively mineralized and 3 being least); c) detailed observations of mineral templating of microbial cells within the tube walls of DSM-210, ranging from partial coating (zone 3) to pervasively-infilled (zone 1. Note the second layer forming within the enclosed microspherules (arrow).

Figure 2.6 Light and BSE images showing variation of Fe-minerals in tube fossil specimens DSM-200 (a-c) and DSM-300 (d-f). The surrounding matrix in all images was made up of aluminosilicates with highly variable amounts of iron; a) goethite micronodules conglomerating into massive buildup goethite at the tube wall; b) goethite micronodules and pervasively-infilled microspherules within the iron rich aluminosilicate clay matrix (see more of this morphology in Fig. 4); c) morphological difference between iron-rich clay (large, platy) and iron-poor clay (fibrous); d) massive hematite at the tube wall; e,f) hematite pseudomorphs of pyrite rhombs (arrows) within iron-poor clay matrix. Note rhomb pseudomorph amalgamation in (f).

Figure 2.7 Hyolithid (UNLV-LSS-001) from the Cambrian Spence Shale Member, Langston Fm, of Northern Utah. a) Reflected light photograph of Hyolithes cerops (no color alteration). Sample has been polished near center of the hyolithid and the remainder was under epoxy; b-d) BSE images of oxidized pyrite crystals. Brighter areas represent pyrite, medium areas represent iron oxides, and darker areas represent clays.

Figure 2.8 Non-Metric Multidimensional Scaling (NMDS) plots of Microprobe data and Raman spectra of Deep Spring Formation specimens DSM-200 (red dots/lines), DSM-300 (green dots/lines), and a Spence Shale hyolithid (blue dots) with interpretation. a) NMDS plot made using FeO, SiO2, Al2O3, and SO3 weight percentages, with a table of four points to illustrate differences in chemistry; b) Raman spectra of DSM- 200 (red) and DSM-300 (green) showing important peaks used for mineral identification.

Figure 2.9 Iron biogeochemical cycle in relation to phase precipitation based on the sediment/microbial zonation model showing sources of sedimentary iron phases and zones of preservation. Addition of the microaerophilic zone into the model provides a small horizon for dynamic cycling of iron phases within the sedimentary column. Ferrumation would dominate the mineralization pathways within the microaerophilic zone.

CHAPTER 3: MICROBIALLY MEDIATED DISSOLUTION OF DETRITAL ORTHOCLASE AND ITS REPRECIPITATION AS AUTHIGENIC AND COMPOSITIONALLY DIVERSE PLAGIOCLASE—A POTENTIAL MICROBIALITE DATING TECHNIQUE?

ABSTRACT The Ediacaran to Cambrian Deep Spring Formation at Mount Dunfee in central Nevada has gained renewed importance with the recent discovery of non-mineralizing tube fossils within the Ediacaran-age Esmeralda Member. However, highly diverse stromatolitic microbialite reefs throughout the Esmeralda Member have made it an international attraction to geoscientists for decades. While previous studies have detailed the morphological diversity of the microbialites and their respective environments, I looked into the mineralogical constituents found captured within the carbonate laminae. Detailing both detrital and authigenic phases, I developed a model which explains observed feldspar and quartz morphologies as the result of the complex biological activity inherent within an actively growing microbial mat. A morphologically diverse assemblage of feldspars was found throughout the carbonate matrix, including orthoclase (k-rich endmember), albite (Na-rich endmember), a compositionally diverse Ca-plagioclase (Na/Ca mixture), as well as abundant quartz and a K-rich aluminosilicate. These mineral phases were found intimately associated with each other, often displaying micrographic-like wormy textures typically formed during the intergrowth and coprecipitation of feldspars and quartz during the cooling of magmas when the two phases are simultaneously stable and oversaturated. While micrographic textures have only been previously documented within igneous rocks, with similar textures forming during hydrothermal alteration of feldspars (albitization), I propse that the same textures can form from the coprecipitation of mineral phases from heterogeneous solutions

66 within the mineralizing portions of an actively growing stromatolite. This stromatolite- authigenesis model involves (1) the enhanced dissolution of orthoclase due to organic-metal complexation of Si and Al as well as the bioextraction of K ions for cellular osmotic regulation within the stromatolite, (2) the capture of dissolved ionic species within the complex biochemical EPS framework of the microbial mat, (3) the degradation of organic matter within the mineralizing portions of the stromatolite, and (4) the release of complexed ions into solution where they combine with seawater to produce authigenic albite, Ca-plagioclase, quartz, and a K-rich aluminosilicate. These mineral associations have been confirmed in four stromatolites from the Deep Spring Formation (three stromatolite horizons from the Mt. Dunfee section and one from the Molly Gibson Mine section), and one stromatolite horizon from the Miocene Duero Basin, Spain. Another example has been tentatively identified within a stromatolite from the Flinders Range, Australia.

INTRODUCTION General Background As the first fossil evidence of life on Earth, microbialites provide important manifestations of early life and the influence of microbes on the rock record (Reitner, Quéric, and Arp, 2011). Windows into the past are rare, however, microbialites provide an abundance of easily recognizable biogenic structures which are important reminders of a thriving microbial biota on a world which could not yet sustain complex life. Microbes arrived early and set the evolutionary stage for the development of complex, macroscopic life by effecting nearly all geochemical cycling of nutrients, including the oxygenation of Earth’s atmosphere. Although modern microbialites are restricted to environments uninhabitable to most life on Earth—such as hypersaline, restricted marine and lacustrine basins—their continuous restriction to the photic zone throughout their history allows for a quick paleoenvironmental assessment wherever found in the rock record. Their informative nature, their importance within the history of life on Earth, and their striking appearance make microbialites well sought after by Earth scientists and astrobiologists. The purpose of this project is to expand on the geobiological utility of microbialites by introducing a potential geochronological technique which utilizes the in-situ formation of organominerals within the microbially active region of the growing stromatolite. Microbialites of the middle member Deep Spring formation at Mount Dunfee, Nevada, display a remarkable diversity of form, occurring in nine distinct horizons with closely associated quartz siltstones and sandstones (Fig. 3.1, 3.2) (Oliver, 1990; Oliver & Rowland 2002). For this reason, the Deep Spring Formation has been interpreted as a locus of carbonate sedimentation and stromatolite buildup in the late Proterozoic. Microbialite morphologies include digitate bioherms, inclined bioherms and biostromes, isolated forms, and cryptomicrobial boundstones. Due to their high topographic relief and large geographic expanse, the microbialites

68 of the Deep Spring Formation have been interpreted as Precambrian microbialite reefs exerting significant physical control over the local environment (Oliver, 1990).

Microbial Dissolution of silicates and the resulting precipitates Microbially induced dissolution and precipitation of minerals has been previously reported to aid in many aspects of microbial life. Ehrlich (1996) identified a variety of reasons why microbes would interact with the lithosphere in this way, utilizing both dissolved ionic species and authigenic precipitates either by (1) being used as a source of energy or as a terminal electron acceptor, (2) using them for trace element requirements, or (3) to enhance general competitiveness in the microbial community (Ehrlich, 1996). Microbial dissolution is accomplished through the interaction of organic acids and other metabolites created by a thriving microbial community which break up Si-O and Al-O bonds within silicate minerals or form organic-mineral ligands which pull ionic species directly from the mineral framework (Ehrlich, 1996). Mineral precipitates can benefit microbial communities through (1) removal of harmful oxidative/reductive metabolites from their microchemical environment, (2) the uptake of ionic species during energy metabolism, and (3) the uptake of ionic species for conversion into tissue (Ehrlich, 1996). Authigenic mineral formation occurs both passively in the surrounding chemical microenvironment from the effects of ionic species uptake as well as actively within the microbe through the conversion to salts or oxides or on the protective walls which serve as templates for mineral precipitation (Ehrlich, 1996). Together, these microbe-mineral interactions result in a wide variety of geobiological effects on the rock record. Much attention has been given to the ability of microorganisms and fungal hyphae to affect the dissolution of silicate minerals (Barker, Welch, Chu, & Banfield, 1998; Bennet, Rogers, Choi, & Hiebert, 2001; Buss, Lüttge, & Brantley, 2007; Ehrlich, 1996; Fenter, Zapol,

He, & Sturchio, 2014; Lee & Fien, 2000; Lee, Brown, Hodson, Mackenzie, & Smith, 2008; Liermann, Barnes, Kalinowski, Zhou, & Brantley, 2000; Li, Liu, Chen, & Tang, 2016; Qureshi, Qureshi, Sodha, Tipre, & Dave, 2018; Sánchez-Navaz, Martín-Algarra, & Nieto, 1998; Liu et al., 2006; Ullman, Kirchman, Welch, & Vandevivere, 1996; Visscher & Stolz, 2005; Welch &

Vandevivere, 1994; Welch & Ullman, 1999), showing both increasing and decreasing dissolution rates (compared to nonbiological kinetic dissolution) depending on the microbial assemblage and the biochemical molecules present. However, directly studying microbial precipitates in this environment can be difficult and mostly involves authigenic aluminosilicates, amorphous silicates, and hydrated iron minerals (Bar-Or & Shilo, 1988; Darroch, Laflamme, Schiffbauer, & Briggs, 2012; Decho et al., 2005; Douglas & Beveridge 1998; Konhauser & Urrutia 1999; Konhauser, Fyfe, Ferris, & Beveridge, 1993; Lee et al., 2008; Orange et al., 2014; Phoenix &

Konhauser 2008; Phoenix, Adams, & Konhauser, 2000; Sánchez-Navas et al., 1998; Tazaki, 1997; Tazaki, 2005; Tuck et al., 2006; Good review in Newman et al., 2017). These are just a few examples of how the metabolic activity of bacteria, under specific redox and sedimentological conditions, can produce a variety of authigenic minerals. Microbially-mediated mineral dissolution/precipitation in low-temperature-pressure environments is not only common, but these processes also provide significant amounts of mineralization within sedimentary environments. The best example of these microbially induced sedimentary structures are the microbialites found throughout most of Earth’s history.

Microbialites The experimental study of biofilms and microbial mats is becoming increasingly important in geomicrobiological studies due to their expected influence on the preservation of exceptional fossils and their many potential biomarkers for astrobiological studies. Many of these studies have focused on the importance of microbial mats in preserving the delicate soft- tissues of the Ediacaran biota with the development of the “death mask” hypothesis (Darroch et al., 2012; Gehling, 1999), as well as the importance of biologically influenced mineralization in the preservation of Burgess Shale type fossils (Schiffbauer et al., 2014). The best examples of this type of work comes from the development of taphonomic models involving the precipitation of authigenic iron sulfides during the metabolic activities of sulfur-reducing bacteria (Picard, Kappler, Schmid, Quaroni, & Obst, 2015; Picard, Gartman, & Girguis, 2016; Picard, Gartman, Clarke, & Girguis, 2018; Schiffbauer et al., 2014 ), which likely aided in the creation of “death masks” as well. Iron-oxidizing microbes are well known for their ability to mediate the precipitation of organo-mineral structures (Chan, Fakra, Emerson, Fleming, & Edwards, 2011,

Picard et al., 2015, 2018), iron oxides (Bazylinski, Frankel, & Jannasch, 1988), and iron oxyhydroxides (Miot et al., 2009)—much of which is focused in microaerophilic environments. For the purposes of this study, I will follow the nomenclature of Dupraz et al., (2009) for differentiating microbialites and mineral constituents. Under this nomenclature, a microbialite is any mineral precipitate (organomineral) that forms due to the influence of biological processes.

Microbialites can be either biologically induced forms (i.e., stromatolites and thrombolites) or more passively mineralized organic matter that formed due to the mere presence of life. Within a microbialite, organomineralization is controlled by two components: 1) the alkalinity engine, which is controlled by both biological activity within the microbial mat as well as environmental conditions, and 2) through the interaction of extracellular polymeric substances (EPS) (Deco et al., 2005 ;Dupraz, Visscher, Baumgartner, & Ried, 2004; Dupraz et al., 2009). Biominerals, on

72 the other hand, are those which are strictly precipitated by life with purpose, function, and form, such as mollusc shells or the biominerals of the magnetotactic bacteria (Lefevre and Bazylinski, 2013). Microbial mats are those which do not produce significant organominerals or exert major control on local sedimentary structures. The central role EPS play in the cycling of nutrients throughout the biosphere is becoming increasing clear (Deco et al., 2004; Decho & Gutierreze 2017; Dupraz et al., 2004, 2009). EPS are primary constituents of microbialites and microbial mats, seems to be intimately associated with most marine biochemical processes, and assert considerable control over marine ecosystems, while also influencing micro-scale community structure of microbial consortia. While EPS are highly varied depending on environmental variables, they commonly contain polysaccharides, proteins, lipids, and nucleic acids (Dupraz et al. 2009). By increasing the efficiency of the alkalinity engine, microbial consortia primarily influence the precipitation of carbonates in marine stromatolites and thrombolites; however, EPS have been implicated in a wide variety of mineral precipitates including aluminosilicates, iron sulfides, iron oxides, apatite, and struvite (Decho and Gutierreze 2017). The role of EPS in stromatolite formation involves a complex series of transformations, decompositions, and re-secretions as they become consumed by the ever-growing carbonates and captured detrital grains (Dupraz et al. 2009). Long term studies of the stromatolites at Highborne Cay (Bahamas) has revealed a series of layered microbial consortia consisting of layered biochemical metabolisms which follow a redox gradient (Dupraz et al., 2009; Reid et al., 2000). Environmental conditions, along with these internal biogeochemical conditions, result in daily variability in growth dynamics of the stromatolite as well as different mineral dissolution/precipitation patterns (Reid et al., 2000). Typically, microbial consortia within stromatolites display taxonomic and metabolic layering, in which the highest energy consuming metabolisms appear at the top (such as photosynthetic metabolisms), and the lowest appear in the internal dark layers (such as methanotrophic

73 metabolisms) (Dupraz et al., 2009; Visscher & Stolz, 2005). Oxygen and redox gradients occur in older stromatolites with the most developed microbial zonation, which contain zones of aerobic and anaerobic metabolisms (Visscher & Stolz, 2005).

