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2017 Investigation of Paleoredox Conditions across the Llandovery- Wenlock () Boundary: Implications for the Event and Carbon Isotope Excursion Andrew Thomas Kleinberg

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COLLEGE OF ARTS AND SCIENCES

INVESTIGATION OF PALEOREDOX CONDITIONS ACROSS THE LLANDOVERY-

WENLOCK (SILURIAN) BOUNDARY: IMPLICATIONS FOR THE IREVIKEN

EXTINCTION EVENT AND CARBON ISOTOPE EXCURSION

By

ANDREW KLEINBERG

A Thesis submitted to the Department of Earth, Ocean, and Atmospheric Science in partial fulfillment of the requirements for the degree of Master in Science

2017 Andrew Kleinberg defended this thesis on June 29, 2017. The members of the supervisory committee were:

Seth Young Professor Directing Thesis

Jeremy Owens Committee Member

Angie Knapp Committee Member

The Graduate School has verified and approved the above-named committee members, and certifies that this thesis has been approved in accordance with university requirements.

ii ACKNOWLEDGMENTS

I would like to take the time to thank my thesis advisor, Dr. Seth Young, whose invaluable guidance, patience, and positive encouragement has pushed me well past the precipice of my own capabilities and expectations throughout my Masters career. I am so grateful for the opportunities he has provided me and being given the chance to study my Masters under him. I would also like to take the chance to thank my other committee members: Dr. Jeremy Owens for his equal involvement during my Masters and his light-hearted and humorous, yet engaging and helpful, mentoring approach; and Dr. Angela Knapp for her helpful guidance and providing an oceanographic perspective to this geologic event. Additionally, I could have never of done this without the help of my friends/colleagues Chelsie Bowman and Nevin Kozik whose support and helpfulness throughout my Masters was immeasurably important. I would also like to thank Emily Benayoun, Lance Newman, Adolfo Calero, Claudia Richbourg, and Matt Mcdowell for their help as undergrad assistants during this process. I would also like to thank Burt Wolff of the National High Magnetic Field Laboratory Geochemistry department for his help in running my samples for carbon isotope analysis and Ben Underwood from the Indian State University Stable Isotope Research Laboratory for running my samples for sulfur isotope analysis on their instruments Additionally, I would like to thank the National Geologic Society of America, awarding me with the Charles A. and June R.P. Ross Research Grant, and Southeastern GSA for the partial funding of this project. Finally, I would like to thank my friends and family whose love, support, and encouragement helped me get through some of the harder parts of my Masters.

iii TABLE OF CONTENTS

List of Tables ...... vi List of Figures ...... vii Abstract ...... viii

1. INTRODUCTION ...... 1

2. GEOLOGIC BACKGROUND ...... 6 2.1 Early Silurian Climate and Oceans ...... 6 2.2 Depositional Environment, Facies, and Sequence Stratigraphy ...... 9

3. METHODS ...... 11 3.1 Sample Preparation and Extraction ...... 11 3.2 Stable Isotope and I/(Ca+Mg) ...... 12

4. RESULTS ...... 15 4.1 Newsom Roadcut, Nashville, Tennessee ...... 15 4.2 Pete Hanson Creek II, Roberts Mountains, Nevada ...... 16

5. DISCUSSION ...... 18 5.1 Evaluation of Diagenesis on Primary Geochemical Signatures ...... 19 34 5.2 Mechanism for late Llandovery-early Wenlock δ SCAS Trends: a Global Redox Proxy 21 5.3 Variation in I/(Ca+Mg) records: a Local Redox Proxy ...... 24 13 13 5.4 Factors Affecting Variability of δ CCarb and δ COrg ...... 25 5.5 Redox and Environmental Conditions Across the Llandovery-Wenlock Boundary ...... 26

6. CONCLUSIONS ...... 29

APPENDICES ...... 31

A. DATA TABLES ...... 31 B. FIGURES AND GRAPHS ...... 37

References ...... 46

Biographical Sketch ...... 55

iv LIST OF TABLES

A.1 Geochemical proxy data of Newsom Roadcut field site ...... 31

A.2 Geochemical proxy data of Pete Hanson II field site ...... 35

v LIST OF FIGURES

13 B.1 Generalized Silurian (Wenlock-Ludlow) δ Ccarb curve with and graptolite biostratigraphy, and new high precision U-Pb radiometric dates [Cramer et al., 2011, 2012, 2015] ...... 37

B.2 Timeframe of the Ireviken Carbon Isotope Excursion in relation to Jeppsson (1990) oceanic states, epicontinental stratigraphy, glaciation, and eustatic sea level change [Cramer and Saltzman, 2007]. Detailed biostratigraphy after Calner [2004] and eustatic sea-level curve from Cramer and Saltzman [2005] ...... 38

B.3 Primo and Secundo state model modified after Jeppsson (1990) and Cramer and Saltzman (2005). Red star marks the Pete Hanson Creek II section and blue marks the Newsom Roadcut section. WSBW stands for Warm Saline Bottom Waters ...... 39

B.4 Geochemical crossplots for evaluation of diagenesis for this study, please note that all 13 18 Newsom Roadcut δ Ccarb and δ Ocarb values are replotted from Cramer and Saltzman (2005, 13 18 13 34 2007). A) Plot of δ Ccarb vs. δ Ocarb. B) Plot of δ COrg vs %TOC. C) Plot of δ SCAS vs 18 18 δ Ocarb. D) Plot of CAS(ppm)% vs δ Ocarb ...... 40

B.5 Paleogeographic reconstruction of Laurentia during the middle Silurian ~431 Ma with field sites marked in their respective localities both during this time and in modern day ...... 41

B.6 Roberts Mountains Stratigraphic Column with stable carbon and sulfur isotopes and I/Ca+Mg data. The plotted is from Klapper and Murphy (1975) ...42

B.7 Newsom Roadcut Stratigraphic Column with stable carbon and sulfur isotopes and I/Ca+Mg data. Please note that δ13Ccarb and δ13Corg are replotted from Cramer and Saltzman (2007) and conodont biostratigraphy is from Barrick et al. [1983] ...... 43

13 B.8 δ Ccarb isotope data of Pete Hanson Creek II, Roberts Mountains, NV, Banwy River, Wales, 13 13 δ Corg data from Loydell and Frýda (2007), and the Newsom Roadcut, Nashville, TN, δ Corg data from Cramer and Saltzman (2007). Condont and graptolite biostraigraphy done by [Männik and Aldridge, 1989; Jeppsson, 1997; Loydell et al., 1998; Männik, 2007; Rubel et al., 2007; Cramer and Saltzman., 2006a; 2010]. Grey is due to a confusion of graptolite and conodont correlation ...... 44

B.9 Carbonate Redox Proxies and expected trends during the Ireviken CIE ...... 45

vi ABSTRACT

The Ireviken occurred within the early-middle Silurian, spanning the Llandovery-Wenlock boundary, ~431 million ago and has been proposed to have been initiated at least partly through the transition from late - early Silurian icehouse conditions to middle Silurian greenhouse conditions. This extinction, like many of the biotic crises throughout the Paleozoic, coincides with the rising limb of a subsequent carbon isotope excursion and the magnitude of the Ireviken carbon isotope excursion (CIE) has recorded δ13C values of ≥+4‰ worldwide. Epicontinental seaways have long been thought to have been the most likely place for high enough organic carbon burial to have the capability of producing this positive isotope excursion. However, the Ireviken extinction event and its subsequent carbon isotope excursion had both occurred during times of increased carbonate production, rather than organic carbon burial, in these shallow epeiric seas. This led to the development of a two steady- state ocean climate model, known as primo (P) and secondo (S) states, that could best explain these lithostratigraphic and biostratigraphic trends. This model demonstrates how alternating between climate states can affect ocean circulation via changing the site of deep water formation and ultimately induce anoxia throughout portions of the deep ocean. It assumes of a globally deoxygenated deep ocean with δ13C alone and lacks any direct evidence for pervasive deep ocean anoxia. It is for that reason that this study has conducted the first paired δ34S-carbonate associated sulfate (CAS) and δ13C study along with a carbonate paleoredox proxy, I/(Ca+Mg) ratio analysis, to bring new insights on the paleoredox conditions of the late Llandovery-early 13 Wenlock oceans during the Ireviken extinction and CIE. Covarying positive shifts in δ CCarb and 34 δ SCAS are consistent with the onset of sea-level rise following the end of Late Ordovician-Early Silurian icehouse conditions and represents an overall increase in the fraction of anoxic waters in 13 the global ocean. Trends in δ COrg show overall low oxygen trends and downwelling of nutrient poor water masses resulting from expansion of epeiric seas. I/(Ca+Mg) ratios during this time also indicate local and pervasive low oxygen conditions. Shallow oxic surface waters were shown to be in direct exchange and within proximity to anoxic water masses suggesting a shallow chemocline prior to the Ireviken CIE and an expanded one during. The net effect of organic carbon and pyrite burial results in an overall increase in atmospheric O2, eventually cooling climate, reestablishing thermohaline circulation, and, therefore, ending the Ireviken CIE.

vii CHAPTER 1

INTRODUCTION

Within the past 30 years, the Silurian period has been shown to have been a time of repeated marine that were believed to have been linked to abrupt and substantial changes in oceanography, climate, and the global carbon cycle [Jeppsson, 1990; Samtleben et al., 1996; Bickert et al., 1997; Jeppsson and Calner, 2002; Calner, 2004, 2008; Cramer and Saltzman, 2005; Calner and Eriksson, 2006; Barrick et al., 2010]. Research in sedimentology and biostratigraphy recognized three significant extinctions in the Silurian known as the Ireviken, Mulde, and Lau events (B.1) and, more recently, chemostratigraphic studies have demonstrated that these extinctions were linked to major positive carbon isotope excursions [Samtleben et al., 1996; Bickert et al., 1997; Saltzman, 2001, 2002; Cramer and Saltzman, 2005; Cramer et al., 2010]. The oldest of these extinctions, known as the Ireviken extinction event, occurred within the early-middle Silurian and spanned the Llandovery-Wenlock boundary ~431 million years ago [B.2; Cramer et al., 2010]. The Ireviken extinction has been proposed to have been initiated at least partly through the transition from late Ordovician- early Silurian icehouse conditions to middle Silurian greenhouse conditions [Jeppsson, 1990; Bickert et al., 1997; Díaz- Martínez and Grahn, 2007]. , graptolites, and were among some of the groups most affected during this event, where conodonts and graptolites saw a reduction in biodiversity by 80% and biodiversity decreased by 50% [Calner, 2008]. Other faunal groups that were also affected during the Ireviken extinction event include, but are not limited to, brachiopods, acritarchs, and chitinozoans [Calner, 2008]. This extinction, like many of the biotic crises throughout the Paleozoic, coincides with the rising limb of a subsequent carbon isotope excursion and the magnitude of the Ireviken carbon isotope excursion (CIE) has recorded δ13C magnitudes of ≥+4‰ worldwide [B.1; [Samtleben et al., 1996; Bickert et al., 1997; Saltzman, 2001; Cramer and Saltzman, 2007a; Cramer et al., 2010]. Although the timeframe of the Ireviken extinction itself was much shorter [Cramer et al., 2010], its CIE is comparable in terms of its duration (~1 Ma) when compared to one of the “Big Five” extinction events’ CIEs such as the late Ordovician mass extinction/ CIE [Cramer et al., 2010, 2011, 2012]. The late Ordovician mass

