Florida State University Libraries

Electronic Theses, Treatises and Dissertations The Graduate School

2018 A Geochemical Analysis of Rare Earth Elements Associated with Significant Sedimentary Phosphate Deposits of West- Central Florida Kyle Matthew Turner

Follow this and additional works at the DigiNole: FSU's Digital Repository. For more information, please contact [email protected] FLORIDA STATE UNIVERSITY

COLLEGE OF ARTS AND SCIENCES

A GEOCHEMICAL ANALYSIS OF RARE EARTH ELEMENTS ASSOCIATED WITH SIGNIFICANT

SEDIMENTARY PHOSPHATE DEPOSITS OF WEST-CENTRAL FLORIDA

By

KYLE MATTHEW TURNER

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

2018 Kyle M. Turner defended this thesis on April, 20, 2018. The members of the supervisory committee were:

Jeremy D. Owens Professor Directing Thesis

William C. Parker Committee Member

Vincent J. Salters Committee Member

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

ii

To my family Mom, Dad, Kayla and Bailey To my friends and colleagues Thank you for the love and support

iii ACKNOWLEDGMENTS I would like to express my thanks to my family and friends for supporting me along this journey at Florida State University. Mom, Dad, and Kayla, thank you for supporting me constantly, even when you had other things to worry about. To the faculty and staff at FSU, specifically Dr. Jeremy Owens, Dr. Vincent Salters, Dr. William Parker, and Dr. Seth Young, thank you all for supporting this project and pushing me out of my comfort zone to become a better scientist. Jeremy, you have been a blessing in my academic career, and for you I will forever be grateful. I appreciate your willingness to drop everything and discuss topics whenever necessary. To my classmates, Matthew Schreck, Eric Parrish, Randal Funderburk, Mary Beth Lupo, Ben Davis, Gary Fowler and the Young/Owens lab group, thank you for the support, scientific input and for giving me a few laughs along the way. To the geochemistry department of the National High Magnetic Field Laboratory, specifically Afi Sachi-Kocher and Gary White, thank you for assisting in the analysis of samples, and for all you do for the lab. All geochemical analysis was completed at The National High Magnetic Field Laboratory in Tallahassee, Fl. which is supported by National Science Foundation Cooperative Agreement No. DMR-1157490/DMR-1644779, and the state of Florida. To Mr. and Mrs. Bill Maxwell, thank you for financial support on this project, without it none of this would have happened. To Dr. Tom Herbert and Dr. Linda Lampl, thank you for supporting aspiring geologists, and for assisting in the logistics of this project. I would also like to thank Cascade Environmental for their generosity of using their drill rigs for sample recovery, without it, this project wouldn’t exist. Finally, Bailey Follman, thank you for the love and support over the last 5 years. Being 300 miles apart hasn’t always been easy, but our relationship has grown even stronger thanks to it. I love you.

TABLE OF CONTENTS List of Tables ...... vii

List of Figures ...... viii

Abstract ...... ix

1. INTRODUCTION ...... 1 Purpose ...... 1 Background ...... 2 Overview ...... 2 Previous Work ...... 3 Phosphorite Formation ...... 4 Florida’s Geologic History ...... 6 Sample Location and Setting ...... 7 The Hawthorn Group ...... 7 The Peace River Formation ...... 8 The Bone Valley Member ...... 8

2. METHODS ...... 10 Field Sampling ...... 10 Core Description and Sampling ...... 10 Laboratory Procedures ...... 11 Sample Digestion and Elemental Analysis ...... 11 Strontium Isotopes ...... 11 Thallium Isotopes ...... 12

3. RESULTS ...... 14 Core Description ...... 14 Strontium and Thallium Isotope Analysis ...... 14 Phosphate and Trace Elements ...... 17 Rare Earth Elements (REEs) ...... 18

4. DISCUSSION ...... 20 Geologic Age of the Peace River Formation ...... 20

v REE Trends ...... 21 Environmental Conditions and Depositional Mechanisms ...... 22 Possible Climatic Control on Phosphorite Burial ...... 26 Potential Solution ...... 29

5. CONCLUSION ...... 30

REFERENCES ...... 32

APPENDIX ...... 39

BIOGRAPHICAL SKETCH ...... 55

vi LIST OF TABLES Table 1: Sr isotope data. This data is used to plot samples on the Sr seawater curve in Figure 4...... 17

Table 2: Tl isotope data ...... 17

Table 3: Geochemical data ...... 39

vii LIST OF FIGURES Figure 1: Distribution of major United States phosphorite deposits (modified from Emsbo et al., 2015). The darker shaded green area denotes the extent of the Central Florida Phosphate District...... 3

Figure 2: Schematic showing the wide variety of environments where phosphorite formation occurs. Note that a majority of formational settings are unique to the Phanerozoic (Adapted from Pufahl and Groat, 2017)...... 5

Figure 3: (a) Paleogeographic reconstruction of the Miocene. (b) Zoomed in portion of the southeastern United States in the Miocene. (c) Locations where core recovery took place in the Central Florida Phosphate District. Modified from (Blakey, 2013) (Google Earth, 2018)...... 9

Figure 4: Concentrations of light REEs (Y, La,Ce,Pr,Nd) normalized to phosphate weight percent plotted vs. depth from surface of each sampled core. The symbols on each core correlate with symbols on Figure 5, showing locations where Sr isotope samples were taken. See Table 3 for data used to construct graph...... 15

Figure 5: Sr isotope ratios plotted on a portion of the Cenozoic Sr Seawater Curve. Note that the star represents three samples plotting with modern values while every other shape represents a single sample. Samples had a standard deviation of .000014 with the exception of HC92017 with a SD of .000066; therefore error bars are not visible. Modified from (Koepnick et al., 1988)...... 16

Figure 6: Sample concentration normalized to PAAS for REEs. Iron oxide data retrieved from (Bau and Dulski, 1995)...... 19

Figure 7: Ages of samples plotted on stratigraphic columns from Figure 4. See explanation on Fig. 4 for unit identification. Ages from the tops and bottoms of each core do not correlate with each other...... 22

Figure 8: Schematic showing inputs of P, upwelling and iron oxyhydroxide breakdown within the sediments...... 27

Figure 9: Compilation showing the changes in climate, geologic events, and phosphate deposition through the Phanerozoic, modified from (Takishima et al., 2006)...... 28

viii ABSTRACT Rare earth elements (REEs), yttrium and uranium are important industrial resources in many technology sectors; therefore, demand and production will likely continue to increase. Increased international market prices have led to new exploration for REE mineral resources in North America. It has been proposed that phosphorite deposits are a viable economic source of REE but the overall concentrations, depositional conditions, and ages are relatively unconstrained. Phosphorite is commonly associated with nutrient replete seawaters and sedimentary deposition, which is driven by upwelling, and/or continental delivery of bio-essential elements that are deposited on continental shelf regions. This project analyzed three sonic drill cores to better constrain, spatiotemporal REE concentrations of the Miocene-Pliocene aged samples from the Peace River Formation, which is associated with North America’s largest phosphate deposit. This project presents concentration data from a lateral (west-east) transect collected in central Florida that documented phosphatic sands, silts and clays. Newly obtained cores contained samples with well-rounded quartz sands, dolomitic silts, teeth, bones, and marine fossils commonly found in a near shore depositional environment. Sedimentary archives of the area are highly enriched in REEs, yielding concentrations nearing 200 ppm for some REEs. Our analysis confirms a previous study that phosphate grains, teeth, bones, and bulk sediment indicate REE are not associated with and/or sourcing from biogenic components, but rather associated with phosphate grains through secondary diagenetic processes. Together, a variety of factors could explain enrichment seen in Miocene phosphatic sediments including nutrient-rich runoff, oceanic upwelling, and sediment reworking. Although concentrations do not reach values as high as other mining sources, the relative ease of extraction from sedimentary deposits may make them a valuable source. A compilation of major global phosphate deposits through geologic history documents that a majority (based on tonnage phosphorite) of the burial are associated with climatic transitions from icehouse to greenhouse conditions.

ix CHAPTER 1 INTRODUCTION

Purpose As the global demand for Rare Earth Elements (REEs) increases, the search for new REE deposits is essential (Long et al., 2010; Emsbo et al., 2015). The central-Florida phosphate district offers a unique opportunity for recovery of REEs as ancillary elements associated with phosphorite production of phosphorus nutrient used for fertilizers (Emsbo et al., 2015). The Miocene aged Bone Valley Member, Peace River Formation of the Hawthorn Group is a pebble phosphate deposit found in portions of west-central Florida’s Polk, Hillsborough, Manatee and Hardee counties (Scott, 1988). Though minimal research has focused on the association of REEs in these central-Florida phosphorite deposits, available data suggests that REEs are present in economically valuable quantities within the Bone Valley Member (Emsbo et al., 2015). If this trend reflects the entire central Florida phosphate deposit, the region offers a significant additional source to the global demand for REEs in a relatively easy to access mineral deposit. Extrapolating findings and assuming similar enrichment mechanisms, these types of deposits could be a significant source of REEs as there are many significant sedimentary phosphorite deposits globally throughout the Phanerozoic. Rare earth elements are defined as the lanthanide series elements with atomic numbers 51-71 (La, Ce, Pr, Nd, Pm, Sm, Eu, Gd, Tb, Dy, Ho, Er, Tm, Lu) plus Y (Van Gosen et al., 2014; Bank et al., 2016). It has been long understood that REE enrichments appear in coal, and are associated with processed coal known as coal ash (Zhang et al., 2007), with current research being devoted towards finding new resources for REEs (Goodenough et al., 2016). REEs are important in that they are used in the commercial production of electronics, defense technologies, and green energy systems (Bank et al., 2016). Globally, these industries have expanded significantly causing a domestic concern for the amount of REE reserves available. Studying the underlying mechanism for the formation of enriched REEs in phosphorite deposits, better constraints on the time periods and locations of similar deposits can be constructed.