Microbial Dissolution Microbes have been shown to directly dissolve mineral surfaces to bioextract necessary ionic constituents of life (Iverson et al., 1978; Qureshi et al., 2018). Qureshi and colleagues (2018) detailed the ability of several fungi and bacteria strains to bioextract potassium ions from feldspars, with specific attention on organic molecules responsible as well as chemical characteristics which increase dissolution rates. They found that both fungi and bacteria strains are able to bioextract potassium from feldspars at a rate from 10-40ppm depending on major microbial metabolites present and chemical conditions of the medium. Similarly, Ivarson et al., 1978 showed the ability for the iron-oxidizing bacterium Thiobacillus ferrooxidans to bioextract Na and K ions from various sources including glauconite, illite, microcline, and albite, resulting in the authigenic formation of jarosite and natrojarosite. In the Ivarson et al. (1978) study, the resulting authigenic mineral formation was controlled by the cation source—jarosite formed in the presence of glauconite, illite and microcline while natrojarosite formed when albite was available for cation bioextraction. Additionally, while the Liu et al. (2006) study showed no enhanced dissolution of feldspars by Bacillus mucilaginosus, it did show the important role polysaccharides and organic acids play in complexing with SiO2 to drive dissolution of silicate minerals through enhanced solubilization of the mineral surfaces (Liu et al., 2006). In an ingenious experimental setup, Lee et al. (2000) showed that bacterial presence (both viable and non-viable) enhanced dissolution of gibbsite over sterilized controls. More importantly, they showed that lysis products may be as important in dissolving minerals as metabolic products (metabolites). However, lysis products from viable experiments showed an increased dissolution capacity over non-viable experiments. Similarly, bacterial dissolution of minerals seems to be taxon-specific (Lee et al., 2000). Ullman (1996) showed that only specific microbial metabolites will increase dissolution rates—organic acids acting at Al sites on the mineral surface—while others, such as alginate and

75 poly-aspartate, significantly reduce dissolution by decreasing surface reactivity and surface area (Ullman, 1996). Focusing on the mineral bytownite, Vandevivere et al. (1994) and Welch and Vandevivere (1994) examined the change in dissolution potential of several metabolites in near- neutral pH, showing that 8 metabolites increased dissolution (Vandevivere et al., 1994; Welch & Vandevivere, 1994). Gloconate-promoted dissolution was also demonstrated for albite, quartz, and kaolinite (Welch & Vandevivere, 1994). These studies show that metabolites primarily composed of sugars showed no effect on the dissolution of feldspars and other silicate minerals; uronic acids and high molecular weight polyaspartate significantly decrease dissolution rates, while low molecular weight polyaspartate significantly increase dissolution rates (Vandevivere et al., 1994; Welch & Vandevivere, 1994). The increase in dissolution rates affected by polyaspartate likely is caused by complexing with framework ions in solution, thereby keeping ions out of solution and increasing proton-promoted dissolution (Welch & Vandevivere, 1994). High-and low-molecular-weight metabolites resulted in reduced dissolution potential due to the binding of these metabolites to mineral surfaces (Welch and Vandevivere 1994, Vandevivere et al., 1994). However, Chen et al. (2018) showed an increase in mobilization and biotransformation of As/Fe-rich mine soil with the addition of low-molecular-weight metabolites into the system (Chen et al., 2018).

Microbially-induced precipitation of silicates In a study utilizing Focused Ion Beam and TEM instruments, Lee et al. (2008) investigated the formation of authigenic mineral coatings around 1.1-thousand-year-old feldspar grains collected from soil. They documented the authigenic growth of several mineral phases including an Fe-K aluminosilicate (smectite), which likely grew in an “amorphous and probably organic-rich matrix” (Lee al., 2008, p. 1319), likely referring to EPS. Most previous studies have looked only at one side of the microbe-mineral interaction coin, studying either microbially induced dissolution or microbe induced mineralization. In an impressively thorough experimental study, Newman et al. (2017) developed a model for the authigenic precipitation of aluminosilicate on cyanobacterial cells within a cultured microbial mat with known substrates for dissolution and ion utilization. In the Newman et al. (2017) study, the chemistry of authigenic precipitates was dependent on the introduced substrates (i.e., glass beads, illite powder, and quartz) as well as the amount of agitation the cultured cyanobacteria experienced. As one of the few studies that used a wholistic approach—experimenting with both dissolution of substrates and reprecipitation of authigenic minerals—the Newman et al. (2017) study provides a steppingstone toward the further understanding of biologically induced precipitation of minerals in the sedimentary environment. Importantly, due to the only source of iron in their experimental quartz-substrate setup being biological, the inclusion of iron-rich aluminosilicate precipitates in their results illustrates the ability of biologically sourced ions to become incorporated into the geological realm (Newman et al., 2017). Several studies have identified a variety of Mg-silicates within modern microbialites, including kerolite and stevensite (Burne et al., 2014; Gérard et al., 2018; Pace et al., 2016; Zeyen et al., 2015). These Mg-silicates have been shown to precipitate around microbial cells very early, even before carbonate mineralization of the mat has begun (Pace et al., 2016) and are preceded by amorphous gels.

In their study of the modern stromatolites of Lake Dziani Dzaha, Western Indian Ocean, Gérard et al. (2018) found that within the complex biogeochemical environment of a growing stromatolite, two separate microbial taxa were responsible for separate authigenic mineralization within the mat (Gérard et al., 2018); alphaproteobacteria were responsible for the precipitation of aragonite (the primary mineralogical constituent of the stromatolites) and pleurocapsales were responsible for the precipitation of the Mg-silicate. The Mg-silicate precipitate formed after the microbes had accumulated silicon and magnesium within their sheaths before desiccation in the mineralizing zones of the stromatolite. Similar results were found within the modern stromatolites of the Great Salt Lake, in which a Mg-silicate mineral was found to precipitate before aragonite precipitation of the mat began (Pace et al., 2016).

Authigenic feldspars in carbonate rocks The study of authigenic silicates within carbonate rocks goes back to the mid-nineteenth century with the discovery of authigenic albite within a Triassic dolostone in France (Drain, 1861; Kastner & Siever, 1979). However, Drain’s theory of low temperature authigenic feldspars within sedimentary rocks was not accepted until the mid-twentieth century, when a broader understanding of these minerals began to form (Kastner & Siever, 1979). More recent studies, along with more than a century of research, have collected evidence to support the low- temperature formation of authigenic feldspars (see review in Kastner & Siever, 1979). Observations of high optic axial angles, along with the Kastner (1971) study using EPMA to show the end member chemistry of authigenic feldspars (KAlSi3O8 and NaAlSi3O8), documented that feldspars can grow in temperatures below 100°C (Kastner & Siever 1979). Once a low temperature of formation was proposed by Drain (1861), paleoenvironmental interpretations moved from strictly marine (Grandjean, 1909) to saline-dependent waters (Reynolds, 1929). The relative importance of water chemistry vs detrital component availability has been a major debate surrounding the growth of authigenic feldspars in carbonate rocks. However, the chemical composition of the water is likely dependent to some degree on the presence of detrital material and its dissolution (at least micro-spatially). Buyce & Freidman (1975) suggested that the necessary K and Na ions come from the in-situ dissolution of detrital grains, while Fuchtbauer (1956) suggested the importance of seawater for the availability of Na and K ions (Buyce &

Freidman, 1975; Fuchtbauer, 1956). A useful (but very-out-of-date) summary of authigenic feldspar research can be found in Kastner and Siever (1979).

Authigenic albite is typically thought to occur in high-temperature hydrothermal brine environments during albitization (Bernard et al., 2012; Gold, 1987; Kasterner and Siever, 1979). However, the presence of albite in some sedimentary rocks suggests that the formation temperature can reach as low 70°C, such as those authigenic albites found within the Green

River Formation (Cole, 1985; Nuccio & Johnson, 1988). Kasterner and Siever (1979) outlined the characteristics of authigenic feldspars in hydrothermal environments, with later emending by Mu et al. (2016), including (1) end-member purity of authigenic feldspars, (2) lack of luminescence, (3) high isotopic values, (4) non-twined crystal behavior. Turner and Fishman (1991) concluded that the diagenetically altered ash beds of the Bushy Basin member of the Morison Formation resulted from syndepositional, low-temperature diagenetic forces of unusual pore-water chemistry and potentially elevated pore-water temperatures due to the solar heat capture exhibited by some saline ponds (Turner & Fishman, 1991). In their model, a lateral hydrogeochemical gradient existed across the basin producing a variety of authigenic mineral assemblages. These include authigenic albite and illite, both of which had previously been thought to occur only in high temperature geochemical environments and have been shown to precipitate dependently from hydrothermal solution (Mu et al., 2016; Turner & Fishman 1991). Similarly, Cole (1985) and Williamson (1987) had interpreted authigenic albite within the Green River Formation and Miocene tuffs, respectively, as syndepositional albite. Both Mu et al.

(2016) and Turner and Fishman (1991) emphasized the importance of the illite-albite associations. However, Mu et al. (2016) described their authigenic albites as having formed during hydrothermal alteration of arkosic sandstones at higher temperatures ranging from 90- 120°C compared to the ~70°C of the Green River Formation (Cole, 1985; Williamson, 1987).

Dating of authigenic feldspars in sedimentary rocks The radiogenic decay of 40K to 40Ar has had wide ranging applications for studies of Earth and planetary science (Jourdan et al., 2014). The half-life of the 40K to 40Ar decay of 1250 Ma is ideal for dating older rocks. Using this decay, older rocks will contain relatively less K and more Ar than younger rocks. The advent of the 40Ar/39Ar dating system has increased the efficiency of the K/Ar geochronological system, in which any rock with measurable amounts of potassium can be dated. Kastner and Siever (1979) briefly mentioned the possibility of Ar/Ar dating of authigenic feldspars, however they cited the difficulty of abundant detrital cores as precluding any reliable dating. Ar/Ar dating of authigenic feldspars has produced unreliable results for bulk separates, owing mostly to the difficulties of separating authigenic and detrital phases (Jourdan et al., 2014; Lee & Savin, 1985; Mark, Parnell, Kelley, & Sherlock, 2006; Mark et al., 2008). In these previous studies, all attempts to isolate and separate authigenic feldspars only included authigenic overgrowths on detrital grains within moderately (at least) hydrothermally altered rocks. The potential alteration of plagioclase to sericite (Verati & Jourdan et al., 2014) is a possible alteration which may influence the dating of authigenic albite. Parnell et al. (2011) analyzed authigenic orthoclase grains that occurred within a fluid escape structure in the Mesoproterozoic Stoer group, Scotland. They concluded that the grains provided a reliable date across four localities. Albite grains collected from the Lowville limestone, Pennsylvania, were identified as being formed authigenically during initial burial and early diagenesis based on four lines of evidence, (1) pure chemical composition (low potassium), (2) idiomorphic habit, (3) dearth of associated detrital material, (4) the cosmopolitan distribution within the rock (Honess and Jefferies 1940).

METHODS Sample collection and preparation Stromatolite specimens were collected, prepared, and imaged from the Mount Dunfee, Nevada (Fig. 3.1), and the Molly Gibson Mine, California (Fig. 3.2) localities of the ca. ~542 Ma Deep Spring Formation, the ca. 1.44-1.40 Ga Tieling Formation, China (Fig. 3.3), the ca. 760 Ma Flinders Range, Australia (Fig. 3.4), and the middle-late Miocene Duero Basin sedimentary succession (Fig. 3.5). All samples were cut to billet size and polished as preparation for imaging in the Scanning Electron Microscope (Fig. 3.6).

Scanning Electron Microscopy (SEM) Samples were polished down to .25-micron grit level and were carbon coated after polishing to provide a conductive surface for Scanning Electron Microscopy. The suite of Scanning Electron Microscopes of the Electron Microscopy and Imaging Lab (EMIL) at the University of Nevada Las Vegas was used to analyze the stromatolite samples for this study. Initial observations were conducted using a JEOL JSM-5600 with energy dispersive spectroscopy (EDS) for chemical analysis. The JEOL JSM-6700F Field Emission Scanning

Electron Microscope (FESEM) was used for high resolution imaging. Backscatter Electron (BSE) imaging was done to determine the relative mineral densities in real time. A mixture of point scanning and elemental mapping was accomplished with EDS. General working conditions for the SEM was 15 kV accelerating voltage with a variety of working distances for specific purposes. Due to the highly polished nature of the samples, secondary electron imaging was not used to produce images for this study.

Electron Probe Microanalysis Data were collected through wavelength dispersive X-ray spectroscopy (WDS) using an electron probe microanalyzer (EPMA; JEOL Superprobe). Compared to the semi-quantitative geochemical method of EDS, WDS provides fully quantitative data on elemental composition of minerals with comparison to standards, thus providing more accurate measurements of mineral composition and more reliable statistical analysis. Point analyses and elemental maps were collected from the two selected Deep Spring specimens. Point analyses provide the most accurate quantitative chemical data, while elemental maps allow for interpretation of the distribution of elements (and interpreted phases) within the specimens. Point analysis measured the “major” element package, which includes Mg, Ti, Ca, Na, Si, Fe, K, Al, P, Mn, Cl, F, Cr, Ba, and S. The probe condition was using 15kV, 10 nA, and 1 μm beam spot.

Mineral Separation To isolate both detrital and authigenic minerals from the carbonate matrix of the stromatolites, I macerated the one large stromatolite sample (Fig. 3.1C, 3.6) in hydrochloric acid until all carbonate was dissolved. I then diluted the remaining HCl until only water remained and filtered out the mineral residuals. Once separated from the stromatolite, a variety of techniques were used to separate albite, quartz, and orthoclase from each other. I first tried Lithium Metatungstate (LMT) heavy liquid mineral separation, however the mineralogical densities between the orthoclase and the albite were too similar to get reliable separation. I then tried hydrofluoric acid etching of the mineral grains, along with staining using Na3CO(NO2)6. I also placed grains onto carbon tape in a semi-grid and imaged with SEM/EDS to determine grain chemistry. Grains were then handpicked from the carbon tape.

PHREEQC Modeling Simple geochemical modeling was done using PREEQC Version 2, a program designed to perform aqueous geochemical calculations, to determine the stability of albite within the geomicrobiological system of the growing stromatolite. The model was setup using general Seawater parameters. Four assumptions were made for this model, (1) Biologic uptake of K completely removes K from the system initially, (2) there is a locally concentrated release of complexed Si and Al ions into solution during organic degradation, (3) addition of Na ions occurs from seawater sources, and (4) Si and Al are source only form the dissolution of detrital feldspars and quartz in the microbial mat. The baseline solution composition used in the PHREEQC simulations consists of seawater levels of calcium (.412 g/kg) and sodium (10.76 g/kg), relative amounts of Al and Si assuming complete dissolution of K-feldspar (38 g/kg of Aluminum, 100 g/kg of carbon, and 114 g/kg of silicon), and 100 g/kg oxygen. Chemical microenvironments created by mineral-attached microbes were completely ignored for these studies but likely have a large influence on mineral stabilities. Instead, the modeling consists of homogeneous solutions in a batch reactor.

RESULTS Geochemistry

Microprobe Data

I collected compositional data using wavelength dispersive x-ray spectroscopy to determine the mineral phases of the various silicates found within the stromatolites of the Mount Dunfee section of the Deep Spring Formation. Figure 3.7D summarizes these data by showing a typical microprobe dataset for each identified mineral. The entirety of the microprobe dataset was plotted onto the feldspar ternary diagram (Fig. 3.7A), showing the clear distinction between albite, orthoclase, and the gradational nature of Ca-plagioclase assemblage. Along with these feldspars, quartz, calcite, and a K-rich aluminosilicate are also found in varying amounts depending on the stromatolite and the location within each stromatolite. Mineral formulae were calculated based on the typical microprobe dataset for albite, orthoclase, Ca-plagioclase, and quartz (Fig 3.7C). The albite formula is often inconsistent due to the Na migration during electron beam interaction, calcite inclusions within the grain, and analysis of multiple crystals in the same beam spot. The plagioclase within the stromatolites contains a wide range of Ca, from .072 wt% to 11.29 wt% (Fig. 3.7A). Some of this calcium can be explained through the inclusion of nearby calcite in beam spot and calcite inclusions within the albite grains. Even though plagioclase has a wide variety of calcium inclusion (Fig 3.7A), albite-Ca-plagioclase mineral associations are often sharp in backscatter electron imaging, with a defining line between the Ca-plagioclase core and albite rims (Fig. 3.7B, Fig. 3.8H-I) or involving complex mixtures of albite and gradational Ca-plagioclase with the same defining line between them (Fig. 3.8E) and may include quartz within the mixture (Fig. 3.9A-C). This clear differentiation in backscatter imaging may be the result of the gap in the ternary diagram (Fig. 3.7A) between 90/10 (Ab/An) and ~83/17 (Ab/An).