1 extinction and Hirnantian CIE lasted for ~1.4 Ma and had occurred over a 3nd-order sea-level cycle, with contrast to the Ireviken extinction and CIE which occurred over a 3rd-order sea-level cycle as well, with a δ13C recorded magnitude values of ≥+7‰ [Kump and Arthur, 1999; Sheehan, 2001; Brenchley et al., 2003; LaPorte et al., 2009; Young et al., 2010]. Athough the Ireviken extinction/CIE and the late Ordovician mass extinction/Hirnantian CIE differ in magnitude (HICE is >3‰ larger) and severity of extinction, the coincidence of biotic extinction and large perturbation of the global carbon-cycle are clear [King, 1937; St. John, 1966; Huff and Turkmenoglu, 1981; Sheehan, 2001; Young et al., 2010; Cramer et al., 2011, 2012]. Traditionally, global carbon cycle modeling during positive CIEs have focused on the importance of sequestration of 12C enriched organic matter as the main driver for positive CIEs [Arthur et al., 1987; Kump and Arthur, 1999; Cramer and Saltzman, 2007b]. It is for that reason that positive CIEs are often associated with deposition of organic rich black shales in epeiric sea settings [Arthur et al., 1987; Cramer and Saltzman, 2007b]. These shallow epeiric seas are considered highly productive in terms of primary producers which; when organic matter starts to decay, increases aerobic respiration/oxygen demand in the water column resulting in a decrease dissolved oxygen, and thus inducing anoxia and the preservation of enough organic matter to drive these positive CIEs [Arthur et al., 1987; Kump and Arthur, 1999; Cramer and Saltzman, 2007b] However, the Ireviken CIE contradicts this model in the sense that this CIE was found to have occurred during periods of increased carbonate production rather than organic carbon burial in these epeiric sea settings [Bickert et al., 1997;Saltzman, 2001; Cramer and Saltzman, 2005; Cramer et al., 2010]. In fact, many other CIEs throughout the Paleozoic (e.g. SPICE, GICE, Mulde CIE, Lau CIE, etc.) have also been found to have occurred during extensive carbonate deposition in shallow epicontinental seaways [Saltzman et al., 1998; Jeppsson and Calner, 2002; Young et al., 2005; Barrick et al., 2010]. This led Jeppsson (1990) to develop an oceanographic- climate state model that could explain the large-scale lithologic, biostratigraphic, chemostratigraphic trends seen in Gotland, strata during the Ireviken extinction. It wasn’t until later that Cramer and Saltzman (2005) would revise the model to account for carbon isotope trends. This model is referred to as the Primo (P) and Secundo (S) state model and demonstrates how alternating between climate states can affect ocean circulation by changing the site of deep water formation from the high latitudes to the low latitudes and ultimately induce anoxia throughout portions of the deep ocean [Jeppsson, 1990; Cramer and Saltzman, 2005].

2 Therefore, the organic carbon burial responsible for the Ireviken CIE did not happen in shallow epicontinental seaways, but instead offshore in the deep ocean, where instead a healthy and extensive carbonate factory was able to flourish in these epeiric settings globally [Cramer and Saltzman, 2005]. Although the proposed cause behind the Ireviken extinction and CIE is explained well within the context of the modified P and S state model of Cramer & Saltzman (2005, 2007), it assumes a deoxygenated deep ocean based upon carbon isotope data alone and lacks any direct evidence for pervasive deep ocean anoxia during this time [Jeppsson, 1990; Cramer and Saltzman, 2005, 2007]. The hypothesized spread of anoxic bottom waters were accompanied by a strengthened pycnocline which greatly decreased nutrient cycling to the overlying water column [Cramer and Saltzman, 2007b]. The result was a significant reduction in primary production and a potential collapse of several marine food webs thus leading to the Ireviken extinction event. However, to date no paleoredox proxy analyses have been done to assess the validity of carbon burial/anoxia model. δ34S and δ13C are investigated because they are tracers of long-term carbon and oxygen cycles linked by enhanced burial of organic matter and pyrite [Berner and Raiswell, 1983; Lyons and Gill, 2010]. The link of these two long-term cycles is through microbial sulfate reduction (MSR), a process that dominates under anoxic conditions. Organic matter acts as the limiting reagent where an increase in organic matter burial can lead to an increase in MSR and, therefore, an increase in pyrite formation and burial [Lyons and Gill, 2010]. If more deep water settings were to become anoxic during the Ireviken CIE, then enhanced organic carbon and pyrite burial should have occurred. The sequestration of 34S depleted pyrite and 13C depleted organic matter, due to biologically mediated isotopic discrimination, would then leave enriched isotopic signatures with respect towards the heavier isotopes (13C and 34S) in the global oceans. This can 13 34 be measured by looking at δ Ccarb and δ S carbonate-associated-sulfate (CAS) trends in bulk carbonate and would expect to have positive values if pervasive deep ocean anoxia was prevalent during the Ireviken extinction and CIE. This is seen in late SPICE event in which 13 34 covarying shifts of δ Ccarb and δ SCAS are due to increased sequestration of organic matter and pyrite brought on by anoxia/euxinia (H2S dominated anoxic waters) [Gill et al., 2007]. High primary production eventually led to increased anoxia/euxinia and MSR activity, which then fueled pyrite production and burial [Gill et al., 2007].

3 The destruction of Paleozoic deep ocean sediments due to subduction has proven to be problematic in terms of finding adequate sampling areas when conducting any deep time study that proposes to test changing conditions in these environments. Fortunately, epeiric seas were prevalent globally throughout much of the Silurian and allowed for the development and preservation of large carbonate successions suitable for stable isotopic and geochemical analysis on both the Laurentian and Baltic cratons [Cocks, 2001]. In order to assess the local effects of potential redox conditions in these Silurian epeiric sea sediments, a local carbonate redox proxy is needed.Duee to redox sensitive iodine , Iodine/Calcium + Magnesium ratio analysis (I/(Ca+Mg)) of carbonate rocks can be used to trace past local changes in dissolved oxygen [Lu et al., 2010]. Iodine is highly sensitive to free oxygen conditions and occurs as iodate under a well oxygenated water column [Lu et al., 2016]. Iodate concentrations decreases with dissolved oxygen and, under sub-oxic and anoxic conditions, iodide becomes the preferred species [Lu et al., 2016]. In a lab setting it has been shown that I/Ca ratios in synthesized calcite linearly increases with increasing iodate concentrations, whereas there is no observed increase with increasing iodide concentrations [Lu et al., 2010]. This means that iodate is preferentially incorporated into carbonates allowing for reconstruction of dissolved oxygen gradients in past water columns [Lu et al., 2010]. Thus, depletion in local seawater oxygenation would be reflected with an overall decrease in iodate concentrations and, therefore, lower I/(Ca+Mg) values through carbonate successions in concert with deeper ocean redox changes during the Ireviken extinction and CIE. This study has conducted the first paired δ34S-carbonate associated sulfate (CAS) and δ13C study along with an Iodine/Calcium ratio analysis, a recently developed carbonate paleoredox proxy [e.g., Lu et al., 2010], to bring new insights and constraints on the paleoredox conditions of the late Llandovery-early Wenlock oceans during the Ireviken extinction and CIE. The geochemical proxies proposed here can infer things about both global and local paleoredox conditions in Silurian oceanic environments. For this investigation, two study areas in North America have been chosen along the southern and western margins of Laurentia (B.5). These sights were selected due to previous conodont and graptolite biostratigraphy, along with δ13C work, correlating these research areas within the timeframe the Ireviken extinction event had occurred [Berry and Murphy, 1975; Klapper and Murphy, 1975; Barrick et al., 1983; Saltzman, 2001; Cramer and Saltzman, 2005]. This study will be able to directly test if Ireviken Extinction

4 and CIE were the result of a proliferation of deep ocean anoxia and increased organic carbon sequestration due to a shift in deep water formation from the high latitudes to the low latitudes [Jeppsson, 1990; Cramer and Saltzman, 2005, 2007].

5 CHAPTER 2

GEOLOGIC BACKGROUND

2.1 Early Silurian Climate and Oceans

It has been proposed that the Ireviken extinction event occurred during a major climatic shift and deglaciation as the Silurian world transitioned away from icehouse condtions that began and climaxed within the Hirnantian of the Late Ordovician [Jeppsson, 1990; Calner, 2004; Cramer and Saltzman, 2005]. It is now known that the Late Ordovician icehouse period did not end with the Hirnantian glaciation, but actually extended well into the Silurian [B.2; Grahn and Caputo, 1992; Jeppsson and Calner, 2002; Cramer and Saltzman, 2007; Díaz-Martínez and Grahn, 2007]. Stratigraphic evidence has shown that as many as four ice sheet advances occurred during the Llandovery Epoch [Grahn and Caputo, 1992; Caputo, 1998; Cramer and Saltzman, 2007]. This evidence comes from glacially derived diamictites found in the Nhamunda Formation in Brazil and have been well-dated using graptolite and chitinozoan biostratigraphy within interfingering shale units, demonstrating that they are Llandovery to early Wenlock in age and predating the Ireviken carbon isotope excursion (CIE) [Grahn and Caputo, 1992; Cramer and Saltzman, 2007b]. Similarly, conodont biostratigraphy from the thin Sacta carbonate member of the Kirusillas Formation in Bolivia, that immediately overlies the glacially derived Cancaniri Formation, has confidently been identified as early Wenlock in age and deposited during the Ireviken CIE [Diaz-Martinez, 1997]. This is due to the occurrence of the conodont Ozarkodina sagitta rhenana, this conodont species range was exclusively during the Ireviken CIE, and thus showing the glaciation responsible for the Cancaniri Formation diamictites had preceded the rising limb of the Ireviken CIE [Jeppsson, 1990; Calner, 2004; Cramer and Saltzman, 2005]. In addition, a major unconformity near the Llandovery-Wenlock boundary is present globally, documented everywhere from Nevada to Gotland, Sweden, and has been interpreted as a major eustatic sea level fall linked to icesheet advance on Gondwana [B.2; Cramer and Saltzman, 2005; Cramer et al., 2010]. Not only are extensive and thick carbonate sequences seen directly above this unconformity in low latitude shelf and epeiric sea settings, but 13 a positive shift in δ Ccarb values is recorded globally (Ireviken CIE) in these sequences (B.3). It