1 Background

Overview REEs can be found in a variety of geologic settings including major rift systems (European REE mettalogenetic belts; Goodenough et al., 2016), continental carbonatites (Mountain Pass Mine, California; Mariano, 2012), and deep-sea hydrothermal vents (British Geological Survey, 2011; Goodenough et al., 2014; Bank et al., 2016). Currently over 95% of REE sources and production are from igneous and metamorphic mineral deposits that are associated with secondary hydrothermal fluids (Zhi Li & Yang, 2014). Importantly, phosphorite deposits have been suggested as a solution to the increasing REE demand (Long et al., 2010; Humphries, 2010). However, a vast majority of REEs reside in alkaline igneous rocks where the extraction from these sources can be expensive and environmentally hazardous (Emsbo et al., 2015). Additionally, placer deposits often concentrate REE bearing minerals, which have been significant economical sources of REEs (Long et al., 2010; Bank et al., 2016). Consequently, placer deposits are relatively rare in the Phanerozoic and contain only trace amounts of REEs (Castor, 2008; Sengupta and Van Gosen, 2016). Thus, enriched REE sedimentary deposits present a possibly new opportunity for extraction with less energy, hazard and greater profits (Emsbo et al., 2015) but their enrichment mechanism remains unconstrained. Phosphorite deposits are abundant throughout the geologic record with numerous deposits in the United States. These deposits range from the Northwest to the Southeast (Fig. 1; Emsbo et al., 2015) with age ranges of ~450 Ma to ~2 Ma, respectively. Their distribution across the U.S. has been directly attributed to ancient sea-level variations in that phosphorite deposits are thought to be related to ancient productive continental margins (Emsbo et al., 2015). Permian phosphorite deposited in Idaho, Utah, Wyoming and Montana along with Miocene deposits of Florida, Georgia, and South Carolina could be potential REE resources (Emsbo et al., 2015) but little is known about the depositional settings for each of these deposits. There have been two studies by the United States Geologic Survey (USGS) that have assessed REE recovery for a few samples but lack spatiotemporal resolution to constrain the mechanistic underpinnings for REE enrichment. Thirty million metric tons of phosphate rock are produced annually in the US, most of

2 which is manufactured for phosphate fertilizer (Jansinski, 2011; Emsbo et al., 2015). Importantly, the Miocene Florida deposit accounts for over 85% of US production which is dominantly extracted from the mineral francolite, a calcium-rich fluorapatite (Long et al., 2010; Emsbo et al, 2015).

Figure 1: Distribution of major United States phosphorite deposits (modified from Emsbo et al., 2015). The darker shaded green area denotes the extent of the Central Florida Phosphate District.

Previous Work Phosphorite housed REEs have been analyzed using high precision mass spectrometry to provide economic identification. Previous works by USGS geologists have reported REEs concentrations from a set of samples obtained from global deposits, varying in both age and lithology (Long et al., 2010; Emsbo et al., 2015). Initial studies indicated that yttrium (Y) and lanthanum (La) are present in concentrations as high as 0.1% in Bone Valley (similar locality to this study; Long et al., 2010; Emsbo et al., 2015) suggesting that the mechanism for enrichment is similar for all phosphorite enriched sediments. Further

3 studies of francolite indicate that REEs appear to be replacing Ca in the lattice structure of the mineral, or are being adsorbed on to the surface due to weathering and post- depositional interaction with seawater (Clarke & Altschuler, 1958; Van Kauwenbergh & McClellan, 1990; Long et al., 2010; Emsbo et al., 2015). Other factors that have been suggested to control the amount of REEs present include the presence of biogenic components within the strata (Zanin & Zamirailova, 2009) and the physiochemical conditions under which phosphorites form (Emsbo et al., 2015). The Bone Valley Member contains a relatively high amount of fossil derived phosphate; therefore, biogenic components may have an impact on the fractionation of REEs (Dubinin, 2004). Previously analyzed data from Bone Valley indicate REE concentrations ranging from 10-300 ppm (Emsbo et al., 2015); however, enrichments of this magnitude lack a mechanism.

Phosphorite Formation Previous work has shown that sedimentary phosphorite deposits are directly impacted by changes in the biogeochemical cycling of P, which has changed greatly throughout Earth’s history (Figure 2; Pufahl and Groat, 2017). Additionally, there have been dynamic changes in the redox state of the ocean-atmosphere system that have affected the phosphorus cycle throughout the geologic record (Reinhard et al., 2017). Preserved sedimentary phosphorites are typically Phanerozoic in age, which suggests a relationship with biological utilization and/or the dominant redox state of the ocean- atmosphere system. Importantly, phosphorus is a bio-essential nutrient required for primary producers, thus a driver of primary productivity. The ocean source of P is dominantly via runoff of weathered continental rocks as both dissolved and particulate P (Benitez-Nelson, 2000; Filippelli, 2008; Pufahl and Groat, 2017). In oceanic settings, P regeneration from the deep ocean to the photic zone is associated with anoxic conditions due to re-mineralization of organic carbon thus the release of phosphorus (Van Cappellen and Ingall, 1994). This process allows for rapid recycling of P back into the system driving elevated concentrations in the zone of primary production (Benitez-Nelson, 2000; Filippelli, 2008; Pufahl and Groat, 2017). Thus this recycling shows the direct coupling of the P systematics to the carbon and oxygen biogeochemical cycles (Berner et al., 1993).

4

Figure 2: Schematic showing the wide variety of environments where phosphorite formation occurs. Note that a majority of formational settings are unique to the Phanerozoic (from Pufahl and Groat, 2017).

Sedimentary phosphorite deposition is controlled by two main factors, the amount of P entering the system through runoff and the degree of recycling due to coastal upwelling (Pufahl and Groat, 2017). Coastal upwelling brings nutrient replete waters, including elevated phosphorus, to the photic zone which can then lead to elevated primary production and then eventually deposition on the seafloor to eventually form phosphorite minerals. Many of the largest phosphorite deposits on earth are found near continental margins, and epeiric seaways where upwelling was prevalent (Pufahl and Groat, 2017).

Dissolved P is typically delivered in the form of HPO42-, H2PO4-, or PO43-, while particulate P, usually unavailable for biological uptake, is in the form of physically weathered igneous and metamorphic minerals (Pufahl and Groat, 2017). In some cases, particulate matter in the form of clay minerals, iron oxyhydroxides or detrital remains containing P, can be biologically utilized via dissolution and/or microbial respiration (Filippelli, 2008). Clay minerals and Fe oxyhydroxides both commonly adsorb P on their surface, hence when interaction with seawater or reduced pore water occurs; P is released (Mills et al., 2004; Drummond, 2015). Though upwelling occurs in many locations throughout the modern oceans, mainly western continental margins, upwelling itself does not guarantee phosphorite formation as dictated by modern observations. Upwelling must be present on the distal shelf and delivering P-rich waters towards the surface (Pufahl, 2010). Coastal (distal shelf) upwelling

5 typically occurs as a response to increased amounts of surficial winds pushing upper portions of the surface ocean away from land creating a ‘void space’. This can be observed today on the NW African Coast (McGregor et al., 2007), the eastern Pacific in both California (Bograd et al., 2009) and Peru (Busch and Keller, 1981) and on the western edge of the Bay of Bengal (Shetye et al., 1991). Phosphorus-rich waters in coastal regions drive productivity and sedimentary organic matter availability, in return fueling the precipitation of francolite. Examples of this feedback forming phosphorite can be observed in the Permian Phosphoria formation as well as Eocene North African (Morocco) phosphorite deposits. Francolite deposits form in large scales when pore water within the sediments becomes saturated with phosphate, producing sand sized phosphatized sediments, like those commonly seen in Miocene aged Florida phosphates (Froelich et al., 1988). Though upwelling is a driving factor in phosphorite formation, it can also form in areas with minimal amounts of upwelling, if any at all. Lack of upwelling delivering organic material commonly results in suboxic bottom waters, thus iron oxyhydroxides play a large role in phosphorite deposition (Heggie et al., 1990; Jarvis et al., 1994). Iron oxyhydroxides are dissolved when buried below the anoxic boundary, typically near/below the sediment water interface, releasing adsorbed P (Pufahl and Groat, 2017). Dissolved phosphorus concentrations then increase as a cyclical process occurs, where dissolved P is readsorbed onto Fe oxyhydroxides above the redox boundary. Then, burial occurs yet again where P is re-released into the system, eventually allowing for francolite precipitation (Pufahl and Groat, 2017).

Florida’s Geologic History The breakup of supercontinent Pangaea in the early Mesozoic (200 m.a.) caused the North American plate, including proto-Florida, to drift away (Allen and Main, 2005). A foundation carbonate platform continued to accrete through the Eocene, seen in the oldest surficial rocks in Florida (Avon Park Formation) (Lane, 1994). The Florida Georgia channel system divided present day Florida from the mainland of North America via a system of shallow waterways running east and west across the northern portion of the state (Figure 3a,b). As the Appalachians experienced uplift and erosion during the Oligocene, large amounts of sediment were transported towards the coast, filling in the channel

6 (Huddleston, 1988; Scott, 1988). Siliciclastic sediments continued southward transport, eventually depositing on top of the previously formed carbonate platform. Direct evidence of this process is present across the state in the form of carbonates containing sand lenses (Lane, 1994). Mixing of siliciclastic sediments with carbonates formed Florida’s Hawthorn Group during the late Oligocene/Miocene, also the same time phosphorite deposits accumulated (Scott, 1988). Rounded land pebble phosphatic sediments are evidence for physical weathering processes in high-energy coastal environments. Siliciclastic input from the Appalachians continued through the late Pliocene, where sediments deposited on the Florida Platform consist mostly of sands, silts, and clays rather than carbonates (Lane, 1994).