In comparison, the lack of defining lines between some albite and some Ca-plagioclase within the mineral assemblages of the Molly Gibson Mine is due to the complex textures and compositional gradient between them (Fig. 3.9).

PHREEQC

Preliminary PHREEQC modeling results suggest that albite is thermodynamically supersaturated in the model conditions, with saturation indexes ranging from 17.0-102.64. Albite and Ca-plagioclase are often roughly equally supersaturated to kaolinite, which is a common product of feldspar dissolution in soils. Results show that the addition of seawater-sourced Na ions always results in the supersaturation of albite, assuming the necessary Si and Al ions are available. Similarly, Ca-plagioclase is consistently supersaturated due to the high levels of Ca within a stromatolite.

Unusual Mineralogical Features

Albite-orthoclase Textures

Many examples of semi-rounded orthoclase grains from the Mount Dunfee section of the Deep Spring Formation were found to be coated with a patchy rim of albite (Fig. 3.6, Fig. 3.8A- D) which often contains captured calcite (Fig. 3.8E-G). These core-rim textures were identified from Scanning Electron Microscopy of both thin sections (Fig. 3.8A-C) and also on grains collected from maceration of the stromatolite (Fig. 3.8D). Albite rims rarely fully envelope the orthoclase grains. They provide a patchy coating on most grains. Examples of this core-rim association of orthoclase and albite were only found at the Mount Dunfee section. The Flinders Range stromatolite shows potential orthoclase-albite core-rim textures (Fig. 3.4). If a better polish, or a different cut through the sample had been made, it is possible that the albite within the calcite matrix would have been touching the orthoclase grain (Fig. 3.4).

Albite-Ca-plagioclase Textures

Several examples of Ca-plagioclase-albite core-rim textures were found throughout two of the three stromatolite horizons of the Deep Spring Formation at Mount Dunfee (Fig. 3.8H-I). Within the Mount Dunfee samples, this feldspar association was in the form of Ca-plagioclase cored grains found closely spaced within a small area of the sample and nowhere else (Fig. 3.8H- I). The albite rim is often more advanced on the Ca-plagioclase cores (Fig. 3.8H-I) than it is for the orthoclase cores (Fig. 3.8A-D). Similar to the orthoclase cores, calcite is often captured between the Ca-plagioclase core and the albite rim (Fig. 3.8H-I), however, unlike the orthoclase cores, some of the Ca-plagioclase cores show carbonate capture similar to albite (see later in Results, Fig. 3.8I) and often shows large pore spaces (Fig. 3.8H-I).

The albite-Ca-plagioclase mineralogical textures within the Molly Gibson Mine stromatolite are much more intricate, displaying a complex mixture of albite-quartz-Ca- plagioclase-calcite in relatively large grains (50-100 μm diameter grains) (Fig. 3.9A-C, F). The albite-Ca-plagioclase associations within the Molly Gibson Mine stromatolite sample show a remarkable compositional gradient in which the plagioclase ranges from pure albite to a feldspar with 10% calcium (Fig. 3.9A-B).

Calcite Capture

Stromatolite-matrix calcite is often found contained within albite (Fig. 3.8E-G, Fig. 3.10A, B,C D, G, 3.7B) and more rarely within Ca-plagioclase (Fig. 3.8E,I)), but never within orthoclase or quartz. Calcite inclusions are often found between the orthoclase core and the albite rim (Fig. 3.8A-C, G). These calcite grains are highly diverse in size and morphology, and they are chemically indistinguishable from the surrounding calcite matrix of the stromatolite.

Albite and Quartz Micrograins

Deep Spring stromatolites contain abundant, morphologically diverse, and generally 1-5 μm size grains of albite and quartz with a nearly ubiquitous distribution (Fig. 3.8J-K, Fig. 3.11). Due to the similar density of quartz and albite (albite = 2.61-2.63; quartz = 2.6-2.65), backscatter electron (BSE) imaging on the SEM results in the same grayscale color for both phases. However, elemental mapping of key areas shows that both quartz and albite appear evenly dispersed throughout much of the stromatolite matrix with the exact same morphological diversity; this makes them impossible to identify with BSE alone. These small grains of quartz and albite often capture carbonate grains and seem to amalgamate into larger grains.

Micrographic-like textures

Micrographic textures occur in igneous rocks from the co-precipitation of feldspars and quartz during rapid crystallization of magmas, typically during sudden changes in pressure of a magma chamber. Micrographic texture indicates the relatively rapid precipitation of quartz and a feldspar from the same solution (Seclman & Constantinescu, 1972). The mineralogical associations of feldspars, quartz, and calcite within the Deep Spring stromatolites display very similar textures to the micrographic textures rarely seen in some igneous rocks (Fig. 3.10). Most commonly, albite and calcite show remarkable micrographic-like textures (Fig. 3.8A-I, Fig.

3.10A-C, F), often with quartz being mixed in. Quartz-calcite (Fig. 3.10C,D,G, Fig. 3.8K), albite-quartz (Fig. 3.10C,D,G), and, most remarkably, the albite-Ca-plagioclase-quartz-calcite micrographic-like textures (Fig. 3.9). Micrographic-like textures found in the stromatolite laminae differ dramatically from those found within the interstromatolite zone between columnar stromatolites in both overall morphology, composition, and amount of phases intermixing. Albite-calcite micrographic-like textures within the stromatolite typically include a small albite grain within the stromatolite calcite matrix (Fig. 3.8A-I, Fig. 3.10A, B, F). Comparatively, albite-calcite micrographic-like textures in the interstromatolite zone (Fig. 3.10C) occur with greater amounts of albite relative to calcite in a detrital matrix with a mix fine and coarse grains (left side of 12B). An unusual feature of the Molly Gibson Mine stromatolite is the presence of mica grains show evidence of quartz alteration (Fig. 3.9E). Mica grains typically show little to no alteration except the quartz replacement features.

Cathodoluminescence

Within the Mount Dunfee stromatolites, the orthoclase and large quartz grains in the interstromatolite zone display bright luminescence (Fig. 3.12). Albite within the interstitials and

91 the calcite matrix display dull luminescence (Fig. 3.12). No other samples were imaged using cathodoluminescence.

Etch pits on orthoclase grains

Etch pits were found on orthoclase grains which were chemically macerated from the calcite matrix of the stromatolite using HCl and imaged using the SEM (Fig. 3.8D). These etch pits are similar to previously reported, microbially mediated etch pits on feldspar surfaces

(Barker et al., 1998; Rogers et al., 1998; Fisk et al., 1998; Furnes et al., 2004). The etch pits found in the Deep Spring orthoclase are roughly 5 microns long, 3 microns wide, and 5 microns deep (Fig. 3.8D). Etch pits were only found on orthoclase grains with patchy albite rims (Fig. 3.8D).

K-rich aluminosilicate

A heterogeneously distributed K-rich aluminosilicate can be found in curvilinear concentrations throughout all Deep Spring Formation stromatolites as well as the Tieling Formation (Fig. 3.13). Rare examples of these K-rich aluminosilicate grains embedded into albite were found within the Mount Dunfee stromatolite (Fig. 3.13C). Due to the small microcrystalline nature of the mineral and the close proximity to other minerals, I was unable to obtain precise chemical measurements on these K-rich aluminosilicates using the microprobe, and no identification of the mineral phase was determined. Elemental mapping of the Deep Spring Stromatolites shows the linear distribution of these K-rich aluminosilicates (Fig. 3.11A-

D).

Crystal Twinning

The albite minerals found within the stromatolite samples of the Mount Dunfee section of the Deep Spring Formation do not show twinned crystal behavior. No other stromatolite samples were prepared for twinning analysis.

Other Stromatolite Localities The majority of the previously described data were from the Mount Dunfee (Fig. 3.1) and Molly Gibson Mine (Fig. 3.2) sections of the Deep Spring Formation. In an attempt to determine whether the above micrographic-like textures are common in other stromatolites, I prepared and imaged several other stromatolites from around the world which were provided to me by other researchers. Other localities include the Tieling Formation, China (Fig. 3.3), the Flinders Range, Australia (Fig. 3.4), the Duero Basin, Spain (Fig. 3.5), and the Chuar Group in the Grand

Canyon. Similar mineralogical assemblages (but not textures) were found in most stromatolite which were prepared and imaged. Differences between localities exist in the mineralogical and texture relationships. The stromatolite from the Chuar Group did not show any signs of feldspar or quarts.

Duero Basin

The early-middle Miocene lacustrine sediments from the Duero Basin, Spain, contain abundant stromatolites with concentrations of quartz, orthoclase, albite, and an amorphous Mg- silicate within an organic matrix (Fig. 3.5, see Fig. 11 of Sanz-Montero & Rodríguez-Aranda (2009) and Fig. 14 of Sanz-Montero, Rodríguez-Aranda, & Garcia del Cura (2009)). Sanz- Montero & Rodriquez-Aranda (2009) previously reported detrital “microperthic” textures within these stromatolites, which included albite-orthoclase textures. In a stromatolite sample provided by Maria Sanz-Montero, mineralogical textures within authigenic opal are found which consist of microbial cells mineralized within silica nanospheres and dolomite (Fig.3.5) (Sanz-Montero,

Rodriquez-Aranda, & Garcia del Cura, 2008) within the an organic matrix of “fossil extracellular polymeric substances (ESP)” (Sanz-Montero & Rodríguez-Aranda, 2009, p. 146). This organic matrix shows elevated levels of Mg and Si, possibly representing the amorphous Mg-silicate so often found within microbial mats (Burne et al. 2014; Pace et al. 2016; Tosca et al. 2011; Zeyen

94 et al. 2015). Indeed, examples of this Mg-silicate is found surrounding a detrital orthoclase grain (Fig. 3.5B) in a similar textural relationship to the albite-orthoclase rim-core textures seen within the Mount Dunfee stromatolites (Fig. 3.8A-D). The amorphous Mg-silicate at the bottom left of Fig. 3.5C shows similar wavy textures to the K-rich aluminosilicate within the Deep Spring Formation (Fig.12.B). The highly unusual wormy opal textures seen within this sample are localized into small concentrations (Fig 3.5C). Some opal grains display a gradient from less alteration (less wormy opal textures) to hollow spheres. These hollow spheres were previously reported to occur within these stromatolites and interpreted as mineral-coated microbial cells (Sanz-Montero et al., 2008, 2009).

Flinders Range

Examples of albite-orthoclase-quartz associations were found within a small Baicalia burra stromatolite sample provided by Carol Dehler of Utah State University. While the polishing of the sample was not ideal, the stromatolite contained small areas with concentrations of albite, orthoclase, and quartz in the calcite matrix (Fig. 3.4). No examples of direct contact between albite and orthoclase grains were found, however quartz was in direct contact with orthoclase in a core-rim texture (Fig. 3.4). It is likely that a different cut of the section would result in direct contact of the albite and orthoclase in a similar core-rim texture as described from the Deep Spring stromatolites (Fig. 3.8A-D). More investigation of this sample is needed to fully characterize the mineral assemblage.

Tieling Formation

The columnar stromatolites of the Mesoproterozoic upper Tieling Formation (ca. 1.44 to 1.40 Ga) formed from the trapping of fine-grained carbonate mud in the sticky surface of

95 growing microbial mats (Tosti & Riding, 2017). The stromatolite sample from the Tieling Formation, China, showed the lowest amounts of feldspar and quartz, compared to other stromatolite samples (Fig. 3.3). This is in line with previous observations of the stromatolites (Tosti & Riding, 2017). No significant textural relationships were identified between the orthoclase, quartz, and K-rich aluminosilicate found within the stromatolite (Fig. 3.3). No albite was found; however more time and a better polish of the section may provide better results. Previous studies of the Tieling Formation stromatolites have shown extensive glaucony mineralization (Tang et al., 2017), which may or may not be biologically mediated (Purnachandra Rao et al., 1993; Tang et al., 2017).

Chuar Group

One stromatolite (Inzeria stratifera) specimen from the basal Jupiter Member of the

Galleos Formation (ca. <770 ma), Chuar Group, did not contain any silicates. More investigation of this sample is needed to fully characterize the mineral assemblage.

Grain separation The largest stromatolite specimen from the Mount Dunfee section of the Deep Spring Formation (Fig. 3.1C, Fig. 3.6) was completely dissolved in hydrochloric acid and the mineral residuals were collected. I attempted to separate orthoclase from albite using a variety of methods, which are detailed below.

Lithium Metatungstate Heavy Liquid

I attempted to separate albite from orthoclase grains from the maceration residuals using Lithium Metatungstate Heavy Liquid. These attempts resulted in a no significant change in albite concentration.

HF Etching and Na3CO(NO2)6

Separation of albite from orthoclase was attempted using HF acid etching and

Na3CO(NO2)6 staining. Some orthoclase was found to have a very slight yellow coloration after staining. Yellow-stained grains were removed from the residuals. However, SEM/EDS analysis of the remaining residuals resulted in only a slight increase in albite concentrations, with a composition of 50% albite, 20% quartz, and 30% orthoclase.

SEM imaging and Hand Picking

The most successful attempt to separate albite from orthoclase in the maceration residuals was the result of placing the grains in a semi-grid arrangement on carbon tape and imaging each grain with SEM/EDS point analysis. I was then able to pick off individual grains of albite from the carbon tape, giving me a 100% separation of albite from orthoclase. However, due to the

97 need for SEM/EDS analysis and the associated costs, this process is not practical on a large enough scale to produce a datable quantity of albite for Ar/Ar dating.

DISCUSSION Alternative Hypothesis for the Origin of Micrographic-like Textures The unusual micrographic-like characteristics of the feldspars, quartz, and calcite within the stromatolites of the Deep Spring Formation suggest unusual processes were at play in their formation. Herein, I present alternative hypothesis to the origin of these micrographic-like textures: (1) igneous detrital grains, (2) albitization, and (2) stromatolite-authigenesis. The later hypothesis is a novel model for the authigenic, syndepositional origin of albite and Ca- plagioclase which utilizes previously reported microbe-mineral interactions (dissolution/precipitation) to explain the complex textures seen within the Deep Spring stromatolites.