6 has been demonstrated that the Ireviken CIE occurred during an expansion of epieric seas and carbonate platform environments brought on by a major eustatic rise in sea level (B.2). Epicontinental seaways have long been thought to have been the most likely place for burial of copious amounts of organic carbon that could produce a positive carbon isotope excursion, due to the fact that modern continental shelf/margin settings are the places where the overwhelming majority of organic matter is being buried today [Berner and Canfield, 1989; Berner, 2004]. However, contrary to this idea, the Ireviken extinction event and its subsequent CIE both occurred during times of increased carbonate production rather than organic carbon burial in these shallow epeiric seas [Calner, 2008]. This has been problematic in terms of finding a reasonable explanation for the large-scale lithologic, biostratigraphic, and chemostratigraphic trends seen not only during the Ireviken event, but other Silurian extinction events and CIE’s as well. The cyclicity of conodont extinctions and carbonate successions seen in Gotland, Sweden strata led Jeppsson [1990] to develop an oceanographic model that could explain these observed patterns. This model was later refined by Cramer and Saltzman [2005] to incorporate changes in 13 the long-term carbon cycle that were documented from the associated positive δ Ccarb excursions (e.g., Saltzman, 2001). The Jeppsson (1990) model describes two steady ocean-climate modes known as primo (P) and secondo (S) states (B.4). It demonstrates how alternating between climate states can affect ocean circulation via changing the site of deep water formation and ultimately induce oceanic stratification and anoxia throughout portions of the deep oceans [Jeppsson, 1990; Cramer and Saltzman, 2005, 2007]. Primo states are proposed to occur during times of icehouse conditions, characterized by cold high latitudes and humid low latitudes resulting in vigorous thermohaline circulation and a well-oxygenated deep ocean [Jeppsson, 1990, 1998]. During the late Llandovery, and prior to the Ireviken event, the high latitudes experienced polar cooling brought on by the Cancaniri Glaciation [Cramer and Saltzman, 2007b]. Deep waters would have formed at the poles where the sinking of cold dense surface waters depleted in oxygen and essential nutrients were recycled, replenished, and then delivered to overlying water columns at zones of upwelling (B.4a). This, in combination with enhanced nutrient delivery from continental runoff at the lower latitudes, allowed for highly productive photic zones and deposition of organic rich black shales along these coastal margins (B.4a). High nutrient availability is also supported by massive bedded cherts found within the basal Roberts Mountains Formation, Nevada, [Klapper and

7 Murphy, 1979; Berry and Murphy, 1979] indicative of intense ocean upwelling along the western Laurentian margin [Pope and Steffen, 2003; Pope, 2004] The high nutrient flux, terrigenous input, and lower salinity due to increased precipitation at the lower latitudes inhibited carbonate production and favored conditions for primary producers [Calner, 2008]. This resulted in poor conditions for carbonate factories globally, leading to deposition of marls and argillaceous limestones along continental shelves (B.4a). Due to thermohaline circulation and a well oxygenated deep ocean, the overall fraction of anoxic waters within the global ocean was still too low to perturb the global carbon cycle resulting in isotopically light and unvarying δ13C values during this time [Saltzman, 2001; Cramer and Saltzman, 2005]. In contrast, Secundo states (B.4b) occur during greenhouse conditions and are characterized by having warmer high latitudes and arid low latitudes. As stated above, the early Wenlock was thought to have experienced climatic warming that significantly reduced the flux of sinking surface waters at the high latitudes [Cramer and Saltzman, 2005]. This resulted in diminished thermohaline circulation and significantly reduced the formation of deep water at these high latitudes [Cramer and Saltzman, 2007b]. Conversely, sites of deep water formation likely became more prevalent in lower latitudes where the increased aridity resulted in higher rates of evaporation in the flooded continental interiors and induced, salinity-driven, halothermal circulation [B.4b; Bickert et al., 1997]. Within tens of thousands of years, the sinking of oxygen deficient warm saline bottom water in these low latitude epeiric seas resulted in a strengthened pycnocline, stratifying the deep ocean and creating conditions favorable for widespread deeper ocean anoxia (B.4b). The downwelling of saline water masses and reduction in nutrient availability allowed for an expansion of carbonate platform environments throughout low- latitude epicontinental shelf settings [Copper and Kershaw, 1998; Cramer and Saltzman, 2005]. It has been proposed that more prevalent deep ocean anoxia resulted in sequestration of 13 isotopically light organic carbon that subsequently led to the increased δ Ccarb values observed during the early Wenlock and thus the driver of the global carbon cycle perturbation, Ireviken CIE [Cramer and Saltzman, 2005]. Incorporation of global carbon cycle perturbations (trends in 13 δ Ccarb values) into the Jeppsson (1990) model provided evidence for the mechanism for 13 alternating lithologies and associated trends in δ Ccarb values recorded in these epicontinental seaway environments [Cramer and Saltzman, 2007b].

8 2.2 Depositional Environment, Facies, and Sequence Stratigraphy

The Newsom Roadcut field site is located west of Nashville, TN, (north of I-40 along McCrory Lane) and is interpreted to have been deposited in a shallow inner to mid carbonate rimmed-shelf environment with direct connection to the Rheic Ocean (B.5). The shelf itself was located just off the western flank of a broad structural uplift (B.5), known as the Nashville Dome, that made up the southern extent of the Cincinnati Arch, and is interpreted at this time to be a fault-bounded positive topographic feature overlying a Neoproterozoic rift basin [Beaumont et al., 1988]. The Central Appalachians had experienced several tectonic episodes (i.e., Taconic, Acadian, and Alleghenian) prior to and after the deposition these carbonates, however, this part of the Cincinnati Arch experienced minimal thermal heating and only minor structural uplifts during these events [Beaumont et al., 1988]. This is further confirmed as explained in the discussion section later when looking at potential diagenesis. The , a skeletal wackestone to packstone containing horizons of black nodular chert (B.7), is the oldest unit at this locality. The major Late Llandovery unconformity separates it from the overlying Wayne Formation (Maddox Member). Conodont biostratigraphy from the Maddox Member has constrained the timeframe of its deposition to span from the latest Llandovery to Early Wenlock [Barrick et al., 1983]. The base of the member correlates with Pterospathodus amorphognathoides condont biozone, this conodont that has been found globally at or near the base of the Ireviken CIE [Saltzman, 2001; Calner, 2004;Cramer and Saltzman, 2005, 2007; Cramer et al., 2010]. Due to this section being within an continental interior basin it experienced a longer erosional/non-depositional effects during late Llandovery eustatic sea-level lowstand, and therefore the basal Wayne Formation fails to record Ireviken CIE baseline and some of the rising limb δ13C values (B.7).The base of the Maddox Member consists of a calcareous ferruginous siltstone unit, interpreted to represent the onset of eustatic sea-level rise, that shortly transitions into an argillaceous packstone at ~5.2 meters. Continued early Wenlock sea-level rise is indicative of the shift in lithology to a skeletal packstone to wackestone, eventually reaching max highstand as interpreted by the change to skeletal wackestone around 10m. The transition back to skeletal packstone to wackestone represents falling stand and an overall eustatic sea-level fall (B.7).

9 The Pete Hanson Creek-II section is found in the Roberts Mountains of central Nevada (B.5) where the lower portion of the Roberts Mountains Formation was collected. This section was located off the western coast of Laurentia, connected to the Panthalassic Ocean, deposited in an upper slope setting near the shelf-slope break (Johnson and Murphy, 1984; Harris and Sheehan, 1997). A thick horizon of bedded black chert, indicative of a highly productive photic zone, marks the base of the section (B.6) and represents upwelling off the Laurentian margin [Pope and Steffan, 2003]. The section experiences a major change in lithology into an extensive carbonate lime mudstone facies (B.6), coinciding with a rise in eustatic sea-level and changing oceanographic conditions, that deposition likely continued until maximum highstand. Towards the top of the section around 40m, it transitions into thinly-bedded argillaceous wackestone which has been interpreted to have been deposited during late HST to falling stand (FSST).The conodont biostratigraphy done on the Roberts Mountains Formation [Berry and Murphy, 1979], places the timing of the chert within the P. celloni conodont biozone during the Late Stage and the base of the lime mudstone within the P. amorphognathoides biozone, which again coincides with the onset of the Ireviken extinction event [Calner, 2004;Cramer et al., 2010]. Though the upper portion of this section is lacking conodont biostratigraphic control (B.6), a previous δ13C study has recorded the complete Ireviken CIE at this section [Saltzman, 2001]. In terms of tectonics, this area did experience some mountain building due to the Antler Orogen during the Late , in which, Lower Paleozoic oceanic rocks of the Roberts Mountains allochthon were thrust onto the North American shelf [Murphy et al., 1984]. However, based on field observations, the Late Llandovery–Early Wenlock strata this study analyzed only experienced uplift and remained tectonically unaltered.

10 CHAPTER 3

METHODS

3.1 Sample Preparation and Extractions

13 For δ Ccarb analysis, bulk carbonate samples were micro-drilled on fresh surfaces targeting micritic matrixes whenever possible while also avoiding any calcite veins and or skeletal grains. To minimize risk of cross-contamination, 0.1 M HCl was used to dissolve any 13 residual carbonate powder off the aluminum oxide drill bit between samples. For δ COrg, 34 δ SCAS, and I/(Ca+Mg) analysis, bulk carbonate samples were first prepared by removing all weather surfaces and veins using a water-cooled diamond blade fitted rock saw. A subset of samples from the Newsom Roadcut field site were further slabbed and polished to produce thin- section billets for petrographic analysis. The remaining samples were then broken into chips and powdered by grinding in an agate mortar and pestle or using a SPEX 8510 Shatterbox with alumina ceramic containers. 13 For analysis of δ Corg, 2-3 grams of sample powder for each sample were weighed out into clean 50 mL centrifuge tubes and reacted with 6.0 M HCl to remove all carbonate minerals. Samples were then centrifuged to separate the insoluble residues. This process was repeated twice and then samples were rinsed three times with 50ml of 18.2MΩ ultrapure water, centrifuged, and left to dry overnight in a 70C oven. Dried residuals were then homogenized within an agate mortar and pestle. Anywhere from 0.3–12 mg of residual powder was then 13 weighed into tin capsules for δ Corg analysis. 34 Approximately 60-180 g of carbonate powder were weighed out for analysis of δ SCAS, using a similar CAS extraction procedure as outlined by Wotte et al. (2012). Each sample was rinsed for 12 hours with 400 mL of 10% sodium chloride solution (NaCl(aq)) for a total of 3 times with samples stirred once between each rinse. Each sample was then rinsed, stirred, and decanted an additional 3 times for 12 hours each, but this time with 400 mL of 18.2MΩ ultrapure water to remove all leached water soluble sulfate from the samples. Samples were decanted after the last rinse and acidified using 6 M HCl for ≤ 2 hours to dissolve carbonate minerals and liberate the incorporated sulfate from the matrix but not oxidize any sedimentary pyrite fraction. The acidified solution was then centrifuged to separate insoluble sample residues from the solution.

11 Sample solutions were then brought up to a pH of 10, via NaOH, to precipitate and remove dissolved iron oxides and metal-organic complexes from solution. These were separated from solution via vacuum filtration using clean 47 mm GF/C and/or GF/F filters. Once filtered, a few drops of concentrated nitric acid was added to the remaining solution to adjust the pH to 4, after which excess amount 0.2 M BaCl solution was added to each sample and left to sit for 48-72 hours to precipitate barite (BaSO4). Sample solutions containing barite were then decanted, centrifuged, and transferred to pre-baked dram vials where they were then placed in an 70C drying oven overnight. I/(Ca+Mg) analysis was carried out using a similar procedure similar as described by Lu et al. (2010). Approximately 3-6 mg of carbonate powder was weighed out for each sample, 2+ dissolved in 3% HNO3 solution, and then diluted to keep a consistent matrix of 50 ±5 ppm [Ca ] for precise iodine measurements (ppm). Due to the volatility of iodine, tetramethylammonium hydroxide (TMAH) was used to stabilize it and decrease the memory effect during ICP-MS measurements [Muramatsu and Wedepohl, 1998; Tagami and Uchida, 2005]. For that reason, a background matrix of 0.5% HNO3 and 0.5% TMAH was made and used for standards and samples. Iodine volatility loss in this matrix was found to be negligible within one to two days [Lu et al., 2010]. As an added precaution, however, samples were analyzed [I, Ca, Mg] no longer than 2 hours after sample preparation.