Sample Location and Setting The Central Florida Phosphate District is located inland of Tampa Bay on Florida’s west coast. Lithologies in this region consist mainly of the Bone Valley Member and Peace River Formation of the Hawthorn Group covering ~2070 square kilometers (Lane, 1994; Hurst, 2012; FDEP, 2014). This locality has been extensively mined for phosphorite since the late 1800’s, with expansion currently reaching into southern portions of the region (FDEP, 2014). Samples for this study originated from eastern portions of Manatee county and central western portions of Hardee county (Figure 3c).

The Hawthorn Group Deposition of the Hawthorn Group occurred in a marginal marine environment, with lithologies dependent on depositional age (Oligocene, Miocene, or Pliocene) and location (Long et al., 2010; Hurst, 2012). Primarily this deposit consists of carbonates with regional occurrences of dolostones, phosphatic sands, and clays (Scott, 1988). The North Florida Hawthorn Group consists predominately of carbonates as opposed to South Florida, specifically southwestern Florida, where the lower member is typically carbonates underlying a younger, more siliciclastic member (Scott, 1988). The Central Florida Phosphate District is made up of two Hawthorn Group sub-lithologies, the Peace River Formation, and the Bone Valley Member.

7 The Peace River Formation First identified by Scott in 1988, the Peace River Formation includes the uppermost portion of the Hawthorn Group in southwestern Florida (Scott, 1988). Consisting mainly of interbedded quartz sands, carbonates, and clays, this unit is easily distinguishable from other portions of the Hawthorn Group due to the vast amount of siliciclastics. Clay lenses are common within the Peace River Formation, ranging in thickness from a few centimeters to a meter, often containing phosphate grains (Scott, 1988). Silt sized dolomite (dolo-silt) and carbonate (calcilutite) are identified as two of the dominant Peace River lithologies, both containing phosphatic sediment (Scott, 1988). Along with the vast amount of micro granular phosphate present, fossils including shark’s teeth, mollusks, bivalves, and corals are present in the Peace River Formation (Scott, 1988). Additionally, chert is present in southwest Florida, especially in the lower portions of the Peace River Formation near the contact with the Arcadia Formation. Biostratigraphic framework suggests the lower boundary of the Peace River Formation indicates a late Early to late Middle Miocene age (Webb and Crissinger, 1983).

The Bone Valley Member In central Florida lies the Bone Valley Member of the Peace River Formation (Scott, 1988), with the term coming from the vast amount of recovered vertebrate fossils (Scott, 1988). The unit is identified as a member due to limited areal extent of the Bone Valley, and intertwining with other units (Matson and Clapp, 1908; Scott, 1988). Unlike the Peace River, it is strictly a clastic unit, containing large amounts of phosphorite gravels (Scott, 1988; Long et al., 2010). Phosphorite in Bone Valley can range in color from black, to brown and occasionally white in localities where groundwater leaching has occurred (Scott, 1988). Near the contact with the Arcadia Formation in lowermost portions, clay matrix percentages increase, thought to be the remains of “argillaceous carbonate rock” (Altschuler et al., 1964; Strom and Upchurch, 1983; Scott, 1988). Carbonate rip-up or rubble beds also commonly occur near the bottom of the Bone Valley, where large clasts of both carbonate and dolomite are found in a clay bed matrix (Scott, 1988). Leached phosphorite beds are thought to be associated with groundwater interaction, where most of the initial phosphate has been removed from the unit (Scott, 1988) but tracking this

8

Figure 3: (a) Paleogeographic reconstruction of the Miocene. (b) Zoomed in portion of the southeastern United States in the Miocene. (c) Locations where core recovery took place in the Central Florida Phosphate District. Modified from (Blakey, 2013) (Google Earth, 2018).

process is difficult. Fossilized remains of fauna from the lowermost portion of Bone Valley indicate a depositional age of Early to Middle Miocene (Scott, 1988). A majority of the remains found across the region, typically in the middle to upper portions, indicate a middle to late Miocene age range based on biostratigraphy which provides a general age (MacFadden and Webb, 1982; Scott, 1988). Fossilized remains located in channel fill deposits have been recovered from Bone Valley, likely due to reworking of strata throughout the Pliocene/Pleistocene (Webb and Crissinger, 1983; Scott, 1983). The wide range of Neogene aged vertebrate fossil assemblages discovered in Bone Valley have provided evidence for a Miocene-Pliocene age.

9 CHAPTER 2 METHODS

Field Sampling Sonic drilling was used to recover drill core from the Central Florida Phosphate District’s unconsolidated sediments. This technique was used to permit the highest recovery rate possible for drill core. A series of east-west borehole locations were chosen, ideally recording a transect across the region. Cores were drilled in June 2017 in locations 27°35’40”N and 81°49’19”W for HC02017, 27°35’41”N and 81°57’13”W for HC92017, and 27°35’23”N and 81°05’28”W for MC02017. All cores were collected on public land on the road right-of-way. Cores were drilled using a standard truck-mounted sonic rig, and extracted in 3mm thick plastic sleeves from the casing pipe. The MCO2017 core was drilled to a depth of 20.5 meters below the surface, with near >95% core recovery beginning at 2m. The HC92017 core was completed with a final depth of 19.8m with >95% recovery beginning at the same depth of 2m. Finally, the HC02017 core was completed with a final depth of 27.4m and a recovery of >95% beginning at 3m in depth. Samples were contained in plastic sleeves and transported in aluminum cores boxes with drain holes.

Core Description and Sampling Core materials in plastic sleeves were air-dried to remove any excess water that had been introduced to sediments during the recovery process. Prior to sampling, plastic sleeves were split open, and detailed lithological descriptions were completed. Samples were then collected for geochemical analysis. Approximately 25 samples were collected for analysis from each core for a total of 82 samples. Samples were placed in plastic weigh boats and dried in a 100°C oven overnight to remove all remaining water content. Following water removal, samples were examined under a microscope to identify and collect phosphatic grains, pebbles, teeth, and bones from the matrix and placed in 2mL Nalgene bottles. Bulk samples were extracted for analysis as well. All samples were powdered and homogenized using a trace metal clean agate mortar and pestle. Between each sample the mortar and pestle was cleaned with high purity silica sand to reduce and geochemical contamination. Once powdered, ~10mg samples were placed in acid cleaned 7mL Savillex® Teflon beakers with DI water to reduce sample loss from static.

10 Laboratory Procedures

Sample Digestion and Elemental Analysis Sample dissolution was performed in a state-of-the-art trace metal clean lab at the

Nation High Magnetic Field Laboratory (NHMFL). Trace metal clean acids of HCl, HNO3 and HF were used for a standard, heated, triple acid method for complete digestion of the

sample. Following digestion, all samples were dried down, brought up in dilute HNO3, and dried down once again prior to final dilution. Final sample dissolution was in 0.5mL of 0.5M

HNO3. Samples were then diluted in 0.4M HNO3 and a range of elements (P, Al, Ca, Ti, Mn, Fe, Sr, Y, Ba, La, Ce, Pr, Nd, Sm, Eu, Gd, Tb, Dy, Ho, Er, Tm, Yb, Lu, Tl, and U) were analyzed using an Agilent 7500cs Inductively-Coupled-Plasma Mass-Spectrometer ICP-MS using no- gas mode, housed in the Geochemistry group at the NHMFL. Upon obtaining values from the Mass Spectrometer, corrections for dilutions were calculated. Additionally, P was also measured using a sample split of the same digested solution but required a separate analytical method due to instrument settings as P ionization is only ~5%. The tuning parameters followed the most optimal setting following previous work of Wilbur et al., 2001. Measuring phosphorous concentrations using the same splits allows for direct comparison of elemental concentration to better constrain the relationship with phosphorite material. Additionally, the measurement of P can be difficult due to interferences of S and O with minor mass overlap and the very low ionization efficiency of

P (~5%). Importantly, these samples appear to have very low sulfur (S32) contents and the tune reduces oxide interferences including molecular oxygen. Briefly, the torch sample depth was 11mm, allowing for concentrations of P to be obtained due to increased

ionization. For comparison to previous data the elemental P contents are reported as P2O5 wt.%, which is calculated using stoichiometric conversion which assumes all P was deposited as phosphorite mineral likely as francolite. A USGS SCO-2 standard was run with both methods and elements are generally within 10% of accepted values with the exception of 3 elements; Ca (25.27% error), La (19.91% error), and Lu (21.20% error).

Strontium Isotopes A split of previously dissolved total digest samples were used for Sr isotope analysis. Samples were dried down and re-dissolved in 0.25mL of 2.5N HCl in acid-cleaned 15mL

11 Teflon Savillex® beakers. Quartz cation exchange columns with 200-400 mesh Dowex 50 resin were backwashed with 10mL of 2.5N HCl before samples were loaded. Samples were loaded onto the resin with 0.25mL of 2.5N HCl, followed by 3 rinses, each consisting of the same 0.25mL of 2.5N HCl. Next, the cation exchange column was rinsed with18mL of 2.5N HCl. Subsequently, the Sr fraction was collected in 5.5mL of 2.5N HCl. The Sr fraction with

dried down with 50 4) to remove organics and three consecutive

dry-downs after additionμL of percloricof 50 acid (HClO 3 3 was added to each sample and allowedμL of to concentrated dissolve completely. HNO . Finally, At this 3μL point, of 0.25MSr was of ready HNO to be loaded onto filaments used for analysis. Savillex® beakers containing Sr were taken from the trace metal clean lab to a clean flow bench outside of the lab, where filaments were loaded. Tungsten (W) filaments were constructed and allowed to degas for 4 hours at 3A and at 3.8A for 30 minutes. Samples

3, and Ta phosphate indicator solution. wereSamples loaded were with flashed 1μL untilof solution they glowed containing red, andSr, HNO a white residue was left on the center of the filament. Filaments containing sample were then loaded into a Thermo Finnigan MAT 262 TIMS for isotope analysis. Samples were reported relative to the E&A standard with an accepted value of 87Sr/86Sr=0.708000 and corrected for mass bias using 86Sr/88Sr=0.1194.