Igneous Origin

The striking textures found within the stromatolite samples of this study likely result from vastly different geochemical conditions which are known to produce known mixtures of quartz and single compositional feldspars such as myrmekitic, granophyric, and micrographic textures (Barker, 1970; Smith, 1974,). The use of the term “micrographic-like” textures throughout this chapter emphasize the similarity between the micrographic texture seen in igneous rocks and those seen within the Deep Spring Stromatolites. Therefore, the feldspar- quartz-calcite textures detailed in this study may have originated within a magma chamber long before being deposited into the stromatolite microbial mats. The orthoclase-albite and Ca- plagioclase-albite core-rim textures are commonly found within magma chambers which have experienced a significant change in chamber dynamics. Similarly, the micrographic-like textures seen within the studied stromatolites have obvious similarities to previously described complex textures between feldspars and quartz, resulting from the changing dynamics of a magma chamber (hence the term “micrographic-like”). For example, within the Deep Spring formation,

99 micrographic-like textures are seen amongst albite-quartz (Fig. 3.10C,D,G), albite-calcite (Fig. 3.8A-I, Fig. 3.10A-C,F), quartz-calcite (Fig. 3.8K, Fig. 3.10C,D,G), and albite-Ca-plagioclase- quartz-calcite (Fig. 3.9) mineral assemblages. For this alternative hypothesis to be correct, the complex mineralogical associations within the stromatolites would have originated within a distant magma chamber, been uplifted and weathered, and the eroded sediment carried to sea and deposited into the stromatolites at both the Mount Dunfee section and the Molly Gibson Mine section of the Deep Spring

Formation. After deposition, dissolution of the detrital micrographic grains would occur, and calcite precipitation infills the voids. Similarly, the same sequence of events would have had to occur at the Miocene Duero Basin section, Spain, where the detrital “perthitic” grains (Sanz- Montero & Rodríguez-Aranda, 2009, see discussion for reinterpretation of these perthitic textures) and unusual wormy opal textures would have been deposited into the stromatolites and undergo similar dissolution and calcite precipitation. I view it as unlikely that two stromatolite- bearing formations would have experienced the deposition of grains with highly unusual igneous textures and undergone similarly rare diagenetic conditions. Additionally, the highly unusual wormy opal textures in the Duero Basin stromatolites, previously attributed to authigenic precipitation with EPS mat structure (Sanz-Montero et al., 2008) have no known igneous examples of similar texture. The most parsimonious explanation is that they are authigenic within the stromatolites.

Furthermore, the striking micrographic-like textures found within the stromatolite sample from the Molly Gibson Mine section of the Deep Spring Formation expand these unusual mineralogical relationships beyond the simple two-phase mixing of typical micrographic textures in igneous rocks and those seen in the Mount Dunfee stromatolites. Within the Molly Gibson stromatolite, I found many examples of grains which contain complex mixtures of albite, pure quartz, pure calcite, and a gradient of Ca-plagioclase (Ca-plagioclase) ranging from 3 wt% to 10

100 wt% calcium in SEM/EDS data (Fig. 3.7A,C-D, Fig. 3.9 A-B). These complex mineralogical mixtures are unlike anything previously reported from igneous mineralogy research. In comparison, the Ca-plagioclase found within the Mount Dunfee section appears merely as cores, with roughly 4 wt% calcium from microprobe data (Fig. 3.7, 3.8). This suggests that smaller scale variability within the sedimentary environment resulted in these textures. While the two Deep Spring sections display similar mineral textures and assemblages, the differences between locations are important clues into their origination. The interstromatolite zone contains less calcite than the stromatolite itself, resulting in albite-calcite micrographic-like textures which contain higher amounts of albite (Fig. 3.10C). The textural and mineral assemblage differences between stromatolite samples within the Mount Dunfee section, and between stromatolite laminae and interstromatolite zone, suggest that small scale environmental conditions are responsible for these differences, as opposed to detrital input of different lithologic features into specific sites of the stromatolites.

Albitization

Kastner and Siever (1979) identified 4 characteristics of authigenic albite in carbonates due to the hydrothermal alteration of feldspars: 1) end-member composition, 2) dull cathodoluminescence, 3) non-twinned crystal behavior, and 4) high positive δ180 values. To generate the data presented herein, I used the electron microprobe to determine the composition of the albite, SEM cathodoluminescence to show the lack of luminescence in albite, thin section analysis to show the lack of twinning crystal behavior of the albite. I was unable to determine the oxygen isotopic values of the albite. Based on these metrics for determining authigenicity for feldspars in carbonate rocks, it is likely that the albite within stromatolites represent authigenic albite. However, the albite grains within the Deep Spring stromatolites differ from those produced by typical albitization in that not all of the albite is perfect endmember albite. Rather,

101 the Deep Spring albite grains display a gradient in plagioclase compositions from pure endmember albite to 70/30 (Al/An) (Fig. 3.7A). The albitization of Ca-plagioclase and alkali feldspar is an alternative explanation for the micrographic-like textures seen within the stromatolites detailed in this chapter (Boles, 1982; Morad et al., 1990). Albitization occurs during the hydrothermal brine alteration of feldspars, resulting in the precipitation of authigenic albite, primarily as internal replacement along crystallographic planes and fractures, and as patchy rims (Boles, 1982; Kastner & Siever, 1979;

Milliken, 1989; Morad et al., 1990). Authigenic albite is always closely associated with the albitized feldspar grain, either through rims or direct pseudomorphic replacement. Albitized feldspar grains display clear grain boundaries after albite replacement has occurred, illustrating the 1:1 replication of detrital feldspars with authigenic albite during solution-precipitation processes (Boles, 1982; Kastner and Siever, 1979). This is in clear contrast to the authigenic albite detailed in this study, in which no internal replacement of detrital grains has occurred (Fig. 3.8A-C, H-I). Additionally, the high abundance of intermixed quartz and albite micrograins (Fig.

3.8J-K) occurring throughout the stromatolite matrix is in stark contrast to albitization studies, as no previous albitization studies have reported abundant micrograins of any phase that are not internally replacing larger feldspar grains. Previous reports emphasize the importance of illite- albite mineral assemblages in albitization processes (Mu et al., 2016; Turner and Fishman, 1991). There is no evidence of this mineral assemblage within the studied stromatolites.

The high purity of albitization-derived authigenic albite is always discussed as an important characteristic (Kastner & Siever, 1979; Miliken, 1989), often with >99% albite purity of the euhedral albite crystals which pseudomorph the original detrital grain. Na migration during electron beam interactions within the microprobe can easily explain the non- stoichiometric nature of some authigenic albite which contain “no petrographic evidence of non- feldspar inclusions” (Miliken, 1989). The authigenic albite within the stromatolites of this study

102 are a mixture of albite of seemingly high purity and a gradation of Na-Ca plagioclase with abundant calcite inclusions (Fig. A, C-D). Experimental studies of plagioclase albitization show the direct replacement of Ca- plagioclase with chemically pure albite, producing a core-rim texture with clearly defined contact between phases (Hövelmann et al., 2010). These experimentally formed textures are very similar to the Ca-plagioclase-albite core-rim textures seen in one stromatolite from the Mount Dunfee section (Fig. 3.8H-I) but they are drastically different from the Ca-plagioclase textures found in the Molly Gibson Mine section (Fig. 3.9A-C,F). Nor are they present among the textures found in the Duero Basin stromatolites (Sanz-Montero & Rodríguez-Aranda, 2009). Similarly, the wormy opal textures seen in the Duero Basin stromatolite have no known igneous comparisons (Fig. 3.5). While albitization could explain the Ca-plagioclase cores at Mount Dunfee, it cannot explain the myriad other textures involving the Na-Ca plagioclase gradient and quartz textures seen in the many other stromatolites in this study.

Determination of Authigenicity

The lack of illite in all analyzed stromatolite samples, the lack of internal replacement, the morphological variation of mineral assemblages, and the presence of a Na-Ca plagioclase gradient in the stromatolites studied here, all suggest that the albite formed through different processes than albitization. Likewise, it is unlikely that the myriad mineralogical textures detailed here originated from igneous sources. Next, I detail a geomicrobiological model which could explain the formation of these unusual micrographic-like textures as syndepositional, authigenic precipitates during the mineralization of the microbial mat. This stromatolite- authigenesis model utilizes microbe-mineral interactions to dissolve and concentrate ions and to precipitate authigenic minerals.

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Stromatolite-authigenesis Model The interpretation of some of the silicates within the stromatolites detailed in this chapter as having formed authigenically with the calcite of the stromatolite matrix requires a biogeochemical model to explain the unusual occurrence of these textures and mineral assemblages. I concur with previous authors in suggesting that authigenic silicates often form under these conditions, but I expand on their models to include plagioclase as an authigenic silicate within stromatolites along with the previously reported quartz (del Buey et al., 2108;

Sanz-Montero & Rodríguez-Aranda, 2009), Palygorskite-sepiolite (del Buey et al., 2018; Perri et al., 2017), and Kerolite (Zeyen et al., 2015), along with the various authigenic, poorly crystalline silicates, oxides, and oxyhydroxides, etc., reported to occur on the surfaces of microbial cells. This stromatolite-authigenesis model, occurring within the complex biogeochemical conditions of the organomineralization of a microbial mat (Dupraz et al., 2009), follows 4 distinct phases: (1) the enhanced dissolution of detrital orthoclase due to organic-metal complexation of Si and Al, as well as the bioextraction of K ions for osmotic regulation of cells within the stromatolite, (2) the capture of dissolved ionic species within the complex biochemical EPS framework of the microbial mat, (3) the degradation of organic matter within the mineralizing portions of the stromatolite, and (4) the release of complexed ions into solution where they combine with seawater-sourced ions to produce authigenic albite, Ca-plagioclase, quartz, and a K-rich aluminosilicate (Fig. 13). Preliminary PHREEQC modeling suggest that albite, quartz, Ca-plagioclase, and calcite become thermodynamically supersaturated in these geochemical conditions, possibly resulting in authigenic precipitation of all phases simultaneously. Below, I detail each step in this model.

Dissolution of detrital phases

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The stromatolite-authigenesis model begins with the detrital input of orthoclase and minor amounts of quartz, based on the presence of large grains within the interstromatolite zone (Fig. 3.14). Dissolution of these detrital grains begins in the intense geomicrobiology activity within the upper photosynthetic section of the microbial mat and continues as the grains become smothered during stromatolite growth. The interaction of the detrital minerals with low molecular weight metabolites has been shown to drastically increase the dissolution rates of silicate minerals (Ullman, 1996; Vanevivere et al., 1994; Welch and Vandevivere, 1994; Chen et al., 2018; Lee et al., 2000; Iverson et al., 1978; Qureshi et al., 2018), resulting in either the partial dissolution of the detrital grains (Fig. 3.14). Etch pits observed on albite-rimmed orthoclase grains from the Mount Dunfee section show the results of this process (Fig. 3.8D). Many mechanisms explain these increasing dissolution rates, from the destruction of toxic molecules in the environment to the direct bioextraction of nutrients (Kraegeloh & Kunte, 2002; Qureshi et al., 2018). Often, this increased dissolution rate may be simply due to passive interactions with microbial metabolites to the mineral environment in which the microbes live. However, Qureshi et al. (2018) demonstrated the direct bioextraction of 34 ppm, 42 ppm, and 28 ppm potassium from the surface of feldspar grains in their experimental setup of several different microbial colonies. Similarly, Chengsheng et al. (2014), Meena et al. (2013), Hashem et al. (2016), Sugumaran et al. (2017), Maurya et al. (2014), and Ren et al. (2015) all reported bioextraction of potassium ions from feldspars, mica, mica, muscovite mica, mica, mica, respectively, reporting potassium solubilization levels of 4.5 ppm, 17.33 ppm, 13.24 ppm, 4.29 ppm, 23.98 ppm, and 7.42 ppm, respectively. The bioextraction of potassium from feldspar and mica results in the concurrent and proportional release of Si and Al from the grain surface. Likewise, Ca- plagioclase (Ca-plagioclase) likely experiences an increased dissolution rate in the early non- mineralization to early mineralizing stages of the microbialite due to the intense alkalinity engine occurring within the microbial mat (Deco et al., 2005; Dupraz et al., 2004, 2009).

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Complexation of ionic constituents with organic molecules

The physicochemical properties of EPS and other microbial metabolites, primarily the cation binding capacity, has a dramatic influence on the composition and morphology of mineral precipitates during organomineralization of the microbial mat (Braissant et al., 2007, 2009; Dupraz et al., 2009; Ueshima & Tazaki, 2001). Studies of modern microbial mats and their carbonate mineralization have led to a detailed understanding of the important role that EPS plays in the early stages of mineralization (Dupraz et al., 2009). The EPS structure of the microbial mat provides a sink for dissolved ions which would otherwise be lost into solution, it provides important structure for early gel formation, and it provides abundant nucleation sites for later stage mineralization of gels (Dupraz et al., 2009). To validate this important point, Ueshima and Tazaki (2001) experimentally demonstrated within a ferrosiliceous solution that the formation of layer-silicates (nontronite) occurred preferentially on EPS and rod-shaped iron hydroxides in the external solution (Ueshima & Tazaki, 2001).

All dissolved ionic species are then complexed into the surrounding organic molecules, either through direct assimilation into the cell bodies for osmotic balancing (Kraegeloh & Kunte, 2002; Qureshi et al., 2018) and other metabolic purposes or through passive interaction with the charged organic molecules in the EPS and metabolites solution (Chen et al., 2018; Iverson et al., 1978; Lee et al., 2000; Liu et al., 2006; Ueshima & Tazaki, 2001; Ullman, 1996; Qureshi et al.,

2018; Vandevivere et al., 1994; Welch & Vandevivere, 1994). This complexation of Si, Al, and K ions concentrates ionic constituents around the dissolving grain where they remain until the organic molecules are no longer able to complex them. Del Buey et al., (2018) showed that hydrated (living) microbial mat material had a higher amount of Mg, O, Si, and Al, than the dehydrated mat (del Buey et al., 2018). Similarly, Obst et al., (2009) showed the preferential complexation of Ca ions into EPS surrounding cells, in which early aragonite crystals nucleated

106 from before the precipitation of calcite (Obst et al., 2009). The cation binding capacity of EPS and other microbial metabolites allows most (if not all) of the dissolved ionic constituents to remain within the general vicinity of the original detrital grain. In addition to ionic binding to EPS, microbial cell walls are known to absorb large amounts of a variety of ionic constituents (Beveridge, 1993; Eushima & Tazaki, 2001; Fein et al., 1997; Konhauser & Uruttia, 1999; Schultze-Lam, Harauz, & Beveridge, 1992; Schultze-Lam et al., 1996; Tashiro & Tazaki, 1999; Tazaki, 2005; Theng & Orchard, 1995; Tuck et al., 2006). The ion binding sites of organic molecules in the form of EPS, metabolites, or microbial cells walls, provide nucleation sites during dehydration and release of absorbed ions. The different paths for the various ions released during orthoclase dissolution results in concentrations of ions in different sections of the stromatolite. Since Si and Al have no biological utility, they quickly become passively absorbed onto EPS and microbial metabolites (Obst et al., 2009; del Buey et al., 2018). Considering the fact that the majority of stromatolite microbial mat is made up of EPS, Si and Al ions will become absorbed quickly and not concentrate in any particular area accept around the original detrital orthoclase grain. Potassium ions, on the other hand, are known to be used in osmotic regulation in saline and hypersaline waters (Kraegeloh & Kunte, 2002). The selective bioextraction of K ions from feldspars grains results in the concentration of K where cells are actively thriving. Therefore, K ions will cycle through the stromatolite biochemical environmental within living cells, leading K to be concentrated into the last vestiges of microbial life within the mineralized portion of the stromatolite. This is likely why dark bands of K-rich aluminosilicate exist throughout the stromatolite (Fig. 3.13) and quartz/albite micrograins are found everywhere (Fig. 3.8J-K).