3.2 Stable Isotope and I/(Ca+Mg) Analyses

13 18 Sample powders for δ Ccarb and δ Ocarb were weighed out between 250 – 350 micrograms (g), transferred to glass exotainer vials, and then baked in an oven at 100C to remove any moisture and any trace volatile organics. Next the vials were capped and flushed o with helium then reacted with 3-5 drops of 100% phosphoric acid (H3PO4) for 24 hours at 25 C to produce CO2 for isotopic analyses. The carbon dioxide from the carbonate powder was then measured using a Gas Bench II Autocarbonate device connected to a Thermo Finnigan Delta Plus XP stable isotope ratio mass spectrometer (IRMS) at Florida State University’s National High Magnetic Field Laboratory. Precision and calibration of data are monitored through routine analysis of the IAEA NBS-19 standard, and in-house standards adjusted relative to NBS-19. All

12 results are reported in standard delta (δ) notation with units reported as parts per thousand or permil (‰) relative to the VPDB (Vienna Pee Dee Belemnite) standard for δ13C as: 13 13 12 13 12 δ CCarb= [( C/ C)sample / ( C/ C)VPDB – 1] x 1000 (Equation 3.1) The analytical precision and standard deviation, based on replicate analyses of lab 13 standards processed with each set of samples, is ±0.05‰ or better for δ Ccarb and ±0.1‰ for 18 δ Ocarb. 13 The δ Corg values were measured using a Carlo Erba Elemental Analyzer coupled to a Thermo Finnigan Delta Plus XP IRMS at Florida State University’s National High Magnetic Field Laboratory. Precision and calibration of data are monitored through routine analyses of in- house standards that are calibrated against IAEA standards. All results are reported in standard delta (δ) notation with units reported as parts per thousand or permil (‰) relative to the VPDB 13 (Vienna Pee Dee Belemnite) standard for δ Corg (Equation 3.1).The analytical precision and standard deviation, based on replicate analyses of in-house lab standards processed with each set 13 of samples, is ±0.2‰ or better for δ Corg and ± 1.5% for %C. Weight % of total organic carbon (TOC) in samples is determined by comparison of voltages of the beam intensities of masses 44, + 45, and 46 CO2 with our samples and known wt.% carbon of the gravimetric standard acetanilide, the uncertainty is ±5% or better [Young et al., 2008]. Homogenized barite samples were then weighed out (0.36-0.44 mg) and excess 34 V2O5 loaded into tin cups for stable sulfur isotopic analysis. Barite samples for δ SCAS analysis were sent to the Stable Isotope Research Facility at Indiana University for isotopic analysis. These samples were measured using a Costech Elemental Analyzer coupled to a Thermo Delta V Plus IRMS. Calibration of samples were done based on laboratory standards calibrated relative to the IAEA S-1 (δ34S, −0.30‰) and NBS-127 (δ34S, +20.3‰) standards. All results are reported in standard delta (δ) notation with units reported as parts per thousand or permil (‰) relative to the 34 standard VCDT [Vienna Canyon Diablo Troilite] standard for δ SCAS as: 34 34 32 34 32 δ SCAS = [( S/ S)sample / ( S/ S)VCDT – 1] x 1000 (Equation 3.2) The analytical precision based on replicate analyses of lab standards and duplicate 34 samples processed was ±0.2‰ or better for δ SCAS. I/(Ca+Mg) was measured using an Agilent Quadro-pole inductively-coupled-plasma mass-spectrometer (ICP-MS). Potassium iodate (Fisher Scientific product P253-100) was dissolved, diluted gravimetrically, and mixed with pure elemental standards of Ca2+ and Mg2+ to

13 make the standard curve in which samples were calibrated too. The accuracy of this procedure was based on measurements of a known reference material (KI1-1; e.g. Hardisty et al., 2014) and was found to be within ±0.5% of the known value. Additionally, precision based on replicate analyses of this same reference material and duplicate samples processed was ±0.08 umol/mol or better.

14 CHAPTER 4

RESULTS

4.1 Newsom Roadcut, Nashville, Tennessee

The Newsom roadcut section was sampled along McCrory Lane located just north of I-40 and due west of Nashville, Tennessee (B.5). Samples were taken starting 4m below a major unconformity/contact in the upper Brassfield Formation and then continuing 12m above in the overlying Wayne Formation, Maddox Member [A.1]. Geochemical analysis done on the Brassfield Formation is for chemostratigraphic correlation and helping to constraining missing time between these formations due to lack of conodont biostratigraphy for the Brassfield 13 Formation in this section. At the base of the Brassfield Fm δ Ccarb values [Table. 1] start at - 0.26‰ and gets as low as -0.42‰ from 0 to 0.4m before steadily rising to +0.93‰ at 3.6m. 34 δ SCAS show relatively stable values showing small variations ranging from +24.9‰ to 13 +26.6‰., unlike δ Ccarb which show a steady, small positive increase. Additionally, as an internal check on CAS extraction methodology, two separate CAS extractions from the same 34 interval (2.4m) yielded barite that produced almost identical δ SCAS values, within +0.1‰ of one another. Sampling of the Maddox Member of the Wayne Formation started immediately above the 13 unconformity (at 4m), that separates the Brassfield from Wayne formations. Previous δ Ccarb 13 and δ Corg values done by Cramer and Saltzman (2005, 2007) for this section are being 34 compared to the new analyses of δ SCAS and I/Ca+Mg values presented here in this study [A.1]. 13 δ Ccarb values start to rise at the base of the Maddox member from +2.4‰ to as high as +3.6‰ within the first 40cm. The first peak is around 5.37m with a value of +3.9‰ and peaks again at 7.62m with a value of +3.9‰. After, the values slowly fall back down to a baseline of +1.0‰ by 13 12.12m. In contrast, δ Corg fluctuates very little at the base of the Maddox, with values staying close to -28.0‰, ranging between -27.8 to -29.3‰ overall, until 6.75m. Values then begin to steadily rise with peak values of -25.5‰ at 10.25m where it then maintains values of around - 26.0‰ through the top of the section. 34 δ SCAS values (A.1; B.7) at the base of the Maddox (4m) start out heavy with the first value of +26.3‰. Values remain steady between 4 to 10m, ranging between +23.6‰ to +25.4‰.

15 34 13 These δ SCAS values correspond to the most positive δ Ccarb values of Ireviken CIE at this 34 section. δ SCAS values remain heavy within the skeletal wackestone facies after corresponding 13 34 δ Ccarb values begin returning to baseline values at 9m. Eventually δ SCAS values decline from +26.0‰ at 10m to low values of +18.7‰ at 14m in the skeletal packstone/wackestone facies. 13 This is noticeably two meters higher in the section after δ Ccarb values have established a post- 34 Ireviken CIE baseline. At 15 m, within the upper Maddox Member δ SCAS values begin rising again to +24.0‰ and this may correspond to the beginning of the mid-Wenlock Mulde CIE which is documented to begin in the upper Maddox Member of the Wayne Formation (Cramer et al., 2006). I/Ca+Mg values [I mol/ Ca+ Mg mol] start relatively high, ~2.43, close to the base of the Maddox member at 4.25m then dropping as low as 1.15 at ~4.5 before reaching a peak of 2.49 at 6.5m [B.7], all this variability occurring within the calcareous ferruginous siltstone facies. 13 34 This decrease at 4.5m doesn’t show any distinct patterns with δ Ccarb and δ SCAS trends, as values remain high prior to and after this decrease in I/(Ca+Mg) ratios (B.7). After, values begin 13 34 to exponentially decrease in concert with high δ Ccarb and δ SCAS values, where I/(Ca+Mg) 13 begins to reach lower values ranging between 0.10 and 0.60 after δ Ccarb starts returning to 34 baseline and δ SCAS reaches its second peak at 10m (B.7). Values then increase back up to 1.12 34 at the top of the section and could correspond with δ SCAS values in terms of recording the Mid- Wenlock Mulde CIE as discussed earlier [Cramer et al., 2006]

4.2. Pete Hanson Creek-II, Roberts Mountains, Nevada

Samples were collected from the Pete Hanson Creek II section in the Roberts Mountains of central Nevada (B.5) and originally litho/biostratigraphically described by Klapper and Murphy (1975); Berry and Murphy (1975). The base of the section starts above a major unconformity followed by 8 meters of bedded chert (B.6) in the P. Celloni conodont zone [Klapper and Murphy, 1975]. The middle to upper part of the section falls into the Crytograptus Perneri graptolite zone, , allowing for time constraint where conodont data is lacking [Klapper and Murphy (1975); Berry and Murphy (1975)]. Overall, 46.5m of the Roberts Mountains formation section was sampled in even intervals whenever possible.

16 13 δ Ccarb values [A.2; B.6] start low in the lowermost lime mudstone/Skeletal wackestone facies, and the P. celloni conodont zone, of the section starting at 0.69‰ and then shifting lighter within 4.5m to +0.19‰. After, values begin to gradually climb in the Pt. amorphognothides zone (B.6) where the first peak occurs at 12m at +4.23‰ and then again within 4.5m of the first peak 13 13 at +4.06‰. δ Corg co-varies with δ Ccarb trends, in which, values start light at -28.6‰ towards 13 the base (B.6), shifting positive around 10.5m, and hitting its first peak close to δ Ccarb’s first 13 peak at 13.5m (~-24.0 ‰). δ Ccarb values remain close to +3.00‰ unitl about 34.5m where they then begin to decline in the thinly-bedded argillaceous wackestone facies to values between 0‰ 13 to -0.64‰ starting at 45m. δ Corg start to shift more negative till around 41m, at the base of the thinly-bedded argillaceous wackestone, where they see a slight shift to more positive values at the top of the section. 34 The δ SCAS (A.2; B.6) values for the lower Roberts Mountains Formation start out at +21.1‰ within the basal carbonates of the section, P. celloni conodont biozone, but shift to +35.3‰ at 6m, Pt. amorphognothides zone, in conjunction with positive shifts in δ13C. Values 13 34 typically then stay between +33.0-39.0‰ and unlike δ Ccarb, δ SCAS values don’t see a return to lighter values at the top of the section. The I/Ca+Mg values (Table.2; B.6) show very little variation within the studied interval at this field site. Values only range between 0 to 0.340 umol/mol throughout the entire section. No real trend can be discerned from this data as these values represent extremely low values and could be the result of very oxygen-poor marine waters in this upper slope setting or due to overprinting primary elemental signatures brought on by diagenesis, see discussion below for more details (Section 5.1).