Thallium Isotopes Additional sample splits of previously dissolved sample were dried down and dissolved in 0.5 ml of 1 N HCl. Thallium purification was done using established ion exchange chromatography protocols (Rehkämper and Halliday 1999, Nielsen et al., 2004 Owens et al., 2017) where most samples only used one column. This procedure has shown to be effective for samples with elevated Tl compared to Pb (Ostrander et al., 2017). Briefly, samples were fully oxidized using a brominated HCl solution the night before anion column chemistry. Micro columns were loaded with 100 μl of AG1X8 anion exchange resin, which

was cleaned using with 1.5mL of 0.1M HCl + 5% SO2. Next, 1.6mL of 0.1M HCl was allowed to pass through the column, followed by the addition of 1mL of weakly brominated HCl to the resin in order to prepare the resin prior to sample addition. Samples were loaded into

micro-columns and a series of weak brominated HNO3 acid rinses were allowed to pass through the columns, removing lead (Pb) from the samples. Purified sample was collected

12 through addition of 0.1M HCl + 5% SO2. Tl isotope fractions were collected in a trace metal clean acid washed Savillex® beaker. Samples were then tr

1:1 HCl:HNO3 eated with 50μL of concentrated HNO3 was thenand placed allowed in each to sit beaker for 3 hoursand dried before down being immediately. dried down. All 50μL samples of concentrated were brought u 0.1 HNO3 + 0.1% H2SO4. A small aliquot of solution was

used to makep to aa 30volume-fold dilution of 500μL for in each sample and geostandard to determine Tl and Pb concentrations using an Agilent 7500cs ICP-MS. Once confirming Pb concentrations were low, samples were spiked with NIST SRM 981 Pb solution in order to correct for mass bias when measuring Tl isotopic ratios. Tl isotopic ratios were analyzed using a Thermo Scientific Neptune Multi-Collector ICP-MS, using equation 1.

205Tl = [(205Tl/203Tlsample – 205Tl/203TlNIST SRM 997)/(205Tl/203TlNIST SRM 997)] x 10,000

(1.) ε

13 CHAPTER 3 RESULTS

Core Description All three cores recovered from the Central Florida Phosphate District have distinct contacts between Quaternary alluvium and the Peace River Formation, with contact depth varying from 4.6 m for HC02017 to 8.5 m for MC02017. Refer to Figure 4 for stratigraphic columns displaying lithological variation with depth. Stratigraphically below the contact appears the Peace River Formation of the Hawthorn Group, which continues the entire length of each core. Deepest portions of HC02017 indicate a nearing of the contact with the Arcadia Formation of the Hawthorn Group, however a well-defined contact is not visible. Both MC02017 and HC92017 show lithologies consisting of sandy clay, dolomitic silt, and clay below the Quaternary alluvium contact, with a carbonate rip-up sequence occurring in both cores at a depth of ~17m below land surface (BLS). The rip-up sequence is composed of shell fragments, shark’s teeth, corals, and coble sized dolomite clasts in a carbonate/clay matrix, with a stratigraphic thickness of ~2m. Below the rip-up sequence, lithologies are composed mainly of dolomitic silt containing both micro granular and pebble sized phosphate varying in color from black to white. Both MC02017 and HC92017 have phosphate throughout all lithologies included in the Peace River Formation. Toward the base of HC92017 sediments are composed of fossiliferous calcilutite, which happens to be present in the bottom portions of HC02017 as well. Core HC02017 contains many of the same units present within the other two cores, however at a depth of ~18m, chert nodules appear within the dolomitic silt, transitioning to calcilutite for the remainder of the length of the core. As previously mentioned, the stratigraphically lowest unit within HC02017 is a fossiliferous calcilutite like HC92017.

Strontium and Thallium Isotope Analysis Strontium isotope ratios (87Sr/86Sr) were obtained for samples collected from the top and bottom of each of the three cores. A total of 6 core samples were analyzed plus two E & A elemental standards. Both analyses of the E & A standard were within error of the accepted value of 0.708000. The top of MC02017 (1-08) and HC92017 (2-06) both yielded

14

Figure 4: Concentrations of light REEs (Y, La,Ce,Pr,Nd) normalized to phosphate weight percent plotted vs. depth from surface of each sampled core. The symbols on each core correlate with symbols on Figure 5, showing locations where Sr isotope samples were taken. See Table 3 for data used to construct graph.

15 values within error of modern seawater of 0.7092. The bottom of HC92017 (2-33) was also within error of modern seawater albeit with a somewhat larger uncertainty. The bottom of MC02017 (1-10) yielded a value of 0.70882, while the top of HC02017 (3-03) yielded a similar value of 0.70876. Finally the bottom of HC02017 (3-30) yielded a value of 0.70834, being the lowest isotopic ratio measured from the collected core (Figure 5). See Table 1 for individual data. Thallium isotopic ratios were also obtained. The minimum 205Tl value was

205 determined to be -10.5, with the maximum value being -4.8. The average ε Tl value was calculated to be -7.9 (Table 2). ε

Figure 5: Sr isotope ratios plotted on a portion of the Cenozoic Sr Seawater Curve. Note that the star represents three samples plotting with modern values while every other shape represents a single sample. Samples had a standard deviation of .000014 with the exception of HC92017 with a SD of .000066; therefore error bars are not visible. Modified from (Koepnick et al., 1988).

16 Table 1: Sr isotope data. This data is used to plot samples on the Sr seawater curve in Figure 4. Core Depth (m) 87Sr/86Sr Runs E & A STD - 0.708001±7 108 E & A STD - 0.708003±7 120 MC02017 8.7 0.709258±7 115 MC02017 17.7 0.708826±8 155 HC92017 7.9 0.709223±7 80 HC92017 19.5 0.709239±33 55 HC02017 4.9 0.708762±7 145 HC02017 26.5 0.708343±7 107

Table 2: Tl isotope data 205 Core Depth (m) Type ε Tl 2σ n MC02017 9.8 Grain -10.4 0.5 3

MC02017 13.1 PO4 Clay -10.5 0.7 3

MC02017 13.4 PO4 Clay -7.5 0.5 3 MC02017 14.6 Grain -9.4 0.5 2 MC02017 16.2 Bulk -8.8 0.2 3 HC92017 9.8 Grain -7.2 0.0 2 HC92017 10.4 Bulk -6.7 0.2 2 HC92017 14.6 Grain -6.4 0.1 2 HC92017 19.5 Bulk -4.8 0.2 2 HC02017 6.4 Grain -7.4 0.6 2

Phosphate and Trace Elements

Phosphate concentrations (P2O5) of each sample varied widely from 2.43 to 38.42 wt.%, which was expected as some samples were isolated grains of phosphate while others were bulk samples. The average phosphate concentration was measured to be 24.79 wt.% among all samples (Table 3). Iron values reached a maximum of 3.36 wt.% for a single sample, however most samples yielded values under 0.80 wt.%. Bulk samples had the

17 highest Fe wt.% values, while granular phosphate samples had the lowest values, with a few yielding below detection limit (BDL) values. Concentrations of Mn reached a maximum of 413.3 ppm with an average for all samples of 102.8 ppm. Aluminum (Al) values were also reported as elemental wt.% with an average of 0.48 wt.%. Calcium (Ca) yielded the highest elemental wt.% values with a maximum of 54.23 wt.%, averaging 14.18 wt.%, due to the fact that francolite is a calcium phosphate. Finally, uranium (U) concentrations reached a maximum of 563.1 ppm, while averaging 160.2 ppm.

Rare Earth Elements (REEs) Concentrations of light rare earth elements (LREEs; which includes elements between La and Eu on the periodic table) yielded a wide range of values (Table 3). It must

be noted that values obtained are total amounts thus not normalized to P2O5 weight percent. Lanthanum (La) yielded a maximum of 132.4 ppm, averaging 35.0 ppm among all samples. Cerium (Ce) yielded a maximum of 178.1 ppm, and an average of 47.2 ppm, while praseodymium (Pm) yielded a maximum of 26.0 ppm, and an average of 6.5 ppm. Neodymium (Nd) values nearly mirrored those for La, with a maximum of 117.7 ppm, and an average of 33.1 ppm. Samarium (Sm) yielded a maximum of 19.7 ppm, averaging 5.1 ppm. Finally, Europium (Eu) yielded a maximum of 4.8 ppm, averaging 1.3 ppm among all samples. Minimum concentrations for LREEs were nearing zero indicating they were present in very minute amounts in some samples. LREE concentrations varied with depth, though a definitive sorting mechanism did not appear, showing random periods of sediment enriched in LREEs (Figure 4). When plotted on a spider diagram, raw sample data, and data normalized to P2O5 wt.%, shows samples are enriched in LREEs compared to other REEs, however, when normalized to Post Archean Average Shale (PAAS), biogenic samples are depleted in LREEs, while phosphatic grain samples are enriched by half an order of magnitude (Figure 6). Though heavy rare earth element (HREEs) concentrations do not reach values as high as LREEs (excluding Yttrium), they are enriched relative to PAAS among all samples (Figure 6). Yttrium (Y) is considered a HREE, and happened to yield the highest concentrations. Y yielded a maximum of 350.5 ppm, with an average of 63.2 ppm. The only samples depleted in Y when normalized to PAAS were samples sourcing from clay.