Early carbonate precipitation and organic degradation

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The demonstration of the higher concentration of Mg, O, Si, and Al ions within hydrated microbial mats compared to the dehydrated microbial mats illustrates the importance of both the concentration of ionic constituents within the mat and their release during dehydration (del Buey et al., 2019; Ueshima & Tazaki, 2001). As carbonate mineralization begins below the actively photosynthesizing upper layers of a stromatolite, organic degradation releases absorbed ions which begin to form amorphous Si/Al gens on the dehydrating EPS and microbial cell walls (Fig. 3.14). Due to the ability for the living cells to move within the mat along biochemical gradients, the K they bioextracted will continue in its bioavailable form during this early mineralizing stage of stromatolite formation. The transitory nature of EPS within the stromatolite leads to its eventual dehydration and degradation during carbonate mineralization underneath the actively growing microbial mat surface. Since the EPS structure of the mat degrades more or less homogenously (lasting longest where living microbial cells are thriving), the complexed Si and Al ions will be released into the closing biogeochemical system of the mineralizing stromatolite, where mineralizing forces transform gels into poorly crystalline precursors (Sánchez-Navas et al., 1998; Dupraz et al., 2009) or straight to the albite, quartz, and K-rich aluminosilicate that I see today (Fig. 3.8-3.10). The inclusion of seawater-sourced Na into the stromatolite biogeochemical environmental provides the necessary ionic constituents to precipitate as much authigenic albite as there is dissolved orthoclase-sourced Si and Al. Preliminary PHREEQC modeling suggests that the inclusion of any amount of Na into a system consisting of seawater, plus the dissolved Si and Al from orthoclase dissolution, results in the supersaturation of albite. Similarly, Ca- plagioclase is supersaturated in the model due to the inclusion of Ca. However, PHREEQC results show that Ca-montmorillonite is often the most supersaturated mineral within the system due to the inclusion of Ca, Si, and Al into the model. Ca-montmorillonite was never identified within any of the stromatolite samples.

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It is not possible to determine the exact mineralization sequence with the available data and the nature of organomineralization within the complex biogeochemical conditions within stromatolites, including mm-scale redox changes which effect precipitation/dissolution dynamics (Visscher & Stolz, 2005; Dupraz et al., 2009). The authigenic silicates found within the stromatolite calcite matrix either (1) precipitated in the very late stages of mat mineralization after the majority of calcite had precipitated and mineralized the microbial mat EPS, or (2) began precipitating early during calcite mineralization of the EPS due to the formation of microenvironments located on felspar grains and within the mineralizing calcite matrix (Lee et al., 2008). The ubiquitous presence of micrograins of albite and quartz, typically less than 2 μm in diameter, throughout the stromatolite calcite matrix (Fig. 3.8J-K) suggests that they were co- precipitates with calcite, forming the described micrographic-like textures from the earliest stages of calcite mineralization.

Complete mineralization of microbial mat

As stromatolites grow, new layers of active microbial mat are added to the photosynthetic upper portion of the stromatolite, smothering the lower layers, and forcing them to cycle through geochemical redox gradients (Visscher & Stolz, 2005). Completion of the carbonate mineralization of the mat EPS results in the closing of the chemical system. All absorbed Si and Al become released into the closing system and precipitate, either as poorly crystalline precursors or as the albite, Ca-plagioclase, and quartz that I see today. It is likely in the final stages of mineralization in which the remaining Ca ions become taken up by the precipitating

Ca-plagioclase when carbonate mineralization is less intensely utilizing available Ca. The difference in Ca-plagioclase morphology and chemical diversity between the Mount Dunfee stromatolites and the Molly Gibson Mine stromatolite provides important clues to their formation. The Mount Dunfee Ca-plagioclase occurs primarily as albite-rimmed grains (Fig.

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3.7B, Fig. 3.8H-I) with consistent composition and occasionally with captured calcite (Fig. 3.8I). In comparison, the Molly Gibson Ca-plagioclase occurs as complex mixtures with albite, quartz, and calcite, which display micrographic-like textures and are composed of a compositional gradient of Na/Ca plagioclase (Fig. 3.9). This clear morphological and compositional difference between the two localities likely represents the varying timing in which the phases became supersaturated within the mineralizing mats. The Mount Dunfee albite and Ca-plagioclase likely became supersaturated at different times relative to each other and the other authigenic phases, dependent on both the microenvironmental conditions of the mineralizing mat and on the aggressiveness to which the carbonate precipitation is taking available Ca ions. While I consider the unique micrographic-like textures found within the Molly Gibson Mine stromatolite as evidence of the unusual geomicrobiological origin of these minerals, I do not know what specific conditions would lead to such complex mixtures of phases in comparison to the textures seen at Mount Dunfee. Similarly, the slight differences between the micrographic-like textures seen within the stromatolite laminae and the interstromatolite zone, suggest that small-scale biogeochemical conditions led to these differences. The lower amount of calcite mineralization within the interstromatolite zone (Fig. 3.10C) suggest that a less active microbial community existed in this zone. The curvilinear distribution of K-rich aluminosilicates within the stromatolites suggest that they likely represent the last vestiges of microbial activity in the lower mineralizing sections of the microbialite during mineralization. The concentration of microbes in small curvilinear sections of the microbialite would release all K ions from their biologic constraints after their inevitable demise in the mineralizing environment (Fig. 3.14). The release of potassium ions along with other ions absorbed onto organic molecules (Si, Al) would allow for the precipitation of a K-rich aluminosilicate or a K-Si-Al precursor gel or poorly crystalline precursor mineral.

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The concentration of larger authigenic albite and quartz within the K-rich aluminosilicate lineations suggest that more intense authigenic silicate mineralization occurred during the later stages of mat mineralization (Fig. 3.13). The inclusion of the K-rich aluminosilicate within some albite grains (Fig. 3.13C) suggests that the two mineral phases were co-precipitates at some point during mat mineralization.

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Stromatolite-Authigenesis Model Justification

Authigenic silicates in carbonates

While previous studies on the albitization of carbonates emphasize the importance of high-temperature diagenesis/low-temperature metamorphosis, especially those on albite replacement of orthoclase (albitization) in carbonates (Saigal and Morad, 1988; Chowdhury and Noble 1993; Norberg et al., 2011; Porto da Silveira et al., 1991; Aagaard et al., 1990; El-Khatri et al., 2015; Alvaro and Bauluz 2008), these studies lack geomicrobiological circumstances which are inherent within microbialites. Thus, the door is open for novel interpretative processes which utilize the intense microbial activity within stromatolites to dissolve minerals, concentrate ionic species, and precipitate authigenic minerals through interactions with the extremely high concentrations of ionically charged biological molecules within the actively growing mat. Historically, micrographic textures were thought to be confined to high temperature igneous processes within a magma chamber. Researchers studying the seemingly authigenic albite within sedimentary rocks speculated that they were the result of hydrothermal brines as low as 70°C (Cole, 1985; Kastner & Siever, 1979; Nuccio & Johnson, 1988). The resulting mineral textures are similar to those seen in igneous micrographic rocks with the same mineralogies—authigenic albite, quartz, with detrital orthoclase. This study shows evidence which could expand the environmental circumstances in which authigenic albite, Ca-plagioclase, and quartz create intricate micrographic-like textures through the intense microbe-mineral interactions within mineralizing stromatolites.

Many previous examples of silicates, often in the form of Mg-silicates or amorphous aluminosilicates, have been reported from both ancient and modern microbial mats (del Buey et al., 2108; Perri et al., 2017; Sanz-Montero & Rodríguez-Aranda, 2009; Souza-Egipsy al. 2005; Zeyen et al., 2015). A recent paper by Zeyen et al., (2015) detailed their finding of amorphous

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Mg-silicates within recent lacustrine microbialites and suggest, along with others (Bontognali et al., 2014; Leveille et al., 2000a; Obst et al., 2009; Ueshima & Tazaki 2001,), that these amorphous silicates are precursor phases to a variety of more crystalline phases such as kerolite, chlorite, and even carbonates within microbialite biochemical frameworks. Leveille et al., (2000a) had previously proposed something similar, suggesting that the dehydration of EPS within the microbialite created a Mg-Si-O gel as a precursor to the fine-grained kerolite that they observed to be closely associated with bacterial EPS. Similarly, Obst et al., (2009) showed that an amorphous or nanocrystalline precursor CaCO3 phase precipitated within microbial mat EPS before a more crystalline calcite phase formed (Obst et al., 2009). Obst et al. (2009) and Perri et al. (2017) found palygorskite and potential precursor Ca-Mg-Si-Al-S amorphous nanoparticles within hypersaline microbial mats from Qatar (Obst et al., 2009; Perri et al., 2017). In the case of authigenic albite, it has been suggested that analcime may provide an early, hydrated precursor (Kastner and Siever, 1979). The thorough work of Tazaki and colleagues on the interaction between cellular EPS substrates and mineral precipitation has revealed the powerful nucleation ability of these biomolecules (Eushima & Tazaki 2001; Tazaki 2005, Schultze-Lam, Fortin, Davis, & Beveridge, 1996a; Schultze-Lam, Harauz, & Beveridge, 1992; Tashiro & Tazaki 1999; Theng and Orchard 1995). Tashiro & Tazaki (1999) (and later elaborated on by Eushima & Tazaki (2001)) described a model which invokes the importance of metal-ligand bonds in the interaction between EPS and dissolved ions to concentrate mineral constituents before precipitation has begun and resulting in the preferential precipitation of layer-silicates within the EPS (Tashiro & Tazaki, 1999; Eushima

& Tazaki, 2001). The potential importance of biological activity in the dissolution and precipitation of authigenic feldspars was downplayed in the Buyce & Friedman (1975) paper, in which authigenic orthoclase was found to be associated with cryptalgal structures (Buyce & Friedman,

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1975; Mazzullo 1976). Key components of the Mazzullo (1976) model of the authigenic feldspar involved the seawater-derived source of K ions which came from weathering of silicate rocks in the Adirondack lowlands (Mazzulo, 1976). The addition of enough K ions into the system to allow for the precipitation of authigenic orthoclase is difficult, given the relative lack of potassium in the ocean. In contrast, the precipitation of authigenic albite within similar environments is easier to reconcile considering the vast amount of sodium in marine waters.

Previously Misinterpreted examples of stromatolite-authigenesis

Two similar models to the stromatolite-authigenesis model were proposed by Sanz- Montero & Rodríguez-Aranda (2009) and del Buey et al. (2018), in which they detail the presence of detrital and authigenic silicates within Miocene and modern microbialites in Spain, respectively. Sanz-Montero & Rodríguez-Aranda (2009) detailed the authigenic formation of quartz, a poorly crystalline Mg-silicate, dolomite, calcite, Al-rich amorphous phases, and barite within minimally buried (~100m depth), Miocene, lacustrine microbialites. Their model shows the origin of these minerals as the result of microbial activity within the forming stromatolite dissolving detrital “microperthitic” orthoclase-plagioclase grains and the precipitation of the authigenic phases (Sanz-Montero & Rodríguez-Aranda, 2009). However, reinterpretation of the plagioclase as albite using their figure 13 elemental maps, shows the albite as a likely authigenic co-precipitate along with quartz and the other authigenic phases (Fig. 13, Sanz-Montero &

Rodríguez-Aranda, 2009). As seen within figure 11 & 13 of Sanz-Montero & Rodríguez-Aranda (2009), the albite within the “microperthite” grain is disassociated from orthoclase and surrounded by matrix calcite in the upper-right corner of the figure and coating the orthoclase in the bottom-right corner (Figure 11 & 13 of Sanz-Montero & Rodríguez-Aranda , 2009). Similarly, albite is found within the orthoclase grain directly connected to what they determined was authigenic quartz, seemingly precipitating within fractures of the detrital orthoclase. These

114 mineralogical textures are the same that I describe from the Deep Spring Formation as authigenic albite-quartz-calcite micrographic-like textures. In addition to the reinterpretation of “microperthitic” textures in the Duero Basin stromatolite, I prepared and imaged a new sample from the Duero Basin which shows highly unusual mineralogical textures. The matrix of the stromatolite is composed of an amorphous Mg-silicate within the “carbonaceous masses” that “are thought to represent fossil extracellular polymeric substances” (Sanz-Montero & Rodríguez-Aranda , 2009) within the Duero Basin stromatolite. This further illustrates the likely authigenic nature of these mineral assemblages. The textural similarity between the Mg-silicate organic matrix within the Duero Basin stromatolite (bottom left corner of Fig. 3.5C) and the K-rich aluminosilicate of the Deep Spring Formation (bottom left corner of Fig. 3.13B) suggests that the wavy linear concentrations of these phases result from the alignment of flat grains during mat mineralization. Since the amorphous Mg-silicate within the organic matrix of the Duero Basin stromatolite is clearly authigenic and biologically sourced, we can assume that the K-rich aluminosilicate of the Deep

Spring Formation is similarly produced. Based on my reinterpretation of the data presented in Sanz-Montero & Rodriquez-Aranda, (2009), and in light of the data presented in Sanz-Montero, Rodriquez-Aranda, & Garcia del Cura (2008), as well as the new images I collected (Fig. 3.5), the most likely explanation for the origin of the observed mineral associations in both the Deep Spring Formation and the Duero Basin is that they are the product of authigenic, co-precipitation of albite, calcite, quartz, and opal during microbial mat mineralization.

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Ar/Ar Dating The presence of abundant, individual (non-cored) authigenic albite within a carbonate matrix (Fig. 3.8J-K) allows for a potentially reliable separation of authigenic phases for 40Ar/39Ar dating. If the authigenic albite within the stromatolites of this study do indeed result from syndepositional geomicrobiological processes (Fig. 3.14), then the authigenic albite will record a Ar-Ar age of the formation of the stromatolite and the formation of the sedimentary rock unit. This could provide a new dating method for stromatolite-containing strata.

I made multiple attempts to separate the authigenic albite from the detrital orthoclase within the stromatolites, but with little success. A large stromatolite sample was completely dissolved in hydrochloric acid to isolate the silicate minerals. I used Lithium Metatungstate Heavy Liquids to separate the albite, quartz, and orthoclase residual grains, but the amount of albite concentration I could achieved was limited by the similar mineral density of the two minerals. I experimented with hydrofluoric acid etching of the residuals in combination with

Na3Co(NO2)6 staining to make identification of orthoclase easier under the microscope. The etching experiments yielded only a slight increase in albite concentrations within the residuals. Additionally, I sprinkled grains onto carbon tape which held a metal grid, and I analyzed the grains using SEM/EDS to determine composition. Albite grains were then collected from the carbon tape. While this technique did result in a 100% albite sample, the SEM and microscope time required to gather the necessary albite sample for Ar/Ar dating was unachievable within the current scope of the project. While researchers have successfully dated authigenic feldspars within carbonates, difficulties separating authigenic from detrital grains has made it unreliable (Kastner & Siever, 1979). When successfully separated and dated, authigenic feldspars often represent geothermal brine interactions. For this reason, little work has been done on isolating and dating authigenic feldspars within carbonates. No one has reported dating authigenic feldspars from stromatolites.

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This can be viewed as (1) a sign of presumed failure, (2) non-reporting of negative results, (3) no one has attempted to date them, or (4) no one has even looked for them. Previous studies of authigenic feldspars have focused only on micritic carbonates. The ability of microbial consortia to change the dissolution and precipitation dynamics of minerals may allow for a more reliable dating of authigenic feldspars within the Deep Spring Formation microbialites due to the lack of detrital orthoclase cores and remnant fragments in most authigenic albites within the Deep Spring stromatolites.