17 CHAPTER 5

DISCUSSION

The δ13C values from the Pete Hanson Creek II section document the full extent of the Ireviken CIE (rising limb, peak values, falling limb/baseline values), however the Newsom Roadcut only records a small part of the Ireviken CIE rising limb (B.7) due to this widespread regional unconformity (discussed above). A chemostratigraphic investigation of the underlying Brassfield Formation, presented here estimates that almost the entire and Telychian 13 Stages are missing here between the Brassfield and Wayne formations (B.7). A δ CCarb shift from -0.3‰ at the base of the section to +1.0‰ at the top may represent the very onset of late –Early Aeronian CIE [Kaljo and Martma, 2000; Poldvere, 2003; Cramer et al., 2011; Frýda and Štorch, 2014]. This would also be consistent in terms of age with conodont biostratigraphy from the Brassfield Formation in the southwest Ohio region [Schneider and Ausich, 2002]. Thus, a rough estimate of ~11 m.y. is missing from this section at the unconformity between the Brassfield and Wayne formation’s. In North America, the Ireviken CIE not only shows a consistent relationship with 34 34 13 lithologic changes, but also shares a similar relationship with δ SCAS. δ SCAS and δ CCarb both shift positively almost immediately after the major unconformity at the base of the Newsom Roadcut and an abrupt sediment facies shift near the base of the Pete Hanason Creek II section 34 [B.6; B.7]. Similarly, I/(Ca+Mg) values show a somewhat analogous trend to that δ SCAS and δ13C in terms of environmental implications. Before this data can be addressed, however, evaluation of potential diagenesis must first be looked at to determine its effect on primary 34 geochemical signatures presented in this study. Subsequently, δ SCAS and I/(Ca+Mg) data presented here will be evaluated in terms of changes in paleoceanographic and paleoredox conditions that could explain these trends in stable isotopic and elemental geochemistry from both sections. This study will also place these environmental constraints and implications on both global and local paleoredox conditions during the Ireviken CIE within the broader context of the previously proposed P and S state oceanographic model [Jeppsson, 1990, 1998; Cramer and Saltzman, 2005]. This new geochemical data provides the first examination of paleoredox 34 conditions throughout the late Llandovery–early Wenlock. Furthermore, these paired δ SCAS and

18 13 13 I/(Ca+Mg) measurements with δ CCarb and δ Corg analyses provide the first direct evidence of changing paleoredox conditions during late Llandovery–early Wenlock, suggesting that expansion of oceanic anoxia (possibly euxina) were the driving mechanism for the Ireviken extinction event and CIE.

5.1 Evaluation of Diagenesis on Primary Geochemical Signatures

Due to the chemical nature of carbonate rocks, diagenesis has the potential to affect the preservation of primary geochemical signatures. One of the strongest lines of evidence for primary geochemical signatures being preserved is the consistency of isotopic trends observed at both study sections, both of which can be correlated independently using conodont 13 biostratigraphy and previous δ Ccarb (Saltzman, 2001; Cramer and Saltzman, 2005). Additionally, many studies have successfully used cross-plots of seawater geochemical proxies to show the degree of alteration in carbonates during diagenesis [Banner and Hanson, 1990; Kaufman et al., 1991; Melim et al., 2001; Allan and Matthews, 2009; Young et al., 2016]. The effects of diagenetic processes on carbon isotope values in carbonates have been well-studied and been shown to be essentially ‘rock buffered,’ less prone to alteration than corresponding oxygen isotope values [Banner and Hanson, 1990; Melim et al., 2001; Allan and Matthews, 2009]. However, negative shifts in carbon isotopes below or at unconformity/ exposure surfaces have been shown to occur when meteoric fluid-rock interactions in carbonates occur, with meteoric fluids rich in dissolved organic matter that are isotopically much lighter than most marine carbonates [Marshall, 1992; Railsback et al., 2003; Jones et al., 2016]. Linearity between 13 18 δ Ccarb and δ Ocarb values often is also indicative of diagenetic alteration caused by fluid-rock interactions with meteoric fluids [Banner and Hanson, 1990; Melim et al., 2001; Allan and Matthews, 2009]. Our data, however, shows no such trend in carbonates from the Pete Hanson 2 13 Creek II section (B.4a) having an R value >0.1. Unlike the Nevada section, previous δ CCarb 18 and δ OCarb work done by Cramer and Saltzman (2005) on the Newsom Roadcut section in 13 18 Tennessee has shown correlation between δ CCarb and δ OCarb trends (B.4a). This section 18 though shows a trend and variability in δ OCarb values similar to trends observed in brachiopods in Gotland, Sweden and is consistent with the view that primary carbon isotope signatures have been preserved [Samtleben et al., 1996; Bickert et al., 1997; Cramer and Saltzman, 2005]. 13 Diagenetic alteration of primary δ COrg signatures can also happen because of oxidative loss and

19 thermal alteration of organic compounds [Kaufman et al., 1991; Young et al., 2016]. Correlation 13 between δ COrg and wt.% TOC is anticipated if differential alteration of organic matter occurred in strata that are lean versus relatively rich in organic carbon [Kaufman et al., 1991; Young et al., 2008]. This study, and previous studies (e.g. Cramer and Saltzman, 2007), finds no such 2 13 relationship (R >0.1) between δ COrg and wt.% TOC (B.4b) at these sites and, therefore, 13 believes primary δ Corg values remain preserved. 34 Additionally, recent studies on the diagenetic effects on δ SCAS values in carbonates [Gill et al., 2008] and late limestones in Nevada [Sim et al., 2015] have shown that although meteoric diagenesis can greatly lower CAS concentrations, ultimately, the 34 34 primary δ SCAS values remain largely preserved [Young et al., 2016]. δ SCAS signatures are also 2 18 34 interpreted to be primary, finding little to no relationship (R >0.2) between δ Ocarb and δ SCAS 2 18 (B.4c) and no relationship (R <0.1) between δ Ocarb and CAS(ppm)% [B.4d; Young et al., 2016]. Another important potential concern to address is incorporation of contaminant sulfate into the CAS extracted fraction. Incorporation of oxidized sedimentary pyrite during diagenesis 34 or chemical extraction of CAS can artificially lower δ SCAS values. Although, CRS extractions of sedimentary pyrite were not performed for this study, detailed petrographic examination of carbonate strata from both sections yielded little to no pyrite. Furthermore, horizons and stratigraphic intervals that had visible pyrite or iron-oxides were specifically avoided. 34 In a similar study as Gill et al. (2008) done on the same rocks as δ SCAS diagenetic investigations, meteoric diagenesis has found to greatly lower Iodine concentrations in carbonates where I/(Ca+Mg) ratios are almost completely analogous to carbonates deposited under an anoxic water column [Hardisty et al., 2017]. In no cases, however, has diagenesis been found to increase I/(Ca+Mg) ratios, due to reducing conditions in pore space, and minimizes the chances of a “false positive” [Hardisty et al., 2017]. Additionally, in some instances, dolomitization of carbonates and precipitation of authigenic dolomites in reducing marine pore fluids can have substantially lower I/(Ca+Mg) ratios in contrast to related primary carbonates that show higher ratios and reflects original formation under a well-oxygenated water column [Hardisty et al., 2017]. Although, it’s also been shown that many Paleozoic dolomites can in fact preserve primary seawater I/(Ca+Mg) ratios; having been formed penecontemporaneously with deposition at the sediment water interface and, therefore, in open exchange with seawater [Fritz and Smith, 1970; Tucker, 1982; Fairchild et al., 1991; Hood and Wallace, 2012; Husson et al.,

20 2015; Hardisty et al., 2017]. From a combination of petrographic observations and overall low ICP-MS measurements of [Mg] (ppm) during I/(Ca+Mg) analysis; dolomite is inferred to make up a very minor carbonate component in samples from both field sites and it’s unlikely that decreases in I/(Ca+Mg) ratios were due to dolomitization of primary fabrics. The lack of any substantial correlation between geochemical datasets and diagenesis proxies also further strengthens the argument that primary I/(Ca+Mg) ratios are preserved as well.

34 5.2 Mechanisms for Late Llandovery–Early Wenlock δ SCAS Trends: a Global Redox Proxy

The main source of carbon and sulfur to the global oceans is from crustal and oxidative weathering of carbonates, sulfur-bearing minerals/evaporites, and highly organic sedimentary rocks [Berner and Raiswell, 1983; Canfield, 2000; Kurtz et al., 2003; Young et al., 2016]. Additionally, oxidation of sulfide from continental-, arc-, and mid-ocean ridge volcanism is another input of sulfur to the global ocean reservoir, however, is negligible when compared to riverine flux [Alt, 1995; Canfield, 2000; Kurtz et al., 2003]. The main removal and output of sulfur from the global oceans is through the deposition of sulfate evaporates and burial of sedimentary pyrite [Bottrell and Newton, 2006]. Sulfurs link to the long-term carbon cycle, is largely controlled through the process of microbial sulfate reduction (MSR), an anaerobic respiratory pathway that flourishes under an anoxic water column, and the gateway to sedimentary pyrite formation [Berner and Raiswell, 1983]. The byproduct of this process is hydrogen sulfide (H2S) and, due to biologically mediated kinetic fractionation, is typically isotopically enriched in regards to the lighter isotope 32S. In the modern euxinic environment of 34 the Black Sea, H2S produced by MSR here has been shown to reach isotopic depletions in S by up to 70‰ [Fry et al., 1991; Sim et al., 2011]. There is little to no isotopic fractionation associated with sulfur and sulfur mineralization from volcanic sources (0-3‰) and formation/dissolution of evaporites, to which, they are considered to have little effect on isotope variability in the marine sulfate reservoir. [Ault and Kulp, 1959; Thode et al., 1961; Fritz and Smith, 1970]. Therefore, changes in fluxes of volcanic sourced sulfur and/or evaporite formation 34 were not a primary driver of the positve shift in δ SCAS we document during the Ireviken CIE. Instead, the primary drivers of sulfur isotopic change of marine sulfate on geologic timescales

21 are sedimentary pyrite burial, riverine input of weathered products of sulfur minerals, and 34 34 34 changes in Δ S (δ SCAS - δ SPyrite) [Garrels and Lerman, 1984; Gomes and Hurtgen, 2013]. Weathering has previously been hypothesized to be a potential driver behind shifts in δ13C during the Ireviken CIE [Bickert et al., 1997], which is applicable to δ34S shifts as well. This hypothesis, however, fails to explain several key factors associated with these δ13C and δ34S trends. For instance, the rising limb of Ireviken CIE has been found to coincide with a major Silurian extinction event and weathering fails to provide a kill mechanism for the extinction [Jeppsson, 1990; Samtleben et al., 1996; Bickert et al., 1997; Jeppsson and Calner, 2002; Calner, 2004, 2008; Cramer and Saltzman, 2005; Calner and Eriksson, 2006; Barrick et al., 2010]. Additionally, the Ireviken CIE occurred completely during a eustatic-sea level rise (Loydell, 1998; Cramer and Saltzman, 2007; Johnson et al., 2010) and, therefore, would have experienced a decrease in continental weathering flux due to overall less land exposure. This is consistent with the lithology observed at the Newsom Roadcut (B.7) where, facies shift from a calcareous ferruginous siltstone (shallow nearshore) to a skeletal wackestone (mid-shelf open marine). Given these facts it appears very unlikely that continental weathering and riverine input 13 34 were the primary drivers of the observed δ CCarb and δ SCAS trends in late Llandovery-early Wenlock strata. The fractionation in sulfur isotopes recorded from paired sulfate-sulfide δ34S 34 34 34 34 34 measurements is often reported as Δ S (Δ S = δ SCAS, SO4 – δ Spyr). The δ Spyr values reflect the average of local water-column pyrite and/or pyrite formed within sediment pore waters that is often further 32S-enrinched, while CAS does not incorporate the evolved sediment porewater sulfate isotopic signature, but rather reflect sulfate formed in marine surface waters (e.g. Lyons et al., 2004).Therefore Δ34S values do not always equal the initial in situ sulfur isotopic 34 fractionation due to MSR (SR) but rather an averaged kinetic isotopic signature, Δ S  SR. 34 34 Several factors that influence Δ S have also been invoked as drivers for shifts in δ SCAS [Habicht et al., 1998; Sim et al., 2011; Gomes and Hurtgen, 2013; Jones and Fike, 2013; Leavitt et al., 2013]. Changes in marine sulfate concentrations (reservoir size) have been used to explain 34 34 stratigraphic changes in Δ S and δ SCAS values [Kah et al., 2004; Fike et al., 2006]. However, for a changing sulfate reservoir to have such a large isotopic effect on biological fractionation, the marine reservoir of sulfate (estimated a range of 4mM – 15mM e.g. Horita et al., 2002; Gill et al., 2007) during the Ireviken CIE would have to increase by an order of magnitude and then