18 Gadolinium (Gd) yielded a maximum of 24.6 ppm, with an average of 6.5 ppm. Terbium (Tb) yielded a maximum of 3.9 ppm, and an average of 1.0 ppm. Dysprosium (Dy) yielded a maximum of 25.4 ppm, and an average of 6.2 ppm. Holmium (Ho) yielded a maximum of 7.5 ppm, and an average of 1.5 ppm. Erbium (Er) yielded a maximum of 33.8 ppm, and an average of 4.9 ppm. Thulium (Tm) yielded a maximum of 5.0 ppm, and an average of 0.7 ppm. Ytterbium (Yb) yielded values reflecting those for Er, a maximum of 41.8 ppm, averaging 4.6 ppm. Finally, Lutetium (Lu) yielded a maximum of 6.6 ppm, and an average of 0.7 ppm. Normalized (PAAS) HREE values for phosphatic grain samples neared single order of magnitude enrichment, while bulk, clay, and tooth samples contained near the same amount of HREEs as average shale.

Figure 6: Sample concentration normalized to PAAS for REEs. Iron oxide data retrieved from (Bau and Dulski, 1995).

19 CHAPTER 4 DISCUSSION

Geologic Age of the Peace River Formation Fossils from the Peace River Formation suggest a late to early middle Miocene age, which has been approximated to be 15-20 Ma (Webb and Crissinger, 1983; Scott, 1988). Strontium isotope data provided here suggests an age of (~0-21 Ma) which was estimated due to the steep inflection of the Cenozoic seawater isotope curve during the Miocene (see Fig. 5; Koepnick et al., 1988). This provides the first reported geochemical evidence for the age of deposition for these samples, which are in general agreement. However, periods of reworking among sediments in the coastal regions of the Florida platform support a wide range of ages for lithological units in central to south Florida. Three of the six Sr isotope ratio analyses (MC02017-8.7m, HC92017-7.9m, and HC92017-19.5m) suggest an age of near modern when compared to the known Sr isotope seawater curve (within the last million years; 87Sr/86Sr close to 0.7092). Interestingly, the lowermost sediment from HC92017, a depth of 19.5m, produced a value indicating the same depositional age as uppermost portions. The standard deviation of this sample was found to be 0.000066, but this still suggests near modern deposition, likely within the last few million years. This result was unexpected but it is possible that 20m of sediment was deposited in ~1 million years and it is likely this core experienced times of limited sediment deposition. Additionally, the value from the surface could have been reset due to reworking or other fluid flows including meteoric waters. Ages suggest that portions of Florida not only experienced serious reworking events, but channel filling events, meaning younger sediments could have been re-deposited alongside sediment that are much older. This Sr isotope dataset provides a low resolution geochemical estimate of depositional age, but it is unlikely that higher resolution will provide a clearer picture as the cores have been reworked numerous times and age estimates are difficult with this method. The deepest sample from core MC02017 (green square Fig. 4, 5), 17.7m in depth from land surface, produced an age of ~11 Ma, placing the period of deposition in the late middle Miocene. It must be noted that this sample had a standard deviation within the accepted range. The HC02017 core top and bottom produced ages older than all other

20 samples for which 87Sr/86Sr ratios were obtained. The top of the core, depth of 4.9m, produced an early middle Miocene age (~14Ma), while the bottom of the core was found to be early Miocene aged (~20Ma). Both samples had the lowest possible reported value for standard deviation (Figure 5). The approximate age for the upper portion of the core might indicate that younger deposited sediment was removed from this location. This is particularly likely as the core was drilled alongside the road, adjacent to both active and inactive phosphate mines. While all ages obtained from Sr isotope are supported by biostratigraphic ages, the age variation was unexpected as a total distance of only 13.5km separated the cores. Thus, based on Sr isotopes, the three cores do not overlap with respect to time and they do not create a single continuous stratigraphic record (Figure 7). This is likely due to the extended periods of winnowing in a coastal environment, as well as the possibility of slow sedimentation rates. From data collected, we can confirm that the Peace River Formation is early Miocene to Quaternary in age.

REE Trends Samples were classified into two groups of samples for which REE analysis were preformed. Groups were defined as biogenic samples (teeth and bones) and non-biogenic

samples (bulk, clay, and P2O5 grains) as they yielded significantly different magnitudes and trends for REE concentrations when normalized to PAAS (Figure 6). Normalized LREEs concentrations ranged from half order of magnitude enrichment to half order of magnitude depletion. The type of sample dictated the concentration magnitude; non-biogenic samples were enriched while biogenic samples were depleted in LREEs. Sample type had a direct impact of the enrichment or depletion of concentrations for HREEs. Phosphatic grain samples were enriched in HREEs when normalized to PAAS, ranging from a half to single order of magnitude enrichment. Bulk samples were just slightly enriched in HREEs, while both clay and teeth samples were depleted by approximately a half order of magnitude. Interestingly, the REE patterns are similar to those observed for Fe oxides but the magnitudes of enrichment for phosphorite samples are greater (~1.5 orders of magnitude).

21

Figure 7: Ages of samples plotted on stratigraphic columns from Figure 4. See explanation on Fig. 4 for unit identification. Ages from the tops and bottoms of each core do not correlate with each other.

Environmental Conditions and Depositional Mechanisms Understanding the mechanism for these larger REE enrichments, exceeding one order of magnitude, is crucial when discussing the phosphorite formation and depositional system. Typical shale (PAAS) contains concentrations of LREEs ranging from 1.1 ppm to 80.0 ppm, while HREEs range from 0.4 ppm to 4.7 ppm (Taylor and McLennan, 1985). In samples from the Peace River Formation, these values are much greater; therefore, an

22 additional enrichment mechanism is required. Data collected suggests that upon

completion of phosphorite formation, REE enrichment occurs, accumulating REEs in P2O5 enriched samples. Three factors are likely involved in phosphorite formation in this depositional setting: (1) dissolved nutrient delivery from continental runoff, (2) upwelling of nutrient rich water to the distal shelf region, and (3) recycling of sediment at the sediment water interface. The formation of Florida and its surrounding platforms began during the Miocene, when siliciclastic sediments filled the Florida/Georgia channel system. This closure allowed for weathered material from the Appalachian mountain system to advance down the Florida carbonate platform (Scott, 1988). It is likely that siliciclastic sediments from igneous rocks of the southern Appalachians were weathered and delivered to the Gulf of Mexico, both as dissolved and particulate matter. This delivery of dissolved and siliciclastic material could potentially drive increased productivity by delivering bio-essential

nutrients to the ocean system. This is supported by the presence of rutile (TiO2) that was observed in the HC92017 core from a portion ranging 9.5m -12m in depth. However, conservative elements such as Fe, Al and Tl are all below typical crustal values, which suggests there is limited delivery of siliciclastic material and visual evidence corroborates this idea (see below for further discussion). Nutrients are typically recycled within the shelf region, though some are deposited further offshore where they eventually fall through the water column into deeper oceanic settings. Primary producers utilize bio-essential elements dissolved in riverine runoff first, while solid phases require additional processes to break down and release nutrients to the water column (Pufahl and Groat, 2017). Evidence of local upper trophic level productivity is found within samples from Florida Miocene sediments in the form of fossilized sharks teeth, corals, and bivalves. However, samples from the Peace River Formation show very little visual evidence of local organic carbon and/or pyrite burial. A second mechanism, upwelling of nutrient-rich waters, has been suggested as a possible delivery mechanism of P and REEs to the shelf region (Fig. 8). Importantly, upwelling waters recycle surficially exported nutrients back to the photic zone. Lastly, recycling of bio-essential elements within the water column, and at the sediment water interface, has been proposed as a mechanism to continue enhanced primary production. For P, this process can be enhanced under anoxic conditions (Van

23 Cappellen and Ingall, 1994); however, there is no known evidence for major organic carbon and/or anoxic conditions in this region during the Miocene. Additionally, there is a loss of P from the water column due to phosphorite precipitation and burial. Therefore, recycling of P alone cannot account for the long-term burial of phosphate, and must invoke one and/or both P delivery mechanisms. Thus, better constraining the depositional environment can provide insight into the possible formation of phosphorite minerals and the associated REE enrichments. All samples document a significant depletion of Ce concentrations (Fig. 6), which is the most obvious REE trend. The normalized Ce values document a negative anomaly for all samples. This suggests deposition in an oxygen-rich water column (Elderfield and Greaves, 1981). Cerium has two valance states (Ce3+ and Ce4+), therefore a negative anomaly commonly occurs in seawater due to oxidation of soluble Ce3+ to Ce4+ (Pattan et al., 2005). Negative Ce anomalies have been documented for many sample and mineral types including phosphorites and iron oxides (Pattan et al., 2005). It has also been documented that this negative anomaly is present for samples deposited under slightly reducing conditions, which could reflect initial conditions in the upper water -column (Pattan et al., 2005). Surprisingly, samples from the Peace River Formation do not yield elevated Fe and/or Mn concentrations which are expected under oxic burial conditions (Table 3). Instead, Fe and Mn concentrations are well below crustal averages (Taylor, 1964), indicating a loss of both elements due to re-mineralization within the sediments. It is common for Mn depletion to occur in the presence of oxygen poor water column and/or sediments due to biological reduction of Mn oxides (Burdige, 2007). Fe oxides are also typically lost in such environments, but average detrital Fe concentrations are relatively high and much more difficult to assess a loss of oxide minerals due to a large detrital flux. Given the low Fe values coupled with below continental crust average Al and Ti values (Taylor, 1964), it is suggested that these samples did not accumulate much siliciclastic material. Coupling the negative Ce anomaly with low Fe and Mn oxide concentrations suggests the water-column was oxygenated, while the sediments and pore fluids were reducing enough to consume Fe and Mn oxides. However, were not reducing enough to preserve organic carbon or pyrite in these sediments. This was likely aided by the