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CONCLUSIONS Based on our observations of the mineral assemblages within six stromatolites from four localities, three formations, and three continents, I detailed a geomicrobiological model which utilizes the intense biogeochemical conditions inherent within a forming stromatolite to precipitate authigenic albite, Ca-plagioclase, quartz, and a K-rich aluminosilicate during calcite mineralization. The model follows 4 phases, as described below (Fig. 3.14): (1) the enhanced dissolution of detrital orthoclase due to organic-metal complexation of

Si and Al as well as the bioextraction of K ions for osmotic regulation of cells within the stromatolite. (2) the capture of dissolved ionic species within the complex biochemical EPS framework of the microbial mat. (3) the degradation of organic matter within the mineralizing portions of the stromatolite. (4) the release of complexed ions into the closed system of the mineralizing mat solution and precipitate authigenic albite, Ca-plagioclase, quartz, and a K-rich aluminosilicate.

In particular, I am most interested in the albite component of the mineral assemblage due to its potential usefulness as a geochronological indicator. I base this model on the textural associations between the various mineral components within the stromatolite to develop a model which explains the albite as syndepositionally authigenic, likely forming at the same time as the stromatolite calcite or soon after with the maturation of silicate precursors such as analcime.

These mineral associations have been confirmed in three Deep Spring stromatolite samples from Mount Dunfee (Fig. 3.1, 3.8, 3.10), the Molly Gibson Mine section of the Deep Spring

Formation (Molly Fig. 3.2, 3.9), Miocene stromatolites from the Duero Basin, Spain (Fig. 3.5), and tentatively stromatolites from the Flinders Range, Australia (Fig. 3.4).

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Figure 3.1 Stromatolites of the Mount Dunfee section of the Deep Spring Formation. A) stromatolitic reef; B) Stromatolite reef horizons; C) stromatolite collected from float which contains both stromatolite laminae (left, gray) and the interstromatolite zone (right, orange); D) regional map showing the location of the Mount Dunfee Section and the Molly Gibson Mine section of the Deep Spring Formation; E) Stromatolite sample (red circle) collected from stromatolite horizon A. F) stromatolite sample (red circle) collected from stromatolite horizon D.

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Figure 3.2 Stromatolites of the Molly Gibson Mine section of the Deep Spring Formation. A) field of stromatolites; B) Detailed picture of the stromatolites; C) stromatolite sample (red circle) collected from the Molly Gibson Mine section; D) regional map showing location of the Mount Dunfee section and the Molly Gibson Mine section of the Deep Spring Formation.

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Figure 3.3 Representative mineral assemblage of the Tieling Formation stromatolite. A) Backscatter Electron image of the polished stromatolite sample; B) Potassium elemental map of section shown in “A”; C) Silicon elemental map of the section shown in “B”; D) Sodium elemental map of the section shown in “A”. Qz = quartz, C = Calcite, Kspar = orthoclase.

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Figure 3.5 Mineral Textures within the Duero Basin stromatolite. A) Diagonally oriented grain shows the wormy opal textures (which are composed of pure SiO2 in EDS spectrum) and associated elemental maps. Note the higher amount of alteration at the bottom left corner of the grain; B) orthoclase grain with a partial rim of Mg-Si organic matrix with little to no dolomite mineralization; C) broader view of grain in “A” showing the distribution of altered quartz grains within the Mg-Si organic matrix with associated dolomite precipitates. Dol = dolomite, Kspar = orthoclase, Mg-Si = amorphous Mg-silicate within organic matrix, Qz = quartz

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Figure 3.6 Stromatolite sample preparation and SEM analysis.

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Figure 3.7 Geochemistry of the Deep Spring mineral assemblage. A) ternary diagram showing feldspar chemistry; B) SEM/BSE image with interpreted mineral assemblage; C) Calculated mineral formulae for typical minerals of each variety; D) typical probe dataset for each mineral.

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Figure 3.8 Mineral textures found within the stromatolite found in float (Figure 3.1C, 3.6) at the Mount Dunfee Section of the Deep Spring Formation. A)-C Examples of albite-rimed orthoclase grains from polished thin section; D) albite-rimed orthoclase grain from maceration residuals showing etch pits; E-G) SEM/BSE images showing examples of calcite captured within albite (Ab) from polished thin section; H-I)SEM/BSE images showing examples of albite (Ab)-rimed Ca-plagioclase grains with carbonate capture (Cc); J-K) SEM/BSE images showing examples of micrograins of quartz (Qz) and albite (Ab) within the carbonate matrix (C) within the stromatolite.

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Figure 3.9 SEM/BSE images showing examples of the complex albite-Ca-plagioclase-quartz-calcite mineral assemblages within the Molly Gibson Mine section of the Deep Spring Formation. A-C) complex mineral assemblages with SEM/EDS data showing amounts of silicon (Si), sodium (Na), calcium (Ca), and aluminum (Al); D) SEM/BSE image showing mica grain with associated albite-quartz-calcite grain with micrographic textures; E) SEM/BSE image showing mica grain and associated quartz and albite grains; F) SEM/BSE image showing small quartz grain with internal Ca-plagioclase (light gray) and nearby albite.

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Figure 3.10 SEM/BSE images of micrographic-like textures and associated elemental maps with phase interpretation from the Mount Dunfee section of the Deep Spring Formation. A-B) albite-calcite micrographic-like textures within the stromatolite laminae of the sample found in float (Figure 3.1C, 3.6), C) albite-quartz-calcite micrographic-like textures from interstromatolite zone of the sample found in float (Figure 3.1C, 3.6); D-G) albite-quartz-calcite micrographic-like textures within the stromatolite laminae of the stromatolite from stromatolite horizon A).

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Figure 3.11 Elemental maps of the three stromatolite samples collected from the Mount Dunfee section of the Deep Spring Formation. A-B) sample found in float (Figure 3.1C, 3.6); C-D) sample collected from stromatolite horizon A; sample collected from stromatolite horizon D.

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Figure 3.12 Cathodoluminescence images from the in-float stromatolite sample from the Mount Dunfee section of the Deep Spring Formation. A) cathodoluminescence image of an albite (Ab)-rimed orthoclase grain (SEM/BSE image in “D”) from the stromatolite laminae; B) SEM/BSE image of mineral assemblage within the interstromatolite zone; C) cathodoluminescence image of the section seen in “C”; D) SEM/BSE image of the grain seen in “A”.

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Figure 3.13 SEM/BSE images showing examples of the K-rich aluminosilicate from the in-float stromatolite sample (Figure 3.1C, 3.6) from the Mount Dunfee section of the Deep Spring Formation. A) linear concentration of the K-rich aluminosilicate within the stromatolite laminae; B) K-rich aluminosilicate within the interstromatolite zone; C) K-rich aluminosilicate crystals (red arrows) captured within albite of the stromatolite laminae.

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Figure 3.14 Stromatolite-authigenesis model illustrating geomicrobiological processes leading to the precipitation of syndepositional authigenic albite and K-rich aluminosilicates.

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CHAPTER 4: THE EARLY STAGES OF STROMATOLITIC GROWTH OF ROCK VARNISH

ABSTRACT This is primarily a descriptive study in which I examined the early stages of rock varnish development on surficial sediments from a dry stream channel in southern Nevada. While extremely rare, I identified a few grains which showed incipient rock varnish development. Here termed “microdots”, incipient rock varnish appeared as discrete spots arranged in a mottled pattern, with each microdot measuring 15-40 μm in diameter and only a few microns in thickness. Microdots are composed of either iron-oxide (or oxyhydroxide) or a mixed-composition of iron- and manganese-oxides (or oxyhydroxides). They exhibit morphological varieties which seem to correlate to their particular chemical composition. SEM observation shows that iron-rich microdots display nodular morphologies while manganese-rich microdots display smooth, wavy morphologies. More advanced rock varnish growth results in Fe-rich botryoidal morphologies on a smooth,

Mn-rich varnished surface, with textures that are similar to previously reported stromatolitic varnish. As a working hypothesis, I suggest that the mottled texture of Fe microdots reflects the distribution of microbial colonies on the grain surface and that the single-composition nature of the Fe microdots results from the periodic stability of iron- oxidizing microbes on the grain surface. Further, I suggest that mixed-composition microdots form during microbial Mn- and Fe-oxidation. This research lays the groundwork for further exploration of the early stages of rock varnish growth.

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INTRODUCTION The thin manganese- and iron-rich rock coatings so often found in dry climates attracted the attention of two polymaths of modern science. Both Alexander Von Humboldt and Charles Darwin found examples of rock varnish during their worldly studies and commented on its characteristics and formation. During his voyage on H.M.S Beagle, Darwin noted “rich brown” rock coatings which “seems to be composed of ferruginous matter alone” (Darwin, 1897). Darwin’s wrote that “the origin, however, of these coatings of metallic oxides, which seem as if cemented to the rocks, is not understood” (Darwin, 1897). Unbeknownst to Darwin, Von Humboldt had seemingly solved the rock varnish mystery nearly a century earlier (Von Humboldt, 1812) while studying the “brownish black crust” (Von Humboldt, 1812), found along the Orinoco River in Venezuela. In developing his hypothesis for the formation of the Oronoco rock varnish, Von Humboldt stated that “we must then recur to the idea of a chemical solution,” after noting the lack of “grains of sand, nor spangles of mica, mixed with the oxides”—an interpretation that holds up well, after decades of modern study.

Since Von Humboldt’s pioneering study of rock varnish, two centuries worth of detailed geochemical observations, using advanced microanalytical technology, have broadened our understanding of rock coatings significantly (Dorn et al. 2012; Dorn, 2009; Dorn & Oberlander, 1981; Engel & Sharp, 1958; Krauskopf, 1957; Esposito et al., 2019; Perry & Kolb, 2004a; Lu et al., 2019; Perry, Dodsworth, Staley, & Engel, 2004; Northup et al., 2010, Xu et al., 2019).

Beyond the early observations of manganese and iron-rich rock coatings, a wide array of compositions have been found to coat rocks in a variety of environmental conditions and attributed biotic contribution (Dorn et al. 2012). Multiple attempts have been made to develop a rock varnish dating technique with varying success and potential future possibilities (Watchman, 2000).

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Rock varnish samples on sediment from the Techatticup Mine and Mill in Nelson, Nevada, were recently shown to play a role in trace element scavenging in arid environments (Sims et al., 2017). The study also showed that the different grain sizes of rock particles contained differing amounts of various trace elements, when analyzed using USEPA Solid Waste 846 protocols. Arsenic was detected in higher amounts within the varnish of silt sized fractions, for example, while Se was found only on coarse fractions (Sims et al., 2017). The trace elements of Cr, Cu, Fe, Mn, Pb, and Zn all show increasing concentrations within the desert varnish with increasing grain size. Similarly, regional studies of rock varnish contamination near coal-fired power plants showed that elements commonly associated with these processes were found highly concentrated in varnished surfaces and not in un-varnished surfaces (Nowinski, Hodge, & Gerstenberger, 2012) with a particular emphasis on levels of mercury (Nowinski, Hodge, Gerstenberger, & Cizdziel, 2013). The sediment grains used for this study were selected based on characteristics necessary for previous studies (Sims et al., 2017) as well as a currently unpublished study by the same author which utilized the suite of microscopes at the Electron

Microanalysis and Imaging lab at UNLV. The intense analysis of these sediment samples under the SEM provided an ideal opportunity to study incipient rock varnish.

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Geological Setting The Techatticup mine, located in the mining town Nelson, Nevada, is situated within the Basin and Range geologic province of the Western United States. This geological setting consists of parallel mountain ranges and valleys, which formed during the region’s tectonic extension and crustal thinning. The stratigraphy of the field area primarily consists of Precambrian basement rock composed of schist and gneiss, overlain by Miocene volcanics (Anderson, 1971; Hansen 1962; Longwells et al. 1965). These volcanics belong to the Pasty Mine assemblage. This assemblage consists of andesite flows which are 15 to 18 million years old, and rhyolitic and andesitic ash-flow tuff. The youngest volcanic rocks in the area the Mount Davis assemblage of rhyolites, basalts, and andesites which are 12 to 15 million years old (Anderson, 1971; Anderson et al., 1972; Darvall, 1991; Morikawa, 1993). Overlying these Miocene volcanics are basins filled with ephemeral wash sediments consisting of poorly sorted, eroded bedrock which is eroded, transported, and deposited during large storm events.

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Biogenic and Abiogenic Contributions in Varnish Formation Many studies of rock varnish have been conducted, resulting in a variety of interpretations regarding its formation. Both biotic and abiotic mechanisms have been invoked, depending on the study and the environment in which the varnish formed. While many studies have identified endolithic microbial colonies in association with varnished surfaces, valid abiotic interpretations have also been proposed. Like most things in science, the process behind varnish formation is likely spatially and temporally dependent on environmental conditions and many different and independent processes are at play. Additionally, the variability of surficial products of rock decay and varnish formation, including Mg-rich coatings, fungal mats, iron films, and silica glazes, hints at the complex processes behind their formation (Dorn & Jeong, 2018).

The Abiogenic Model

Previous researchers have detailed abiogenic models to explain rock varnish formation, focusing on the changes in element solubility during changes in the pH and Eh of dust particles on the rock surface (Engel & Sharp, 1958; Krauskopf, 1957). In a purely abiogenic model, Mn (II) and other enriched elements within rock varnish are interpreted to have been leached from dust particles during wetter periods in which the pH and Eh are lower than background conditions. According to this model, these elements are precipitated onto the rock surface during and after evaporation when the pH and Eh have both risen to background conditions. Xu et al.

(2019) introduced the novel addition of photooxidative mechanisms into the abiogenic model, in which solar rays in the 401-683 nm range removes electrons from the valence band to the conductive band. This helps explain the abiotic oxidation of Mn (II) to produce insoluble species, which leads to the formation of birnessite (MnO2) (Xu et al., 2019). Additionally, the presence of Mn oxides catalyzes the oxidation of Mn (II) to form more Mn oxides (Mn oxides begets Mn oxides, suggesting that a limiting factor in rock varnish formation may be the initial formation of

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Mn oxides). These lines of evidence suggest that abiogenic photooxidative processes may play an important role, alongside other abiogenic and biogenic mechanisms in the formation of rock varnish (Xu et al., 2019). Similarly, Lu et al. (2019) demonstrated the significant phototrophic-like behavior of Fe and Mn (oxyhydr)oxides in mineral coatings (Lu et al., 2019). They used a suite of scanning electron microscopic techniques, micro-Raman spectroscopy, X-ray absorption spectroscopy, and a newly developed photoelectric detection device to show that these coatings “behave as highly sensitive and stable photoelectric systems” (Lu et al., 2019, p. 9741). This mineral semiconductor world of Lu et al. (2019) provides additional capability for the Earth’s surficial environmental pathways, in addition to oxygenic photosynthetic pathways, to absorb incoming solar radiation and transform its energy during redox cycling. Both the photooxidation models of Lu et al. (2019) and Xu et al. (2019) utilize the solubility of Mn in dust particles during the instability of pH and Eh of the rock surface during wet and dry periods to concentrate the necessary ions for rock varnish formation—emphasizing the potentially fully abiogenic formation of the rock varnish.