22 decrease back down to an initial several hundred micromolar reservoir size all within <1 Ma timescale which is unrealistic given what we know about rates of long-term sulfur cycling [e.g, Jones and Fike, 2013]. In order to increase the marine sulfate reservoir size, increased weathering and delivery of oxidized sulfur-bearing via riverine input [Berner and Raiswell, 1983] had to have occurred which, as mentioned previously, was highly unlikely due to the Ireviken CIE occurring during eustatic sea-level rise (B.2). Low metabolic rates of MSR have been found to produce large sulfur isotope fractionations, where high rates of MSR are associated with small fractionations between sulfate and sulfide [Kaplan and Rittenberg, 1964; Chambers and Trudinger, 1979; Sim et al., 2011]. Temperature of sea-water has been found to affect the rate of MSR in which lower temperatures would result in low metabolic rates (larger SR) and higher temperatures would result low MSR rates (smaller SR) [Jones and Fike, 2013]. Sea surface and thus global temperatures would had 34 to have been trending from high low to produce the positive shift δ SCAS values we document during the Ireviken CIE, which is inconsistent with oxygen isotopes/ sea-surface temperatures which indicate a warmer climate [Bickert et al., 1997; Trotter et al., 2016]. Glacial sedimentary records also indicates icehouse conditions ended prior to the onset of the Ireviken CIE [MacQuown et al., 1984; Grahn and Caputo, 1992]. In addition to temperature, availability of labile organic matter has been suggested to affect MSR rates, however, this would require uniform shift labile organic matter delivery on a global scale to cause this positive shift in 34 δ SCAS which seems unrealistic. Furthermore, a lack of a developed terrestrial biosphere limits the type of organic matter used in MSR and therefore, limits the amount of fractionation (different electron donors, different fractionation rates) [Kaplan and Rittenberg, 1964; Hagström 34 and Mehlqvist, 2012]. Until δ Spryite values are analyzed through this interval we cannot definitively rule out changes in Δ34S as a causal factor, however based upon previous Silurian

[SO4] estimates, lithologic, sequence stratigraphic, paleo-temp proxy evidence spanning this 34 34 interval we don’t expect any major changes Δ S coinciding with δ SCAS trends. In the absence of these other potential isotope fractionation drivers, the co-variation 13 34 between positive shifts in δ CCarb and δ SCAS (B.6; B.7), and the Ireviken CIE being association with an extinction event, perturbations in the sulfur and carbon cycle were likely caused by enhanced organic carbon and pyrite burial [Jeppsson, 1990; Samtleben et al., 1996; Bickert et al., 1997; Jeppsson and Calner, 2002; Calner, 2004, 2008; Cramer and Saltzman, 2005; Calner and

23 Eriksson, 2006; Lyons et al., 2009; Barrick et al., 2010;]. An increase in the fraction of sulfur to 13 the global oceans being buried as pyrite is consistent with previous interpretations of δ Ccarb records invoking enhanced organic matter burial under anoxic marine settings (Cramer and Saltzman, 2005, 2007). Futhermore, an expansion of reduced (anoxic/possibly euxinic) marine waters during eustatic sea-level rise provides a unique and plausible linkage between mass extinction event and stable isotope records during the late Llandovery-early Wenlock. Enhanced pyrite and organic carbon burial, due to increased anoxia, have also been recently documented during the late Cambrian SPICE as well (Gill et al., 2011; Saltzman et al., 2011).

5.3 Variations in I/(Ca+ Mg) Records: a Local Redox Proxy

As previously discussed in the introduction, I/(Ca+Mg) ratios are largely controlled by iodine speciation in the water column, in which, carbonates preferentially incorporate iodate into their crystal lattice over iodide [Lu et al. 2010]. Iodine speciation is highly sensitive to changes - in oxygen availability where measurements in anoxic basins have shown that iodate (IO3 ) is completely reduced to iodide (I-) nearly simultaneously with decreases in oxygen [Wong and Brewer, 1977; Wong et al., 1985; Luther and Campbell, 1991; Hardisty et al., 2014]. The main controls on iodate concentration in surface waters are due to photochemical and/or microbial iodate reduction, upwelling of iodate from well-oxygenated deep waters, and in-situ iodate production [Luther et al., 1995]. In the absence of oxygen, microbial iodate reduction would increase, upwelling of iodate would decrease, and in-situ production of iodate via biologically mediated oxidation of iodide would decrease. Therefore, depletion in local seawater oxygenation would be reflected with an overall decrease in iodate concentrations and, thus, low I/(Ca+Mg) ratios. To rule out reservoir fluxes as cause for shifts in I/(Ca+Mg) ratios, below is a brief discussion on the modern iodine cycle. Iodine’s chemical properties as a halogen and sensitivity to free oxygen conditions keeps iodine almost exclusively in the marine biosphere, in which, it makes up over 95% of the total iodine reservoir [Wong and Brewer, 1977; Whitehead, 1984; Déruelle et al., 1992]. Marine primary producers take iodate/iodide from seawater and assimilate it into their bodies [Luther et al., 1995], where it is then recycled to the water column upon decomposition. Modern iodine input fluxes (riverine, volcanic, and atmospheric) and output fluxes (biological emissions, organic carbon burial, carbonate associated iodate burial) are negligible and are 2-4 orders of

24 magnitude lower than the flux from recycled organic matter [Whitehead, 1984; Déruelle et al., 1992; Van Cappellen and Ingall, 1994; Luther et al., 1995; Milliman et al., 1995; Muramatsu et al., 2001; Wong and Cheng, 2001; Moran, 2002; O’Dowd et al., 2002; Snyder and Fehn, 2002; Lu et al., 2010]. Additionally, reservoir changes because of subduction is largely non-existent due to the volatile nature of iodine, in which, before it could be subducted, iodine would be redeposited hydrothermally as an iodine rich brine [Muramatsu et al., 2001]. So even with a residence time of 300 kyr, expected change in marine iodate/iodide reservoir over geologic time is considered minimal [Lu et al., 2010]. I/(Ca+Mg) ratios are thus invoked to reflect changes in iodine speciation due to variability in local paleoredox conditions over time.

13 13 5.4 Factors Affecting Variability δ Ccarb and δ Corg Records

13 13 Decoupled δ Ccarb and δ Corg trends have previously been recorded at the Newsom 13 Roadcut (B.7) by Cramer and Saltzman (2005) and don’t follow suit with δ Corg trends observed at the Bawny River section, Wales (B.8) or the newly documented trends from Pete Hanson Creek II section, Nevada (B.8). It has been shown that local decreases in marine phytoplankton growth rates can ultimately lead to increased primary photosynthetic fractionation leading to 13 13 decoupled δ CCarb and δ COrg curves [Freeman and Hayes, 1992; Young et al., 2008]. Decreased growth rates at the time are consistent with low nutrient availability and downwelling of WSBW [Young et al., 2008] indicative of prevailing S-state conditions (B.3). Both the Pete Hanson Creek II (B.6) and Bawny River, Wales (B.9) section show a coupled positive shift in 13 13 δ Corg values with δ CCarb (~+4‰) indicating that phytoplankton growth rates at these sections were largely unaffected by changing nutrient flux dynamics and responding to factors that influence the magnitude of primary photosynthetic fractionation (p) [Freeman and Hayes, 1992]. This is likely indicative of these deep-water sections (Nevada, Wales) being in closer proximity to a more constant supply/flux of nutrients, as opposed to the epeiric sea setting of the Newsom Roadcut which would have been largely disconnected from marine fluxes of nutrient source (i.e., sites of upwelling). With the Bawny River, Wales and Pete Hanson Creek II 13 sections, δ COrg values likely represents passive photosynthetic fractionation during the Ireviken

CIE and recorded global trends linked to increased pCO2 indicative of the Late Llandovery-Early Wenlock icehouse to greenhouse transition.

25 5.5 Redox and Environmental Conditions Across the Llandovery–Wenlock Boundary

13 34 Low δ Ccarb and δ SCAS values (0.7‰ and 21‰ respectively) observed at the base of the Pete Hanson Creek II section are representative of an overall low global burial flux of isotopically light organic matter and pyrite during the late Llandovery- Early Wenlock. Similar 13 baseline values of ~+1.0‰ in δ Ccarb prior to the Ireviken CIE are observed in age equivalent strata found in Baltic sequences as well [Cramer et al., 2010]. Increased preservation of organic carbon and pyrite likely occurred in certain anoxic coastal upwelling zones, however, due to enhanced thermohaline circulation and a oxygenated deep ocean, the overall fraction of anoxic marine settings within the global oceans during this time was still low and therefore had very 13 little to no impact on δ Ccarb values and the long-term carbon cycle (Cramer and Saltzman, 2005, 2007a,b). A more ventilated deep ocean would oxidize and remineralize both organic matter and H2S (or newly formed pyrite), resulting in circulation and delivery of isotopically 13 34 light carbon and sulfur to the global oceans and ultimately maintaining lower δ Ccarb and δ SCAS 13 values. Similarly, δ Corg values are also lower at the base of the Pete Hanson Creek II section 13 34 and seem to follow suite with observed δ CCarb and δ SCAS trends (B.6). 13 A magnitude shift of +4‰ in δ CCarb is recorded at the Pete Hanson Creek II section, 13 correlating with δ CCarb shifts previously recorded at the Newsom Roadcut by Cramer and 34 Saltzman (2005) as well as shifts globally (B.8). δ SCAS values shift positive and covaries with 13 34 δ CCarb at both field sights (B.6;7). The difference in magnitude in δ SCAS between the Pete Hanson Creek II section (+17‰; B.6) and the Newsom Roadcut (+7‰; B.7) is interpreted to be a reservoir effect due to relatively low sulfate reservoir and contributing to heterogeneity in global δ34S sulfate values (e.g., Gill et al., 2007, 2011; Jones and Fike, 2013; Young et al., 2016). 13 34 The overall covarying positive shift in δ CCarb and δ SCAS at both field sites indicates the fraction of anoxic waters in the global oceans increased exponentially during this time [Lyons et al., 2009]. As sea-level continued to rise, the continued extension of epeiric seas resulted in ieven more downwelling of WSBWs, further shifting values more positive. (B.3). The net effect of organic carbon and pyrite burial results in an overall increase in atmospheric O2 [Lyons and Gill, 2010], eventually cooling climate and reestablishing thermohaline circulation, and therefore ending both positive stable isotope excursions in δ13C and δ34S associated with the Ireviken 34 13 extinction event. The lag time seen between δ SCAS and δ CCarb following their subsequent falling limbs (B.6;7) reflects continued large scale anoxic burial of pyrite post-Ireviken CIE that