24 continued and several reworking events that lead to a progressive loss of organic carbon. This could be similar to modern sediments from the Amazon basin which do not accumulate pyrite even with reducing pore fluids and sulfide due to bioturbation (Aller et al., 1996). One possible explanation for the observed REE data in phosphorites is that the original source is from the delivery of ferromanganese oxides. Though iron oxides are not enriched in REEs to the same degree as the phosphatic samples (Bau and Dulski, 1995), the same overall trends are present. Average iron oxide data shows enrichment in Y relative to LREEs (Bau and Dulski, 1995), as well as showing a higher enrichment factor for HREEs to LREEs (Fig. 6). Additionally, there is a significant amount of P associated with ferromanganese oxides, thus this could be the initial source of P and REEs that eventually form phosphorite minerals. It has been proposed that phosphorite formation occurs at or near the sediment water interface via precipitation and recycling that requires variable redox conditions (Pufahl and Groat, 2017). However, REE enrichment directly from seawater seems unlikely due to the low seawater concentration (Anthoni, 2000, 2006). Therefore, it is more likely that a secondary enrichment is occurring in the form of Fe oxyhydroxide recycling. Though there is negative Ce anomaly and the presence of fossils that are indicative of an oxygenated environment, the lack of visible organic carbon, pyrite and vanishingly low Mn concentrations indicate a redox boundary shift, likely at or near the sediment water interface (Froelich 1989). This hypothesis proposes that the presence of reducing pore fluids due to organic carbon consumption would dissolve Fe and Mn oxides, releasing Fe and Mn as dissolved species which would also release adsorbed P, and REEs back into pore fluids which due to seawater mixing could have a loss of elements. This release could provide the required P for phosphorite precipitation, which could then begin adsorbing associated REEs from the pore fluids. However, PAAS normalized Fe oxide enrichments are minimal compared to the values recorded in these phosphorite deposits. Thus, this requires continued delivery and formation that would allow for enhanced enrichments and when coupled with known reworking of the sediments would allow for continued enrichment of REE but loss of Fe and Mn oxides to the water column. Lastly, the Tl isotope data suggests the values are near seawater for this time period (Nielsen et al, 2004) and not directly from Mn oxides. This

25 suggests there was either limited Mn oxide delivery to these sediments or the Tl associated with phosphorite precipitation occurs near the sediment water interface allowing for an incorporation of seawater values. Interpreting these many lines of evidence for phosphorite formation suggest that there was a reduction of oxygen in the upper few centimeters of the sediments. Importantly, there needs to be at least one climatic process to deliver the required bio-essential nutrients, including P, to initiate and maintain the large phosphorite burial. On a large scale, this could be related to the overall climatic state of the oceanic/atmospheric systems. Either mechanism, upwelling or weathering, could be important during the Miocene as there is major glaciation which affects both ocean circulation and weathering rates. Importantly, there is at least one other Miocene phosphorite deposit, the Monterey Formation of California (Filippelli et al., 1994). Florida deposits could be driven by either upwelling from a strengthened Gulf of Mexico loop current or additional weathering of the Appalachian region, ultimately driven by significant glaciation. Miocene glaciation (Mercer and Sutter, 1982) allowed for a cooler planet, likely causing the Gulf of Mexico loop current to strengthen, allowing for upwelling to occur in areas adjacent to coastal Florida. Sea level models for the Miocene also indicate a ~200m fall in the upper middle to late portions of the period (Haupt and Seidov, 2012). This regression, linked to glaciation, would help drive upwelling on distal portions of the Florida carbonate platform. Generally, the climate state may drive possible global phosphorite formation but it is likely that the overall mechanisms for local deposits vary.

Possible Climatic Control on Phosphorite Burial Phosphorite deposits are found throughout the geologic record, with age ranges from Pre-Cambrian to modern day (Devoto and Stephens, 1979; Van Kauwenbergh, 2010) however, the magnitude, types, and depositional settings of these deposits vary (igneous vs. sedimentary). Not all phosphorite deposits may have enriched REEs due to the complexity of the suggested phosphorite depositional mechanisms. The available data of age, P content, and volume can help calculate a tonnage for a given time period (Devoto and Stephens, 1979; Van Kauwenbergh, 2010). The samples were binned into 13 geologic periods within the Phanerozoic. The Paleogene has the largest phosphorite deposits in the

26

Figure 8: Schematic showing inputs of P, upwelling and iron oxyhydroxide breakdown within the sediments.

world, with Morocco accounting for 83% of the 68,000 million metric tons (mmt) deposited in the Paleogene. Neogene phosphorite deposits make up the second largest amount of global deposits, with nearly 21,000 mmt believed to have been deposited in this era (Chernoff and Orris, 2002). The Cretaceous (7,500 mmt), Triassic (5,700 mmt), Permian (9,600 mmt), Carboniferous (1,100 mmt), Ordovician (6,500 mmt), Cambrian

(4,000 mmt), and Pre-Cambrian (10,500mmt) make up the remains amounts of global P2O5 deposits (Chernoff and Orris, 2002; Van Kauwenbergh, 2010). Pre-Cambrian, Triassic, and Neogene phosphorite deposits were deposited during icehouse conditions on the planet; meaning glaciation was present (Figure 9). These three periods of deposition account for

28% of total global P2O5 deposits, and when Paleogene deposits are added to this total,

account for nearly 79% of total P2O5 volume. Thus, there appears to be a unique correlation between glaciation and phosphorite deposition. The majority, (79%), of global deposits are associated with the transition into or out of major glaciations. Glaciation on the poles would not only lower global sea level due to enhanced continental ice in the high latitudes, but likely drives enhanced physical weathering and stronger upwelling and thermohaline circulation due to strong equator to pole temperature gradients (Anderson, 2005). Both of these would provide increased delivery of nutrients to fuel enhanced primary production likely driving redox variability at the sediment water interface, which has been suggested as a necessary requirement for

27 phosphogenesis, which our data seems to confirm. This may have not been the case for Pre- Cambrian phosphorite deposits in that most of these phosphorite deposits are actually carbonatites (igneous origin).

Figure 9: Compilation showing the changes in climate, geologic events, and phosphate deposition through the Phanerozoic, modified from (Takishima et al., 2006).

28 Potential Solution The US produces nearly 30 million metric tons of phosphate material annually; even a small enrichment in REEs could yield impactful amounts in the form of ancillary mineral

recovery alongside of P2O5 production. Florida produces approximately 80% of total phosphate in the United States, an amount equal to 24,000,000 metric tons yr-1 (Jasinski, 2016). Samples from the central Florida phosphate district contain between 0.1 and 0.01 wt. % total REEs depending on sample type (biogenic vs. non-biogenic). An average total REEs wt. % of 0.03 for phosphatic grain samples was determined from this study and used to calculate a total yearly yield assuming homogenous distribution all phosphorites mined. Thus, 24,000,000 metric tons of Florida phosphate could yield ~7,200 metric tons of total REEs assuming annual production numbers are calculated based on phosphatic material, not bulk sediments. Florida phosphates could significantly impact the global supply despite their limited resource size. Importantly, global demand in 2016 for REEs exceeds 95,000 metric tons (Zhou et al., 2017), meaning that extraction from Florida phosphorites containing 0.03wt. % total REEs could produce ~7.6% of total REE global demand. This is especially important as the REEs are likely being dissolved already thus this capture of REEs should be relatively minimal.

29 CHAPTER 5 CONCLUSION The Central Florida Phosphate District is a unique sedimentary phosphorite deposit in that present location does not place it near an area of continuous deep ocean upwelling, or an epeiric seaway. Determined to be early Miocene to Quaternary in age, the deposit is composed of Hawthorn Group sediments, including both the Peace River Formation and the Bone Valley Member. Localities in the southern portion of the district are comprised mostly of Peace River Formation, with very little Bone Valley Member present. Phosphorite from this locality contains an elevated amount of REEs, with LREEs enrichment reaching a full order of magnitude, and HREEs enrichment nearing two orders of magnitude relative to PAAS. These REE enrichments could be attributed to the process forming Florida phosphorites. Miocene sea levels fluctuated as much as 200m, meaning specific depositional locations varied greatly in water depth throughout the epoch. With such drastic changes in sea level, upwelling of nutrient rich deep ocean waters would have allowed for phosphogenesis to take place on the shelf regions adjacent to the present day Florida coastline. With sea level fluctuation attributed to glaciation, there appears to be a correlation between the world’s largest phosphorite deposits and the decent from icehouse into greenhouse conditions. Glaciation would allow from increased amounts of nutrient entering the ocean, which in return would increase the amount of amount nutrients to be upwelled to shelf regions for phosphogenesis to occur. In order to better predict the value of global phosphorite deposits, a better understanding of the formation mechanism controlling francolite precipitation is necessary. With iron oxyhydroxides scouring the ocean floor adsorbing both phosphorus and REEs, notable concentrations of these elements are present in the initial precipitation of francolite. As recycling of Fe oxides occurs near the sediment water interface, Fe is dissolved back into seawater, allowing for P and REEs to accumulate, eventually reaching levels seen in the Central Florida Phosphate District. Though influenced by storm events,

sea level transgression and regression, and mechanical reworking, P2O5 levels remain elevated while Fe concentrations do not increase. By better understanding the formation of sedimentary phosphorites, the exploitation of these deposits as a possible source of REEs

30 can be evaluated to a greater extent. Assuming homogenous concentrations of REEs among Florida phosphates, deposits may offer a solution to the increasing global demand for these elements, and providing extraction is efficient with high amounts of recovery, phosphorites from Florida may contain up to 7.6% of the global annual demand.