Biogenic Models

Several researchers have proposed key biotic roles in the formation of rock varnish (Dorn & Oberlander, 1981; Esposito et al., 2019; Perry & Kolb, 2004a; Perry, Dodsworth, Staley, &

Engel, 2004; Northup et al., 2010). These biotic models all rely on the Mn-concentration abilities of Mn-oxidizing microbes within microsedimentary basins on the grain surfaces. Using SEM imaging, Mn-concentrating Metallogenium-like microbes are often found inhabiting varnished surfaces which, when cultured, produce manganese films and precipitates (Dorn & Oberlander, 1981; Northup et al., 2010).

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Varnished and non-varnished samples from Black Canyon, New Mexico, were collected and the microbial communities were analyzed to determine the community characteristics associated with varnish surfaces (Northup et al., 2010). Microbial communities were dominated by Chloroflexi, ktedobacteria, or cyanobacteria on the varnish surfaces, while non-varnish- coated surfaces were dominated by actinobacteria and cyanobacteria (Northup et al., 2010). The researchers utilized genetic techniques to characterize the microbiota in relation to rock varnish. Sixty-five percent of microbial cultures collected from these varnish surfaces produced manganese precipitates (Northup et al., 2010). Incipient varnish was found on seemingly non- varnished surfaces in the Northup et al. (2010) study. This early stage incipient varnish was often found near microbial colonies inhabiting microbasins on the rock surface, explaining the similarities in microbial communities between varnished and non-varnished surfaces (Northup et al., 2010). No data were presented on the morphology or composition of the incipient varnish by Northup et al. (2010). Additionally, amino acids were detected within varnish, supporting the biogenic nature of varnish origins (Perry et al., 2003).

Previous research has shown the powerful role that Mn-oxidizing microbes play in the Mn geochemical cycle (Nealson, 1983; Maki, Tebo, Palmer, Nealson, & Stanley, 1987) and their ability to produce Mn precipitates during their metabolic processes (Chafetz, Akdim, Julia, & Reid, 1998; Ehrlich, 2000; Ferris, Fyfe, & Beveridge, 1987; Maki et al., 1987; Stiles, Mora, & Driese, 2001; Spilde et al., 2005). These previous studies, as well as many others for a wide variety of varnish compositions, have provided direct evidence of microbe-mineral interactions in relation to mineral deposition within marine and lacustrine waters, soils, and even caves.

Additionally, laboratory culture-based experimentation has provided abundant evidence of the microbe-mineral interactions during varnish formation (Dorn & Oberlander, 1981; Jones, 1991; Perry et al. 2003). Jones (1991) detailed an unusual rock varnish from a hyper arid region in Peru which was Fe-rich and Mn-poor on the surface of the rock and Fe-rich, Mn-rich within surficial

139 pits. Jones (1991) found the highest concentrations of bacteria within the pits containing Mn-rich varnish and found that fogs or dews could provide the necessary Mn to the microbial communities to allow for Mn precipitation during varnish formation.

Polygenetic Models

Models detailing the origins of varnish often involve both abiogenic and biogenic processes (Dorn & Oberlander, 1981; Dorn, 2007; Esposito et al., 2019; Jones, 1991). The vast environmental, compositional, and morphological differences between varnish sites requires fluidity in interpretation. There seems to be a grab bag of biotic and abiotic processes which work in tandem to produce varnish, each with substitutes depending on environmental and biological conditions at the time. Polygenetic models include all previously proposed processes likely responsible for the precipitation of varnish—deposition of dust into microsedimentary basins on rock surfaces, microbial concentration of manganese, abiotic photooxidative concentration of manganese, silica dissolution and gel formation, etc.

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METHODS My early SEM observations of the grains from the Techatticup Mine and Mill resulted in a clear obfuscation of the early rock varnish signal due to the abundance of iron in the underlying grain. As such, I began this study by observing the grains under a compound microscope, collecting the lightest colored grains with the assumption that the lighter the grain the less iron (as well as manganese) it would contain (Fig. 4.1A). These low-iron grains allowed for easier identification of rock varnish growth on their surface through SEM/EDS element mapping and point scans. Once desert varnish was identified on the grains surface, analysis of the varnish would include imaging with the suite of Scanning Electron Microscopes of the Electron Microanalysis and Imaging Lab at the University of Nevada, Las Vegs, including the JEOL JSM-5600, the JEOL JSM-6700F Field Emission Scanning Electron Microscope, and EDS point scanning and elemental mapping of the varnish (Fig. 4.1).

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RESULTS Due to the time required to properly prepare grain samples for SEM observation in the hunt for rock varnish, and the seemingly rare nature of rock varnish growth on the surficial sediments, the vast majority of grains observed in the SEM showed no signs of rock varnish formation. The few examples of incipient rock varnish found in the Techatticup Mine and Mill wash sediments are concentrated on a few grains (Fig. 4.1). Herein, we show detailed chemical and morphological analysis of the incipient rock varnish.

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Early Stages of Rock Varnish Formation

Morphology

The early stages of rock varnish on the Techatticup Mine wash sediments come in the form of small spherical, mineralized dots, measuring 15-40 μm in diameter and only a few microns in thickness (Fig. 4.1-4.3). I termed these mineral growths “microdots”. Microdot morphological variation within the early stages of rock varnish formation on the Techatticup

Mine sediments come in two varieties: (1) nodular (Fig. 4.2) and (2) smooth (Fig. 4.4). Nodular varieties display a wide range in mineralization, from smooth mineralization with little to no nodules (Fig. 4.3A) to predominantly nodules with smooth mineralization confined to the edges (Fig. 4.3E). Microdot mineralization within this gradient is intermixed on grains that have been covered by microdots (Fig. 4.1A-B, E) with no clear pattern. Endmembers of this gradation can be found situated close enough to be touching, where a non-nodular microdot is seen being covered by a nodular microdot that is mostly covered by nodules (Fig. 4.3F). Several examples of microdots which have been chipped away were found (Fig. 4.1C-E, Fig. 4.2F,G-H), which shows them to be extremely thin coverings on the surface of the rocks, even when extensive nodule formation has occurred (Fig. 4.2H). When microdots have been covered by nodules (Fig. 4.3E-F), the smooth mineralization on the outer edges of the microdots seems to be more thickly mineralized than their non-nodular neighbors (compare the lightly mineralized microdot on the left with the heavily mineralized, nodular microdot on the right in Fig. 4.3F). This suggests that the smooth mineralization of the microdots does not stop when nodular formation begins.

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Composition

Two distinct compositional variants exist throughout the Techatticup microdots—Fe-rich and mixed-composition. The Fe-rich microdots are found in highly concentrated groups mottling the surface of individual grains (Fig. 4.1A-B). In comparison, mixed-composition microdots (Fig. 4.4) are less common and are not nearly as ubiquitous when found. Fe-rich microdots show low variability in EDS quantitative analysis with typical composition of 77 wt% Fe, 23 wt% O, and 0 wt% Mn after elements known to exist in the underlying grain (Si, Al, etc.) were removed from the EDS analysis (Fig. 4.5). The microdots on the grain shown in Fig. 4.1 contain an average of 2.5 wt% vanadium. Comparison of the varnish covering with the underlying grain shows a drastic difference in chemistry (Fig. 4.5). The small amount of iron seen within the underlying grain is likely due to the electrons and X-rays emanating from within the chipped region interacting with the iron of the varnish on the rim. Mixed composition microdots (Fig. 4.4) are less common on the Techatticup Mine wash sediments, but typically contain higher amounts of impurities. They are composed of a large variety of Fe-Mn mixes of up to 27 wt% Mn (3.4 wt% Fe) and as little as 2.3 wt% Mn (13.9 wt% Fe).

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Advanced Stage Varnish While even more rare than the early stages of desert varnish on the Techatticup Mine wash sediments, a few examples of desert varnish in advanced stages of development were found (Fig. 4.6). Below, I detail the morphological and compositional characteristics of the advanced stage desert varnish.

Morphology

Morphological variation amongst the advanced stages of rock varnish is much higher than that of the early stage microdots. No particular patterns were found within the examples; however, some very interesting morphological variants are detailed herein. Botryoidal morphologies were found to be surrounded by smooth surficial varnish covering (Fig. 4.6B). Each botryoid measuring roughly 10 microns in diameter and are often connected in long chains of five or more botryoids (Fig. 4.6B). Other smooth varnish coverings were found which did not contain botryoidal growths (Fig. 4.6A). Morphological variants typically had specific compositions which resulted in their particular morphology, with some variation and overlap in between.

Composition

Large compositional variation exists within small areas of advanced stage desert varnish on the Techatticup Mine wash sediments. The compositional difference between the botryoidal growths and the surrounding smooth varnish can as be high as a 30 wt% difference in Mn and roughly 40 wt% difference in Fe. Smooth varnish contains more Fe and less Mn than the botryoidal overgrowths, which contains comparatively more Mn and less Fe (Fig. 4.6C). Changes in accelerating voltage can create large variation in the resulting EDS spectrum of the advanced stage varnish. Accelerating voltage of 15 kV during SEM/EDS analysis of the

145 varnish results in higher amounts of Mn within the EDS spectrum (Fig. 4.6C), while a lower accelerating voltage of 8 kV resulted in reduction in the amount of Mn within the varnish (Fig. 4.6C). Within the botryoids, a reduction of 16 wt% Mn (37% of total Mn at 15 kV) occurred when dropping the accelerating voltage from 15 kV to 8 kV (Fig. 4.6C). Similarly, a complete removal of Mn from the smooth varnish EDS spectrum occurs at with the same drop in accelerating voltage. (Fig. 4.6C).

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DISCUSSION Incipient Desert Varnish Rock varnish has been assigned a variety of compositions throughout the last half century of detailed studies (Dorn & Krinsley, 2011; Nagy et al., 1991; ). However, three basic components are always present within the varnish laminae: iron oxyhydroxides, manganese oxyhydroxides, and clays. These mineral constituents occur within the thin varnish covering in different concentrations throughout the surface of the rock and the depth of the varnish (Krinsley et al., 2019). Previous authors have discussed the appearance of incipient varnish on rock surfaces which contain no macroscopic signs of varnish covering (Northup et al., 2010). However, details on incipient varnish have not been given, other than to note that they are typically found near thriving microbial colonies on the rock surface (Northup et al., 2010). In this chapter, I have detailed the morphological (Fig. 4.1-4.6) and compositional (Fig.4.5-4.6) characteristics of incipient desert varnish on wash sediments from a hyper-arid region of Nevada’s desert. While I can definitively show the surficial nature of this incipient varnish (Fig. 4.1D, 4.2F,G,H, Fig. 4.5A), I found no direct evidence of microbial colonies inhabiting the sediment grains. This absence of microbial life is likely due to the aggressive washing of the sediment with DI water during previous experimentation. It is likely that the unwashed sediments were inhabited with microbial life.

The Fe-rich microdots found on the grain in Figure 4.1 show a gradational mineralization from light mineralization with little to no nodules (Fig. 4.3A) to heavy mineralization covered with nodules (Fig. 4.3E). Gradational mineralization can be seen between these two extremes (Fig. 4.3B-D). Proximity to other microdots seems to have no effect on mineralization grade, considering the least mineralized microdot I found was directly touching one of the most mineralized microdots (Fig. 4.3F). Similarly, the diameter of the microdot seems to have no

147 correlation on mineralization grade, since all microdots are within a small range of diameters. Variation can be quite high even on a single microdot, where areas with significant nodule mineralization exist next to smooth, lightly mineralized areas without any nodules (Fig. 4.3D). I interpret these patterns as evidence for the biogenic origin of microdots on the Techatticup Mine wash sediments, resulting from the microbial oxidation of Fe on the surface of the grain by discrete groups of iron oxidizing microbes. While this is not a new interpretation, considering the incipient varnish found surrounding “colony-occupied pits” on rock samples from Black Canyon,

New Mexico (Northup et al. 2010), I do detail the morphological and compositional aspects of the incipient varnish for the first time.

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Unusual Compositional Aspects

Incipient Varnish

The absence of Mn within some incipient varnish on the Techatticup Mine wash sediments illustrates the unusual nature of this locality. The small size of the wash sediments that are the focus of this study (no more than 2 mm for the mottled grain in Fig. 4.1) suggest a high level of instability on the desert surface. While we typically think of desert varnish forming on stable rock faces within consistently dry/hot environments with periods of wetter climate, these incipient varnishes are nucleating on grain surfaces which likely experience frequent movement by wind and, less frequently, water. All previous examples of incipient varnish have been found on stable rock faces which experience no movement over geologically significant periods of time. This grain instability likely resulted in less dust accumulation within the microsedimentary basins on the grain surface, effectively cutting off the supply of Mn to the microbial community. However, this should simultaneously cut off the supply of Fe to the microbial community.

The discrepancy in Fe/Mn levels in some incipient varnish may result from the transitory nature of alkaline conditions at the sediment surface (Dorn, 1990), which effectively reduces the ability for the microbial community to precipitate Mn (Jones, 1991). While Fe/Mn ratios typically fluctuate drastically from sample to sample and layer to layer, some Mn has always been observed in association with Fe-rich rock varnishes (Jones, 1991). Therefore, the incipient varnish found on the Techatticup Mine wash sediments (Fig. 4.1) may represent the earliest stage of varnish development ever recorded, in which mineralization has taken place either (1) during a period of high alkalinity and no Mn precipitation, or (2) during a period of habitable stability for iron oxidizing microbes and conditions that are not habitable for Mn oxidizing microbes. However, the catalytic nature of iron oxyhydroxides typically can result in the oxidation and precipitation of Mn (Jones, 1989), and something seems to have been inhibiting this process.

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Further development of the varnish on these grains would allow for periods of habitable stability for both microbial communities to thrive, resulting in precipitation of layers of both Fe-rich varnish and Mn-rich varnish as seen in mixed-composition microdots (Fig. 4.4) and advanced stages of varnish (Fig. 4.6). Therefore, while the Fe microdots lacking any detectable Mn are highly unusual, the incipient nature of the microdots is likely to produce unexpected data and allow for new processes into our interpretation of rock varnish origins. A few examples of mixed-composition microdots were found on the Techatticup Mine was sediments (Fig. 4.4). These mixed-composition microdots are composed of variable Mn/Fe ratios and display smooth textures with a complete absence of nodules (Fig. 4.4), and they always include barium or titanium. The composition is more in line with what is typically expected for desert varnish, as opposed to the Fe microdots which lack Mn (Fig. 4.1-4.3), and probably originate from colonies of microbial Mn- and Fe-oxidizers on the grain surface.

Advanced Stage Varnish

Some advanced stage varnish found on the Techatticup Mine wash sediments displays highly variable compositions within very small areas (Fig. 4.6B) while some areas show spatially consistent composition (Fig. 4.6A). Those varnish surfaces which show consistent composition display a smooth, uniform morphology (Fig. 4.6A), while varnish surfaces display wide variation in composition show several morphological variants. Botryoidal “stromatolitic” varnish was found surrounded by very smooth varnish surfaces, each of which has a different compositional Fe/Mn ratio, and the measured ratios which are very sensitive to SEM conditions.