26 eventually diminished with increasing oxic conditions [Owens et al., 2013]. Additionally, it could reflect a sulfate reservoir being relatively larger than the DIC, and, therefore took longer to 34 13 respond. Peak δ SCAS values at the Newsom Roadcut section continue to occur after δ CCarb values start to return to baseline and coincident right with maximum highstand around 10m within this section (B.7). High sea-level would result an expansion of the chemocline into shallower marine settings (distal to mid-shelf), therefore, increasing the area in which MSR and pyrite burial could occur [Jones and Fike, 2013]. I/(Ca+Mg) ratios at both field sites support locally prevalent low-oxygen conditions (<2.5 umol/mol) as well [Lu et al., 2010]. I/(Ca+Mg) ratios in the Nevada section were very low throughout (B.6) and, in the absence of pervasive diagenesis, this indicates locally prevalent low oxygen conditions [Lu et al., 2010; Hardisty et al., 2017]. This is not entirely unexpected as previous studies have interpreted these strata to have been deposited in an upper slope to near shelf-slope break environments near a coastal upwelling zone [Klapper and Murphy, 1979; Berry and Murphy, 1979]. Low oxygen conditions induced by high primary production quickly reduced Iodate concentrations in the water column and lowered I/(Ca+Mg) ratios to similar values observed in oxygen minimum zones today as well as other Mesozoic oceanic anoxic 13 34 events [B.3; Lu et al., 2010, 2016]. Geochemical trends (Low δ C, δ SCAS, and I/(Ca+Mg) ratio values) in conjunction with lithostratigraphic evidence (black bedded chert) and biostratigraphic evidence in terms of timing (P. celloni conodont zone) at the base of the Pete Hanson Creek II section all fit well within the context of the P and S state model and supports the argument that a P state preceded the Ireviken CIE (B.6; B.3). Additionally, as shown by Lu et al. (2016), I/(Ca+Mg) ratios in shallow-marine carbonates can track the oxic-anoxic interface in the water column relative to the site of carbonate precipitation. This is demonstrated quite well at the Newsom Roadcut field site, a shallow inner to mid carbonate ramp, where trends in I/(Ca+Mg) ratios (B.3) are interpreted to reflect oscillations in the position of the encroaching chemocline relative to sea-level change during and after the Ireviken CIE. Although still low overall, the base of the section starts with relatively high I/(Ca+Mg) ratios that decrease in concert with rising sea-levels, and low-oxygen conditions continued to persist past maximum highstand and during the falling stand systems tract, until increasing again at the top of the section (B.3). This eventual increase in I/(Ca+Mg) ratios during falling stand likely represents a suppression of the chemocline due to falling sea-

27 levels and reestablishment, at least party, of thermohaline circulation after the Ireviken CIE and during the purported Sanda P episode (Jeppsson, 1997; Cramer and Saltzman, 2007). Unlike the Newsom Roadcut section, the Pete Hanson Creek II field site records low, unvarying I/(Ca+Mg) ratios (>0.4umol/mol) throughout the entire section (B.6) suggestive of low-oxygen conditions both during and after the Ireviken CIE. This is consistent with lithologic trends (B.6), thick sequences of organic rich lime mudstones, putting into context this sections deeper bathymetry, deposition occurring on the outer to mid ramp, and, therefore, is a closer proximity to anoxic deep waters (B.3). The fact that I/(Ca+Mg) ratios don’t increase post-Ireviken CIE (B.6) 34 coincides with high δ SCAS values and, therefore, continued pyrite burial post-CIE as discussed earlier [Owens et al., 2013].

28 CHAPTER 6

CONCLUSIONS

34 13 This study has conducted the first paired δ SCAS and δ C study along with I/(Ca+Mg) ratio measurements in the Silurian, to bring new insights into the paleoredox conditions of late 13 34 Llandovery-early Wenlock oceans. Covarying positive shifts in δ CCarb and δ SCAS coincide with the onset of sea-level rise following the end of Late Llandovery glaciation and early Silurian icehouse conditions and represent an overall increase in the fraction of anoxic waters in the global ocean, thus increasing the burial fluxes of organic matter and pyrite [Berner and Canfield, 1989; Lyons et al., 2009]. I/(Ca+Mg) ratios from our study sections support this interpretation indicating local and pervasive low oxygen conditions along parts of the western and sourthern margis of Laurentia. Shallow oxic surface waters were also shown to be in direct exchange and within proximity to anoxic water masses suggesting a shallow chemocline prior to the Ireviken extinction event and expanded during the CIE in similar pattern to the paleoredox dynamics associated with OAE2 carbon isotope excursion [Owens et al., 2017]. 13 Coeval positive shifts in δ Corg from Nevada and Wales (Loydell and Fryda, 2007) are also 13 consistent with overall expansion of anoxic conditions, as δ Corg shift positively indicating a larger flux of organic carbon being buried/preserved under expanding anoxic waters during the 13 13 early Wenlock. Although decoupled from δ Ccarb, δ Corg values at the Newsom roadcut section are consistent with a switching from initial nutrient-rich water mass with higher phytoplankton growth rates to slower growth rates as downwelling of nutrient-poor warm saline waters took over [Freeman and Hayes, 1992; Holmden et al., 1998; Cramer and Saltzman, 2007; Young et al., 2008]. The newly documented geochemical trends presented in this study broadly supports the overall oceanographic model put forward by Jeppsson (1990; 1998) demonstrating that globally marine paleoredox conditions shifted towards more reducing/less oxic conditions. This data presented provide a unique mechanism that links the Ireviken extinction event to the associated CIE via expansion of anoxic/possibly euxinic waters into shallower marine settings and increasing the burial fluxes of organic carbon and pyrite. This study demonstrates how the burial 13 flux of carbon and sulfur during P-States are both too low to cause any global shifts in δ Ccarb

29 34 and δ SCAS. Thermohaline circulation maintains a well oxygenated deep ocean oxidizing and remineralizing both organic matter and H2S (or newly formed pyrite), resulting in circulation and delivery of isotopically light carbon and sulfur to the global oceans and ultimately maintaining 13 34 lower δ Ccarb and δ SCAS (B.3). In contrast, S-states (B.3) have shown that the delivery of low oxygen surface waters combined with stagnant oceanic circulation during the Early Wenlock induced anoxia in the deep ocean and created a sink for increased organic carbon and pyrite sequestration (B.3). As eustatic sea-level continued, epeiric sea expansion exponentially increased the downwelling of oxygen-poor WSBW and an overall increase in the fraction of 13 34 anoxic waters throughout the global oceans is observed. δ Ccarb and δ SCAS shift positive due to the increase in the burial of isotopically light organic matter and pyrite in the deep oceans (B.3). Additionally, sea-level rise and sluggish thermohaline circulation eventually lead to expansion of the chemocline into the overlying water column. Eventually migrating onto shallow-continental shelf environments in which the exposure to euxinic and anoxic waters were toxic to metazoan life. The general geochemical trends during the Ireviken CIE, in addition to P to S-state 13 34 transitions, are summarized in B.9 where δ Ccarb and δ SCAS covary and shift positive and I/(Ca+Mg) decrease overall. This is similar to other proposed Paleozoic OAE’s such as the Late Cambrian SPICE Event [Gill et al., 2011] and Mesozoic such as the Cenomanian–Turonian Boundary Event (OAE2) [Schlanger and Jenkyns, 1976; Arthur et al., 1987; Owens et al., 2013; 34 13 2016; 2017]. Covarying shifts in δ SCAS and δ CCarb have each been recorded for each of these events, all of which have been found to be attributed to enhanced burial of organic matter and pyrite due to the expansion of anoxic and euxinc waters. However, unlike these other two events whose anoxia and CIEs were subsequently induced by increased primary production [Gill et al., 2007; Owens et al., 2017], the Ireviken CIE was a result of increased organic carbon burial due to the proliferation of deep ocean anoxia caused by change in deep water formation from the high latitudes to the low. Although we present some of the first paleoredox proxy evidence for expansion anoxic oceanic environments during the early Wenlock, more work is needed in other carbonate-dominated successions as well as deeper/basinal shale-dominated settings to confirm the global nature and extent of these reducing conditions during the Ireviken extinction event and CIE.

.

30 APPENDIX A

DATA TABLES

A.1 Geochemical proxy data of Newsom Roadcut field site.

13 13 18 * δ Ccarb, δ Corg,and δ Ocarb values 4 meters and above are from Cramer and Saltzman (2005, 2007).

13 18 13 34 Meterage δ Ccarb δ Ocarb δ Corg δ Scas I/Ca+Mg Formation Conodont Zone

0 -0.26 -3.83 24.9 Brassfield Fm. No Data

0.8 -0.42 -3.93 Brassfield Fm. No Data

1.6 0.05 -4.91 26.4 Brassfield Fm. No Data

2.4 0.56 -5.39 24.7 Brassfield Fm. No Data

3.2 0.93 -4.86 26.6 Brassfield Fm. No Data

4 2.4 -3.98 -28 26.3 Wayne Fm.- Maddox P. amorph Member 4.05 2.21 -4.16 Wayne Fm.- Maddox P. amorph Member 4.1 2.15 -4.05 Wayne Fm.- Maddox P. amorph Member 4.12 2.51 -3.97 Wayne Fm.- Maddox P. amorph Member 4.18 2.23 -3.25 -27.81 Wayne Fm.- Maddox K. ranulif Member 4.25 2.21 23.6 2.43 Wayne Fm.- Maddox K. ranulif Member 4.37 3.58 -3.68 Wayne Fm.- Maddox K. ranulif Member 4.5 3.42 -27.95 23.6 1.14 Wayne Fm.- Maddox K. ranulif Member 4.62 3.01 -2.18 Wayne Fm.- Maddox K. ranulif Member 4.75 -28.44 Wayne Fm.- Maddox K. ranulif Member 4.87 3.9 -3.18 Wayne Fm.- Maddox K. ranulif Member 5 -28.69 1.21 Wayne Fm.- Maddox K. ranulif Member 5.12 3.72 -3.61 Wayne Fm.- Maddox K. ranulif Member 5.37 3.92 -2.91 Wayne Fm.- Maddox K. ranulif Member 5.5 24.6 2.12 Wayne Fm.- Maddox K. ranulif Member 5.62 3.92 -3.01 Wayne Fm.- Maddox K. ranulif Member

5.75 -29.26 Wayne Fm.- Maddox K. ranulif Member 5.87 3.82 -3.35 Wayne Fm.- Maddox K. ranulif Member

6 23.9 1.63 Wayne Fm.- Maddox K. ranulif Member

31 A.1 (Continued).

13 18 13 34 Meterage δ Ccarb δ Ocarb δ Corg δ Scas I/Ca+Mg Formation Conodont Zone