31 REFERENCES 1. Allen, G. M., & Main, M. B. (2005). Florida's Geological History1. Florida Cooperative Extension Service.

2. Aller, R. C., Blair, N. E., Xia, Q., & Rude, P. D. (1996). Remineralization rates, recycling, and storage of carbon in Amazon shelf sediments. Continental Shelf Research, 16(5-6), 753-786.

3. Altschuler, Z. S., Cathcart, J. B., & Young, E. J. (1964). Geology and geochemistry of the Bone Valley Formation and its phosphate deposits, west central Florida (Vol. 6). Geological Society of America.

4. Anderson, S. P. (2005). Glaciers show direct linkage between erosion rate and chemical weathering fluxes. Geomorphology, 67(1-2), 147-157.

5. Anthoni, J. F. (2000, 2006). The Chemical Composition of Seawater. Retrieved April 2017, from Seafriends: www.seafriends.org.nz/oceano/seawater.htm

6. Bank, T., Roth, E., Tinker, P., & Granite, E. (2016). Analysis of Rare Earth Elements in Geologic Samples using Inductively Coupled Plasma Mass Spectrometry. Pittsburgh: US Department of Energy .

7. Bau, M., & Dulski, P. (1995). Comparative study of yttrium and rare-earth element behaviors in fluorine-rich hydrothermal fluids. Contributions to Mineralogy and Petrology, 119(2-3), 213-223.

8. Benitez-Nelson, C. R., (2000). The biogeochemical cycling of phosphorus in marine systems. Earth-Science Reviews, v.51, p. 109-135.

9. Berner, R. A., Ruttenberg, K. C., Ingall, E. D., & Rao, J. L. (1993). The nature of phosphorus burial in modern marine sediments. In Interactions of C, N, P and S Biogeochemical Cycles and Global Change (pp. 365-378). Springer, Berlin, Heidelberg.

10. Blakey, R. (2013). Deep Time Maps: maps of ancient earth. https://deeptimemaps.com/

11. Bograd, S. J., Schroeder, I., Sarkar, N., Qiu, X., Sydeman, W. J., & Schwing, F. B. (2009). Phenology of coastal upwelling in the California Current. Geophysical Research Letters, 36(1).

12. British Geological Survery. (2011, November). Rare Earth Elements - Definitions, mineralogy, and deposits. Retrieved 2017, from Minerals UK: http://www.bgs.as.uk/downloads/start.cfm?id=1638.

32 13. Burdige, D. J. (2007). Preservation of organic matter in marine sediments: controls, mechanisms, and an imbalance in sediment organic carbon budgets?. Chemical reviews, 107(2), 467-485.

14. Busch, W. H., & Keller, G. H. (1981). The physical properties of Peru-Chile continental margin sediments--the influence of coastal upwelling on sediment properties. Journal of Sedimentary Research, 51(3).

15. Castor, S. B. (2008). Rare earth deposits of North America. Resource Geology, 58(4), 337-347.

16. Chernoff, C. B., & Orris, G. J. (2002). Data Set of World Phosphate Mines, Deposits, and Occurrences: Part A. Geologic Data. US Department of the Interior, US Geological Survey.

17. Clarke, R. J., & Altschuler, Z. (1958). Determination of the oxidation state of uranium in apatite and phosphorite deposits. Geochemica et Cosmochimica Acta , 13, 127-142.

18. De Voto, R.H., Stevens, D.N., 1979. Uraniferous Phosphate Resources and Technology and Economics of Uranium Recovery from Phosphate Resources, United States and Free World (No. GJBX-110(79) (Vol.2)). Earth Sciences, Inc., Golden, CO (USA).

19. Drummond, J. B., Pufahl, P. K., Porto, C. G., and Carvalho, M. (2015). Neoproterozoic peritidal phosphorite from the Sete Lagoas Formation (Brazil) and the Precambrian phosphorus cycle. Sedimentology, v. 62, p. 1978-2008.

20. Dubinin, A. (2004). Geochemistry of Rare Earth Elements in the Ocean. Lithology and Mineral , 39, 289.

21. Elderfield, H., & Greaves, M. J. (1981). Negative cerium anomalies in the rare earth element patterns of oceanic ferromanganese nodules. Earth and Planetary Science Letters, 55(1), 163-170.

22. Emsbo, P., McLaughlin, P. I., Breit, G. N., du Bray, E. A., & Koenig, A. E. (2015). Rare earth elements in sedimentary phosphate deposits: Solution to the global REE crisis? Gondwana Research , 27, 776-785.

23. Filippelli, G. M., Delaney, M. L., Garrison, R. E., Omarzai, S. K., & Behl, R. J. (1994). Phosphorus accumulation rates in a Miocene low oxygen basin: the Monterey Formation (Pismo Basin), California. Marine geology, 116(3-4), 419-430.

24. Filippeli, G. M., (2008). The global phosphorus cycle: Past, present and future. Elements, v. 4, p. 89-95.

33 25. Florida Department of Environmental Protection. (2014, November 4). Florida's Minerals. Retrieved April 2017, from Florida Department of Environmental Protection:http://www.dep.state.fl.us/geology/geologictopics/Minerals.htm #Mine%20 Locations%20in%20Florida

26. Froelich, P. N., Arthur, M. A., Burnett, W. C., Deakin, M., Hensley, V., Jahnke, R., Kaul, L., Kim, K. H., Roe, K., Soutar, A., and Vathakonan, C. (1988). Early diagenesis of organic matter in Peru continental margin sediments: Phosphorite precipitation. Marine Geology, v. 80, p. 309-343.

27. Goodenough, K., Deady, E., Wall, F., Shaw, F., & Lusty, P. (2014). The importance of tectonic setting in assessing European rare earth potential. European Rare Earth Resources .

28. Goodenough, K., Schilling, J., Jonsson, E., Kalvig, P., Charles, N., Tuduri, J., et al. (2016). Europe's rare earth element resource potential: An overview of REE mettalogenetic provinces and their geodynamic setting. Ore Geology Reviews , 72 (1), 838-856.

29. Google. (2018). Google Earth. Retrieved from https://earth.google.com/

30. Haupt B. and Seidov, D. (2012) Modeling geologically abrupt climate changes in the Miocene: Potential effects of high-latitudinal salinity changes. Natural Science, 4, p. 149-158. doi: 10.4236/ns.2012.43022

31. Heggie, D. T., Skyring, G. W., O’Brien, G. W., Reimers, C., Herczeg, A., Morarty, J. W., Burnett, W. C., and Milnes, A. R. (1990). Organic carbon cycling modern phosphorite formation on the East Australian continental margin: an overview, in Notholt, A.J.G., and Jarvis, I., eds., Phosphorite research and development. Geological Society of London Special Publication 52, p. 87-177.

32. Huddleston, P. F., (1988). A revision of the lithostratigraphic units of the coastal plain of Georgia, the Miocene through Holocene. Ga. Geological Survey Bulletin, 93, p. 56-73.

33. Humphries, M. (2010). Rare earth elements: the global supply chain. CRS Report for Congress.

34. Humphries, M. (2013). Rare earth elements: The global supply chain.

35. Hurst, M. V. (2012). Central Florida Phosphate District. Southeastern Geological Society.

36. Jansinski, S. (2011). Phosphate rock. U.S. Geological Survery Mineral Commodity Summaries.

34 37. Jasinski, S. M. (2016). Mineral commodity summaries 2016: U.S. Geological Survey, p. 124- 125, https://minerals.usgs.gov/minerals/pubs/commodity/phosphate_rock/mcs -2016-phosp.pdf

38. Jarvis, I., Burnett, W. C., Nathan, Y., Almbaydin, F.S.M., Attia, A.K.M., Castro, L. N., Flicoteaux, R., Hilmy, M. E., Husain, V., Qutawnah, A. A., Serjani, A., and Zanin, Y. N. (1994). Phosphorite geochemistry, state-of-the-art and environmental concerns. Eclogae Geologicae Helveticae, v. 87, p. 643-700.

39. Koepnick, R. B., Denison, R. E., & Dahl, D. A. (1988). The Cenozoic seawater 87Sr/86Sr curve: data review and implications for correlation of marine strata. Paleoceanography, 3(6), 743-756.

40. Lane, E. E. (1994). Florida's geological history and geological resources.

41. Long, K. R., Van Gosen, B. S., Foley, N. K., & Cordier, D. (2010-5220). The Principle Rare Earth Elements Deposits of the United States - A Summary of Domestic Deposits and a Global Perspective. Virginia: U.S. Geological Survey Scientific Investigations Report .

42. MacFadden, B. J., & Webb, S. D. (1982). The succession of Miocene (Arikareean through Hemphillian) terrestrial mammalian localities and faunas in Florida. Miocene of the Southeastern United States: Florida Bureau of Geology Special Publication, (25), 186-199.

43. Mariano, A. N. (2012). Rare Earth Mining and Exploration in North America. Elements , 8 (5), 369-376.

44. Matson, C. G., and Clapp, F. G. (1908). A preliminary report of the geology of Florida with special reference to the stratigraphy. Florida Geological Survey 2nd Annual Report, p. 25-173.

45. McGregor, H. V., Dima, M., Fischer, H. W., & Mulitza, S. (2007). Rapid 20th-century increase in coastal upwelling off northwest Africa. science, 315(5812), 637- 639.

46. Mercer, J. H., & Sutter, J. F. (1982). Late Miocene—earliest Pliocene glaciation in southern Argentina: implications for global ice-sheet history. Palaeogeography, Palaeoclimatology, Palaeoecology, 38(3-4), 185- 206.

47. Mills, M. M., Ridame, C., Davey, M., La Roche, J., and Geider, R. J. (2004). Iron and phosphorus co-limit nitrogen fixation in the eastern tropical North Atlantic. Nature, v. 429, p. 292-294.