Variation in accelerating voltage during SEM/EDS analysis changes the area of electron- sample interactions occurring within the sample. Higher accelerating voltages result in more energy for the electrons to “push” their way into the sample, producing a larger interaction area in which the electrons can create the characteristic X-rays we use to determine composition

150 during SEM/EDS analysis. Comparatively, lower accelerating voltages produce smaller electron- sample interaction areas. During SEM/EDS analysis of the botryoidal desert varnish of the Techatticup Mine wash sediments, spot analyses were conducted with a high (15 kV) and low (8 kV) accelerating voltage to determine if a change in varnish composition occurs at depth (Fig. 4.6C). Results showed that higher accelerating voltage resulted in higher Mn/Fe ratios, and lower accelerating voltage resulted in lower Mn/Fe ratios (Fig. 4.6B-C). Lower accelerating voltage for the smooth varnish (Fig. 4.6B-C, point 3) resulted in a complete loss of Mn in the EDS quantitative analysis, a drop of 7 wt% for Mn and a 7 wt% increase for Fe. Similarly, lowering the accelerating voltage on the botryoidal spot analysis resulted in decrease of 15.9 wt% of Mn and an increase Fe in the same amount. These results show that the upper surface of both the botryoidal and smooth morphologies on this grain are either completely depleted or highly reduced in Mn. This Mn-poor layer may be anywhere from 2-5 μm thick based on these data. This is further evidence, along with the formation of incipient Fe microdots which lack any Mn, of the periodic nature of Fe and Mn precipitates during varnish formation. These results are not consistent with previously reported varnish which always contain both Fe and Mn in varying ratios.

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Clay Portion of Incipient Varnish The lack of detailed information about the clay component, which likely exists in the incipient varnish, is twofold: (1) The extreme rarity of the incipient varnish on the grains make thin section analysis impossible. Only one sample was found to have enough covering of incipient varnish to warrant attempted thin sectioning (Fig. 4.1). (2) The similar chemistry of the underlying grain with the likely clay portion, together with the very thin coating of the incipient varnish, makes SEM/EDS analysis impossible due to the previously discussed electron-sample interaction area including several microns below the surface of the varnish covering. Any SEM/EDS analyses of the incipient varnish mostly includes the underlying grain chemistry, making any compositional determination of the varnish silicates impossible.

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Stromatolitic Growth Many previous studies have discussed the remarkable similarity in growth habit and growth processes between rock varnish and stromatolites (Krumbein & Geile, 1979; Krumbein, 1983; Nagy et al., 1991). These similarities include layered, columnar growth, the capture of detrital grains on and between columnar mounds (similar to the detrital sediments found within stromatolite laminae and within the interstromatolite zone in chapter 3 of this dissertation), in situ mineral precipitation due to microbial activity, and complex microbial communities.

However, the struggle to accurately define stromatolites over the recent decades has created much confusion surrounding the term, with the current consensus emphasizing the benthic nature of stromatolites (Kuznetsov, 2015). Under this current consensus, rock varnish would not be considered “stromatolites,” however the similarities in origin and form are numerous. The similarities between proper stromatolites and the observed growth habits of some rock varnish has led researchers to consider rock varnish to be a type of stromatolitic growth, even though they occur under vastly different environmental conditions. Nagy et al. (1991) coined the term

“microstromatolites” for the small nature of the columnar stromatolitic growth habit of the rock varnish they imaged from Twin Peak Mountain Park, near Phoenix, Arizona. Other researchers had previously utilized the term stromatolite to describe various microbially produced rock varnishes and concretions (Krumbein & Geile, 1979; Krumbein, 1983). However, with the current consensus on the definition of stromatolites, this is no longer applicable.

The mottled texture of incipient varnish on the Techatticup Mine wash sediments further illustrates the similarities between stromatolites and rock varnish. The early formation of stromatolites likely begins with such a mottled pattern of mound development, considering the modern distribution of stromatolites at the Highborne Cay, Bahamas (Dupraz et al., 2009; Reid et al., 2000). If the incipient mottled textures seen in Figure 4.1B were to continue development,

153 they would begin to close the distance between microdots and begin forming the columnar stromatolitic varnish textures seen in Nagy et al. (1991).

154

CONCLUSIONS The lack of manganese detected in some of the varnish microdots (~30 um diameter), and its presence in other varnish, suggests that the deposition of the iron- and manganese-oxides do not occur simultaneously, but occur at intervals which are likely biologically dependent which are themselves environmentally dependent. The mottled texture of incipient varnish on the surface of sediment grains correlates to the expected distribution of microbial colonies. Even though no microbial life was detected in SEM analysis of the sediment grains (likely a result of the sediment collection and preparation), previous studies have shown an intimate relationship between microbes and rock varnish. The lack of detectable clay minerals within the incipient varnish is likely the result of mixed signals from the underlying grain in SEM/EDS analysis. Further investigation is needed to detail the clay inclusion within the incipient varnish. While no concrete evidence of microbial involvement in incipient varnish development was found, I concur with previous authors who stress the importance of microbially-mediated Fe- and Mn- oxide and oxyhydroxide precipitation during the formation of rock varnish.

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Figure 4.1 Sample preparation and SEM imaging. A) aluminum billet holding iron-poor grains; B) SEM/BSE image of the grain seen in the square area in “A” showing mottled pattern of incipient varnish; C) SEM/BSE image of the square area in “D” showing the details of a chipped surface of incipient varnish; D) SEM/BSE image of the square area in “E” showing the detail of the smooth and nodular mineralization of incipient varnish; E) SEM/BSE image of the incipient varnish seen in the square area in “B” showing details of the chipped surface.

156

Figure 4.2 Examples of Fe-rich (no Mn) incipient varnish. A) SEM/BSE image of incipient varnish microdot; B-D) Elemental maps showing Fe (B), Mn (C), and Ba (D) of the microdot in “A”; E) SEM/BSE image of overlaying microdots showing light mineralization (right microdot) and heavy mineralization (left microdot); F) SEM/BSE image of Fe-rich microdot showing large chip through its diameter; G) SEM/BSE image of overlap of several microdots; H) SEM/BSE image of high-resolution SEM/BSE image of microdot with surficial chip; I) SEM/BSE image of an Fe-rich microdot showing nodular mineralization and light mineralization.

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Figure 4.3 Incipient Fe-rich varnish gradational mineralization from light (A) to heavy nodular (E) mineralization. A) SEM/BSE image of light, smooth mineralization with chipped surface; B) SEM/BSE image of medium-light mineralization with some nodular formation; C) SEM/BSE image of medium mineralization with medium nodular formation; D) SEM/BSE image of light to medium mineralization with medium nodular formation; E) SEM/BSE image of heavy mineralization with heavy nodular formation; F) SEM/BSE image of two Fe-rich microdots showing nature juxtaposition of mineralization extremes (light to heavy, non-nodular to nodular).

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Figure 4.4 Mixed-composition (Mn/Fe) incipient varnish.

159

Figure 4.5 Chipped incipient varnish surface. A) SEM/BSE image of chipped incipient varnish showing the Fe-rich microdot (bright) and the underlying grain (medium to dark gray); B) SEM/EDS point scan qualitative analysis showing the difference in composition between the incipient varnish and the underlying grain.

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Figure 4.6 Examples of advanced stages of varnish development. A) SEM/BSE image of smooth, Mn-rich varnish and associated elemental maps (Si, Mn, Pb, and Fe); B) SEM/BSE image of smooth and botryoidal varnish; C) SEM/EDS point scan quantitative analysis of points in “A” and “B” showing the change in composition relative to accelerating voltage (15 kV vs 8 kV).

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CURRICULUM VITAE

[email protected]

EDUCATION Doctor of Philosophy in Geoscience Fall 2020 University of Nevada, Las Vegas Dissertation Focus: Precambrian-Cambrian Paleobiology and fossil preservation Bachelor of Science in Geology Fall 2014 Utah State University

Permits/Certifications Bureau of Land Management Paleontological Resources Use Permits • Utah UT20-004C, 2019 • Nevada N-099489, 2019 Utah State Lands Paleontological Permit • Utah 2020-574 OSHA • 40-Hour OSHA HAZWOPER, 2019 • 30-Hour OSHA Construction, 2019 Other • First Aid/CPR/AED, 2019 • ESRI Cartography Certification, 2020

RECENT WORK EXPERIENCE Paleontologist/Geologist, Environmental Scientist October 2019 to August 2020 BEC Environmental, Inc. • Paleontological Resources Field Surveys/Reports • Phase I Environmental Site Assessments • Environmental Reviews • Geologic/hydrologic descriptions Research Assistant in Electron Microanalysis and Imaging Laboratory (EMiL) Fall 2014 to 2018 Geoscience Department, University of Nevada, Las Vegas • Proficient operation of the Field Emission SEM (JSM-6700F), the Microprobe (JXA- 8900), and the general use SEM (JSM-5610), as well as sample preparation techniques (Gold sputter, carbon coater, polishing) • Train new users on SEM use and techniques

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• Conduct maintenance on Scanning Electron Microscopes and preparation equipment

TEACHING EXPERIENCE Teaching Assistant, Geography 104: Physical Geography Lab Fall 2019 Teaching Assistant, Geology 102: Earth and Life Through Time Summer 2018 Teaching Assistant, Geography 104: Physical Geography Lab Spring 2017 Teaching Assistant, Geology 5540/6540: Quantitative Geology Spring 2014 Teaching Assistant, Geology 1115: Physical Geology Lab Spring 2013

RESEARCH

Publications • J.D. Schiffbauer, T. Selly, S.M. Jacquet, R. Merz, L.L. Nelson, M.A. Strange, Y. Cai, and E.F. Smith, 2020: Discovery of bilaterian-type through-guts in cloudinomorphs from the terminal Ediacaran Period. Nature Communications. • T. Selly, J.D. Schiffbauer, S.M. Jacquet, E.F. Smith, L.L. Nelson, B.D. Anderson, J.W. Huntley, M.A. Strange, G.R. O’Neil, C.A. Thater, N. Bykova, M. Steiner, B. Yang, Y. Cai, 2020: A New Cloudinid Fossil Assemblage from the Terminal Ediacaran of Nevada, USA. Journal of Systematic Palaeontology. • M.A. Strange and S.M. Rowland, 2017: The Ediacaran-Cambrian biotic transition in eastern California and southwestern Nevada and its global context. Zzyzx Desert Research Symposium Conference Paper. • J. Luque, R.M. Feldman, O. Vernygora, C.E. Schweitzer, C.B. Cameron, K.A. Kerr, F.J. Vega, A. Duque, M.A. Strange, A.R. Palmer, and C. Jaramillo, 2019: Exceptional preservation of mid-Cretaceous marine arthropods and the evolution of novel forms via heterochrony. Scientific Advances. • E.F. Smith, L.L. Nelson, M.A. Strange, A.E. Eyster, S.M. Rowland, D.P. Schrag, F.A. Macdonald, 2016: The last of the Ediacaran Biota: two exceptionally preserved body fossil assemblages at Mt. Dunfee, Nevada. Geology.

Projects in Progress • Dissertation Chapter: Disentangling pyritization from direct iron oxyhydroxide and iron-rich clay precipitation: preservation of Ediacaran tubicolous fossils, Deep Spring Formation, Nevada. • Dissertation Chapter: Microbially mediated dissolution of detrital feldspar and its reprecipitation as authigenic albite—A potential microbialite dating technique • Dissertation Chapter: Early nucleation stages of desert varnish • Microfossil analysis across the Ediacaran-Cambrian boundary within the Deep Spring Formation, Nevada • A new technique for the interpretation of soft-tissue fossils and its application on hyolithids from the Cambrian of northern Utah

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• Sequence biostratigraphy of the Spence Shale, Langston Formation, northern Utah

HONORS, SCHOLARSHIPS, AND GRANTS University of Nevada, Las Vegas • (GRANT) NASA Astrobiology Early Career Collaboration Grant 2017 • (GRANT) Geological Society of America Student Research Scholarship 2015 • Bernada French Scholarship 2015, 2017, 2018, 2019 • Edwards & Olswang Geology Scholarship 2015, 2016, 2018, 2019

Utah State University • Lillywhite Presidential Scholarship • Geology Department Undergraduate Researcher of the Year 2013 • (GRANT) Undergraduate Research and Creative Opportunities grant • (GRANT) College of Science Mini-Grant • Peter T. Kolesar Scholarship • Clyde T. Hardy Scholarship

PROFESSIOINAL PRESENTATIONS • Geological Society of America Annual Meeting Spring 2018 o MICROBIALLY MEDIATED DISSOLUTION OF DETRITAL SANIDINE AND ITS REPRECIPITATION AS AUTHIGENIC ALBITE AND QUARTZ WITHIN EDIACARAN STROMATOLITES—UNIFYING MICROBIAL DISSOLUTION AND PRECIPITATION PROCESSES.

• Geological Society of America Annual Meeting Spring 2017 o Disentangling pyritization from direct iron oxyhydroxide and iron-rich clay precipitation using Precambrian fossils from the Deep Spring Formation, Nevada.

• Geosymposium- University of Nevada, Las Vegas Spring 2016 o Tracking Carbonate Biomineral Diagenesis in Modern-Pleistocene Mollusks from the Las Vegas Formation using electron microanalysis- Presented by undergraduate Assistant

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o Neoproterozoic fossils from Laurentian ad their preservation by direct precipitation of iron oxy-hydroxides and iron-rich aluminosilicates: a new mode of exceptional preservation

• Geology Society of America Annual Meeting Fall 2015 o First report of Conotubus in Western North America and its apparent preservation by direct precipitation of iron oxides and iron-rich aluminosilicates.

• University of Nevada, Las Vegas Geosymposium Spring 2015 o Paleontology and Taphonomy of the Ediacaran Mount Dunfee Lagerstätten, Nevada

• Advisory Board Meeting Posters- USU February 2012 o Paleo Studies of the Middle Cambrian Spence Shale of Northern Idaho

• Student Showcase- USU May 2012 o Paleo Studies of the Middle Cambrian Spence Shale of Northern Idaho

• Posters on the Hill, Salt Lake City- Utah State Capitol Building January 2013 Paleobiology of Some Middle Cambrian Animals from Northern Utah and a New Technique for the Interpretation of Soft-Tissue Fossils

• Utah Conference on Undergraduate Research (CUR)- USU February 2013 o A New Technique for the Interpretation of Soft-Tissue Fossils and its Application on Cambrian Hyolithids from the Spence Shale of Northern Utah (Oral)

OTHER WORK EXPERIENCE Curator of Geology, Utah State University Geology Museum May 2011 to Summer 2014 • Curate department collections including organization and preservation of specimens • Design the museum and displays • Tour guide • Managed interns GeoSymposium Coordinator 2018 & 2019 • Coordination of roughly 40 student volunteers • Develop program and run the symposium • Solicit sponsorships • Invite keynote speakers • Manage registration and budget Research Assistant in Core Lab May 2010 to Summer 2014

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Geology Department, Utah State University • Member of science crew at Mountain Home, Idaho drill site • Lithologic logging and fracture analysis of thousands of feet of core • XRD sample preparation and data collection Department Lab Tech May 2011 to Summer 2014 • Prepare and run samples for geochemical analysis • Prepare thin sections • Install/fix equipment • Field work • Sample collection Internship at the Museum of Natural History Box Elder County May 2010 to September 2010 • Digitize the collection • Design displays • Remake labels

TECHNICAL SKILLS • Proficient use of Scanning Electron Microscopes and sample preparation equipment • Heavy Liquid Separation using LMT • Microbiological experimentation o Culturing o Novel experimental setup development o Mineralogical analysis • Feldspar Staining • Microsoft Office software suite with skill in Excel data analysis • Quantitative analysis (parametric and nonparametric) o Past 3 o Excel • Adobe Photoshop and Illustrator • Photography • Beginners programming • PhreeqC

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