6.12 3.45 -3.44 Wayne Fm.- Maddox K. ranulif Member 6.25 -28.77 Wayne Fm.- Maddox K. ranulif Member 6.37 2.61 -3.46 Wayne Fm.- Maddox K. ranulif Member 6.5 25.4 2.49 Wayne Fm.- Maddox K. ranulif Member 6.62 2.92 -3.62 Wayne Fm.- Maddox K. ranulif Member 6.75 -28.49 Wayne Fm.- Maddox K. ranulif Member 6.87 3.21 -3.58 Wayne Fm.- Maddox K. ranulif Member 7 24.1 2.11 Wayne Fm.- Maddox K. ranulif Member 7.12 3.37 -3.28 Wayne Fm.- Maddox K. ranulif Member 7.25 -27.17 Wayne Fm.- Maddox K. ranulif Member 7.37 3.15 -3.42 Wayne Fm.- Maddox K. ranulif Member 7.5 24 1.58 Wayne Fm.- Maddox K. ranulif Member 7.62 3.93 -3.22 Wayne Fm.- Maddox K. ranulif Member 7.75 -26.57 Wayne Fm.- Maddox K. ranulif Member 7.87 2.34 -3.7 Wayne Fm.- Maddox K. ranulif Member 8 24.4 1.3 Wayne Fm.- Maddox K. ranulif Member 8.12 3.33 -3.49 Wayne Fm.- Maddox K. ranulif Member 8.25 -26.01 Wayne Fm.- Maddox K. ranulif Member 8.37 3.34 -3.82 Wayne Fm.- Maddox K. ranulif Member 8.5 1.5 Wayne Fm.- Maddox K. ranulif Member 8.62 2 -3.98 Wayne Fm.- Maddox K. ranulif Member 8.75 -26.39 Wayne Fm.- Maddox K. ranulif Member 8.87 2.92 -3.59 Wayne Fm.- Maddox K. ranulif Member 9 0.9 Wayne Fm.- Maddox K. ranulif Member 9.12 2.86 -3.59 Wayne Fm.- Maddox K. ranulif Member 9.25 -25.77 Wayne Fm.- Maddox K. ranulif Member 9.37 2 -3.89 Wayne Fm.- Maddox K. ranulif Member

32 A.1 (Continued).

13 18 13 34 Meterage δ Ccarb δ Ocarb δ Corg δ Scas I/Ca+Mg Formation Conodont Zone

9.5 23.7 1.44 Wayne Fm.- Maddox K. ranulif Member 9.62 2.1 -3.76 Wayne Fm.- Maddox K. ranulif Member 9.75 -25.96 Wayne Fm.- Maddox K. ranulif Member 9.87 1.96 -3.75 Wayne Fm.- Maddox K. ranulif Member

10 26 0.93 Wayne Fm.- Maddox K. ranulif Member

10.12 1.75 -3.93 Wayne Fm.- Maddox K. ranulif Member 10.25 -25.47 Wayne Fm.- Maddox K. ranulif Member 10.37 1.48 -3.82 Wayne Fm.- Maddox K. ranulif Member 10.5 24.3 0.67 Wayne Fm.- Maddox K. ranulif Member 10.62 1.38 -3.76 Wayne Fm.- Maddox K. ranulif Member 10.75 -26.59 Wayne Fm.- Maddox K. amsdeni Member 10.87 1.2 -4.23 Wayne Fm.- Maddox K. amsdeni Member 11 21.9 0.65 Wayne Fm.- Maddox K. amsdeni Member 11.12 1.19 -3.74 Wayne Fm.- Maddox K. amsdeni Member 11.25 -25.95 Wayne Fm.- Maddox K. amsdeni Member 11.37 1.17 -4.09 Wayne Fm.- Maddox K. amsdeni Member 11.5 23.1 0.43 Wayne Fm.- Maddox K. amsdeni Member

11.62 1.08 -3.87 Wayne Fm.- Maddox K. amsdeni Member 11.75 -26.55 Wayne Fm.- Maddox K. amsdeni Member 11.87 1.22 -3.84 Wayne Fm.- Maddox K. amsdeni Member 12 22.7 0.83 Wayne Fm.- Maddox K. amsdeni Member 12.12 0.99 -4.4 Wayne Fm.- Maddox K. amsdeni Member 12.25 -26.3 Wayne Fm.- Maddox K. amsdeni Member 12.37 0.92 -4.08 Wayne Fm.- Maddox K. amsdeni Member 12.5 21.9 0.34 Wayne Fm.- Maddox K. amsdeni Member

12.62 0.96 -4.02 Wayne Fm.- Maddox K. amsdeni Member

33 A.1 (Continued).

13 18 13 34 Meterage δ Ccarb δ Ocarb δ Corg δ Scas I/Ca+Mg Formation Conodont Zone

12.75 -26.99 Wayne Fm.- Maddox K. amsdeni Member 12.87 0.98 -4.14 Wayne Fm.- Maddox K. amsdeni Member 13 21.5 0.25 Wayne Fm.- Maddox K. amsdeni Member 13.12 1.12 -3.66 Wayne Fm.- Maddox K. amsdeni Member 13.25 -26.2 Wayne Fm.- Maddox K. amsdeni Member 13.37 1.06 -3.87 Wayne Fm.- Maddox K. amsdeni Member 13.5 20 0.56 Wayne Fm.- Maddox K. amsdeni Member 13.62 0.91 -3.78 Wayne Fm.- Maddox K. amsdeni Member 13.87 1.03 -3.78 Wayne Fm.- Maddox K. amsdeni Member 14 -25.78 18.7 0.73 Wayne Fm.- Maddox K. amsdeni Member 14.12 1.01 -3.75 Wayne Fm.- Maddox K. amsdeni Member 14.37 0.96 -3.61 Wayne Fm.- Maddox K. amsdeni Member 14.5 22.9 0.1 Wayne Fm.- Maddox K. amsdeni Member 14.62 0.81 -4.27 Wayne Fm.- Maddox K. amsdeni Member 14.75 -25.65 Wayne Fm.- Maddox K. amsdeni Member 14.87 0.5 -4.28 Wayne Fm.- Maddox K. amsdeni Member 15 24 0.18 Wayne Fm.- Maddox K. amsdeni Member 15.12 0.89 -4.39 Wayne Fm.- Maddox K. amsdeni Member 15.37 0.91 -4.21 Wayne Fm.- Maddox K. amsdeni Member 15.5 -26.18 0.37 Wayne Fm.- Maddox K. amsdeni Member 15.62 1.06 -4.36 Wayne Fm.- Maddox K. amsdeni Member 15.75 0.32 Wayne Fm.- Maddox K. amsdeni Member 15.87 1.06 -4.08 Wayne Fm.- Maddox K. amsdeni Member 16 -26.49 20.7 1.12 Wayne Fm.- Maddox K. amsdeni Member

34 A.2 Geochemical proxy data of Pete Hanson Creek II field site.

13 18 13 34 METERAGE Δ CCARB Δ OCARB Δ CORG Δ SCAS I/CA+MG FORMATION CONODONT ZONE 0 0.69 -3.35 -28.6 21.1 0 Roberts P. Celloni Mountains 1.5 Roberts P. Celloni Mountains 4.5 0.19 -12.8 0.22 Roberts P. Celloni Mountains 6 1.05 -12.8 -26.3 35.3 0.09 Roberts P. Amorph Mountains 7.5 0.93 -13.2 -28.8 0.08 Roberts No Data Mountains 9 1.37 -12.2 Roberts No Data Mountains 10.5 2.65 -10.8 -27.2 0.17 Roberts No Data Mountains 12 4.23 -9.8 -26.4 0.13 Roberts No Data Mountains 13.5 3.73 -10 -23.64 0.3 Roberts No Data Mountains 15 3.62 -9.8 -26.81 0.23 Roberts No Data Mountains 16.5 4.06 -8.5 -24.4 36.4 0.24 Roberts No Data Mountains 18 3.98 -8.9 -26.5 0.15 Roberts No Data Mountains 19.5 2.64 -12 0.21 Roberts No Data Mountains 21 3.39 -9.2 -23.9 0.12 Roberts No Data Mountains 27 2.99 -13.6 38.5 0.14 Roberts No Data Mountains 28.5 3.05 -8.7 -26.7 0.21 Roberts No Data Mountains 30 3.28 -8.2 0.34 Roberts No Data Mountains 31.5 3.27 -8.2 -25.6 Roberts No Data Mountains 33 3.06 -7.3 Roberts No Data Mountains

34.5 2.65 -7.4 35.6 0.31 Roberts No Data Mountains

36 2.82 -7.1 34.5 0.24 Roberts No Data Mountains 37.5 1.71 -12 Roberts No Data Mountains 39 0.58 -8.8 -26.6 38.3 0.07 Roberts No Data Mountains

40.5 -0.18 -14.3 39.1 0.15 Roberts No Data Mountains 42 -0.02 -9.9 -28.3 33.4 0.1 Roberts No Data Mountains 43.5 -0.2 -8.1 -24.6 33.6 0.19 Roberts No Data Mountains

35 A.2 (Continued).

13 18 13 34 METERAGE Δ CCARB Δ OCARB Δ CORG Δ SCAS I/CA+MG FORMATION CONODONT ZONE

45 -0.64 -7.8 -25.8 35.3 Roberts No Data Mountains 46.5 -0.07 -3.6 0.15 Roberts No Data Mountains

36 APPENDIX B

FIGURES AND GRAPHS

13 B.1: Generalized Silurian (Wenlock-Ludlow) δ Ccarb curve with conodont and graptolite biostratigraphy, and new high precision U-Pb radiometric dates [Cramer et al., 2011, 2012, 2015].

37

B.2: Timeframe of the Ireviken Carbon Isotope Excursion in relation to Jeppsson (1990) oceanic states, epicontinental stratigraphy, glaciation, and eustatic sea level change [Cramer and Saltzman, 2007b]. Detailed biostratigraphy after Calner [2004] and eustatic sea-level curve from Cramer and Saltzman [2005].

38

B.3: Primo and Secundo state model modified after Jeppsson (1990) and Cramer and Saltzman (2005). Red star marks the Pete Hanson Creek II section and blue marks the Newsom Roadcut section. WSBW stands for Warm Saline Bottom Waters

39

B.4: Geochemical crossplots for evaluation of diagenesis for this study, please note that all 13 18 Newsom Roadcut δ Ccarb and δ Ocarb values are replotted from Cramer and Saltzman (2005, 13 18 13 34 2007). A) Plot of δ Ccarb vs. δ Ocarb. B) Plot of δ COrg vs %TOC. C) Plot of δ SCAS vs 18 18 δ Ocarb. D) Plot of CAS(ppm) vs δ Ocarb

40

B.5: Paleogeographic reconstruction of Laurentia during the middle Silurian ~431 Ma with field sites marked in their respective localities both during this time and in modern day.

41

B.6: Roberts Mountains Stratigraphic Column with stable carbon and sulfur isotopes and I/Ca+Mg data. The conodont biostratigraphy plotted is from Klapper and Murphy (1975).

42

B.7: Newsom Roadcut Stratigraphic Column with stable carbon and sulfur isotopes and I/Ca+Mg 13 13 data. Please note that δ Ccarb and δ Corg are replotted from Cramer and Saltzman (2007) and conodont biostratigraphy is from Barrick et al. (1983).

43

13 B.8: δ Ccarb isotope data of Pete Hanson Creek II, Roberts Mountains, NV, Banwy River, 13 Wales, δ Corg data from Loydell and Frýda (2007), and the Newsom Roadcut, Nashville, TN, 13 δ Corg data from Cramer and Saltzman (2007). Condont and graptolite biostraigraphy done by [Männik and Aldridge, 1989; Jeppsson, 1997; Loydell et al., 1998; Männik, 2007; Rubel et al., 2007; Cramer and Saltzman., 2006a; 2010]. Grey is due to a confusion of graptolite and conodont correlation.

44

B.9: Carbonate Redox Proxies and expected trends during the Ireviken CIE

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54 BIOGRAPHICAL SKETCH

Received a BS in geology and a minor in earth, ocean, and atmospheric science from Florida State University in December of 2014. I’m currently a graduate student at FSU receiving my master’s in geology which I started in the 2015 fall semester. I worked as an intern for the Florida Geological Survey from August of 2015 to March 2015 where I was then hired and use to work as an OPS geologist II.

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