35 48. Nielsen, S. G., Rehkämper, M., Baker, J., & Halliday, A. N. (2004). The precise and accurate determination of thallium isotope compositions and concentrations for water samples by MC-ICPMS. Chemical geology, 204(1-2), 109-124.

49. Ostrander, C. M., Owens, J. D., & Nielsen, S. G. (2017). Constraining the rate of oceanic deoxygenation leading up to a Cretaceous Oceanic Anoxic Event (OAE-2:~ 94 Ma). Science advances, 3(8), e1701020.

50. Owens, J. D., Nielsen, S. G., Horner, T. J., Ostrander, C. M., & Peterson, L. C. (2017). Thallium-isotopic compositions of euxinic sediments as a proxy for global manganese-oxide burial. Geochimica et Cosmochimica Acta, 213, 291-307.

51. Pattan, J. N., Pearce, N. J. G., & Mislankar, P. G. (2005). Constraints in using Cerium- anomaly of bulk sediments as an indicator of paleo bottom water redox environment: A case study from the Central Indian Ocean Basin. Chemical Geology, 221(3-4), 260-278.

52. Pufahl, P. K. (2010). Bioelemental sediments in James, N. P. and Dalrymple, R. W., eds. Facies models, 4th ed.: Geological Association of Canada, p. 477-503.

53. Pufahl, P. K., & Groat, L. A. (2017). Sedimentary and Igneous Phosphate Deposits: Formation and Exploration: An Invited Paper. Economic Geology, 112(3), 483- 516.

54. Rehkämper, M., & Halliday, A. N. (1999). The precise measurement of Tl isotopic compositions by MC-ICPMS: application to the analysis of geological materials and meteorites. Geochimica et Cosmochimica Acta, 63(6), 935-944.

55. Reinhard, C. T., Planavsky, N. J., Gill, B. C., Ozaki, K., Robbins, L. J., Lyons, T. W., ... & Konhauser, K. O. (2017). Evolution of the global phosphorus cycle. Nature, 541(7637), 386.

56. Scott, T. M. (1988). The lithostratigraphy of the Hawthorn Group (Miocene) of Florida (No. 59). Florida Geological Survey.

57. Sengupta, D., & Van Gosen, B. S. (2016). Placer-type rare earth element deposits. Rev. Econ. Geol, 18, 81-100.

58. Shetye, S. R., Shenoi, S. S. C., Gouveia, A. D., Michael, G. S., Sundar, D., & Nampoothiri, G. (1991). Wind-driven coastal upwelling along the western boundary of the Bay of Bengal during the southwest monsoon. Continental Shelf Research, 11(11), 1397-1408.

59. Takashima, R., Nishi, H., Huber, B. T., & Leckie, R. M. (2006). Greenhouse world and the Mesozoic ocean.

36 60. Taylor, S. R. (1964). Abundance of chemical elements in the continental crust: a new table. Geochimica et cosmochimica acta, 28(8), 1273-1285.

61. Taylor, S. R., & McLennan, S. M. (1985). The continental crust: Its evolution and composition. Lon on: Blackwell.

62. Van Cappellen, P., & Ingall, E. D. (1994). Benthic phosphorus regeneration, net primary production, and ocean anoxia: a model of the coupled marine biogeochemical cycles of carbon and phosphorus. Paleoceanography, 9(5), 677-692.

63. Van Gosen, B. S., Verplanck, P. L., Long, K. R., Gambogi, J., & Seal II, R. R. (2014, November). The Rare Earth Elements - Vital to Modern Technologies and Lifestyles. Retrieved April 2017, from USGS Mineral Resource Program: http://pubs.usgs.gov/fs/2014/3078/pdf/fs2014-3078.pdf

64. Van Kauwenbergh, S. J., & McClellan, G. (1990). Comparitive geology and mineralogy of the southern United States and Togo phosphorites. Geological Society of London - Special Publications , 52, 139-155.

65. Van Kauwenbergh, S. J. (2010). World phosphate rock reserves and resources (p. 48). Muscle Shoals: IFDC.

66. Webb, S. D., & Crissinger, D. B. (1983). Stratigraphy and vertebrate paleontology of the central and southern phosphate districts of Florida. Central Florida Phosphate District Field Trip Guidebook (Geological Society of America, Southeastern Section), 28-72.

67. Wilbur, S., Bellevue, W. A., & McCurdy, E. (2001). Determination of trace levels of phosphorus in environmental samples with 7500c ICPMS system. Agilent Technologies, Inc, 1-4.

68. Zanin, Y., & Zamirailova, A. (2009). Rare earth elements in supergene phosphorites. Geochemistry International , 47, 282-296.

69. Zhang, Y., Jiang, Z., He, M., & Hu, B. (2007). Determination of trace rare earth elements in coal fly ash and atmospheric particulates by electrothermal vaporization inductively coupled plasma mass spectrometry with slurry sampling. Environmental Pollution , 148 (2), 459-467.

70. Zhi Li, L., & Yang, X. (2014). China's Rare Earth Ore Deposits and Benefication Techniques. China Western Mining Co.; Geological Survey of Finland. Milos: 1st European Rare Earth Resources Conference.

37 71. Zhou, B., Li, Z., & Chen, C. (2017). Global Potential of Rare Earth Resources and Rare Earth Demand from Clean Technologies. Minerals, 7(11), 203.

38 APPENDIX

Table 3: Geochemical data

Core: Type: Depth (m) P2O5 Al Ca Mn Fe Y wt.% wt.% wt.% (ppm) wt.% (ppm) MC02017 Tooth 8.5 38.10 2.22 16.31 49.3 0.83 69.1 MC02017 Tooth 13.7 35.96 0.15 20.61 67.8 0.27 12.1 MC02017 Bone 15.8 35.23 0.19 13.95 131.1 0.38 10.6

MC02017 PO4 Clay 13.4 32.16 0.46 4.36 93.9 0.19 76.5

MC02017 PO4 Clay 13.3 30.11 0.43 1.88 93.7 0.18 67.5

MC02017 PO4 Clay 13.1 33.00 0.59 10.17 77.9 0.45 64.5 MC02017 Blk. Grain 9.8 36.94 2.69 19.82 62.7 0.92 191.2 MC02017 Blk. Grain 8.7 32.71 1.33 4.87 111.0 0.12 125.3 MC02017 Blk. Grain 12.2 38.42 0.77 7.71 71.5 0.29 105.1 MC02017 Blk. Grain 17.7 32.56 0.80 10.06 103.3 0.91 6.7 MC02017 Blk. Grain 8.8 37.57 2.11 26.07 64.6 1.04 133.7 MC02017 Blk. Grain 14.6 34.34 1.05 8.63 121.8 2.54 118.4 MC02017 Blk. Grain 14.6 24.42 0.50 7.79 74.4 1.65 75.8 MC02017 Bulk 16.2 35.79 1.38 37.98 405.2 3.36 122.5

MC02017 PO4 Clay 13.4 7.15 0.75 BDL 34.1 0.16 15.5 MC02017 Bulk 17.7 13.70 1.22 5.42 122.9 0.98 59.1 HC92017 Wht. Grain 5.5 36.38 0.19 2.91 69.7 0.03 3.6 HC92017 Wht. Grain 6.4 35.32 0.26 2.08 76.5 0.03 22.2 HC92017 Wht. Grain 6.4 31.20 0.21 0.71 68.7 0.00 BDL HC92017 Wht. Grain 7.3 32.15 0.32 2.31 71.8 0.01 14.7 HC92017 Wht. Grain 7.9 25.65 0.46 0.74 72.1 0.36 83.2 HC92017 Wht. Grain 7.9 35.30 0.15 2.97 87.3 0.06 4.5 HC92017 Wht. Grain 9.1 30.09 0.17 1.38 86.2 0.02 29.6

39

Table 3: Continued

Core: Type: Depth (m) P2O5 Al Ca Mn Fe Y wt.% wt.% wt.% (ppm) wt.% (ppm) HC92017 Blk. Grain 9.1 30.05 0.08 1.28 77.3 0.04 BDL HC92017 Wht. Grain 9.4 13.29 1.72 0.13 55.8 0.75 40.0 HC92017 Wht. Grain 9.8 31.40 0.33 1.75 67.4 0.03 18.5 HC92017 Wht. Grain 9.8 26.04 0.47 2.77 67.5 0.21 2.1 HC92017 Brn. Grain 10.4 22.69 0.41 2.84 179.5 0.82 BDL HC92017 Brn. Grain 11.3 26.63 0.39 2.58 97.2 0.11 20.8 Brn. Grain 11.9 20.83 1.00 4.49 38.1 0.34 52.0 HC92017 Brn. Grain 11.9 30.09 0.35 2.83 52.8 BDL 15.6 HC92017 Brn. Grain 12.8 27.76 0.28 2.96 117.9 0.01 20.8 HC92017 Brn. Grain 13.4 26.90 0.34 2.31 85.5 BDL 39.9 HC92017 Brn. Grain 13.4 29.18 0.38 4.63 69.5 0.00 31.0 HC92017 Wht. Grain 14.3 24.22 0.29 4.61 65.5 0.24 10.6 HC92017 Wht. Grain 14.3 19.35 0.59 3.57 50.4 0.61 28.7 HC92017 Brn. Grain 14.6 26.93 0.15 7.96 88.6 0.20 0.1 HC92017 Wht. Grain 14.6 2.43 0.14 3.79 182.4 0.08 BDL HC92017 Wht. Grain 14.6 5.04 0.37 6.72 158.7 0.06 12.7 HC92017 Brn. Grain 14.6 27.17 0.22 6.61 53.4 0.15 22.1 HC92017 Wht. Grain 14.6 6.37 0.24 5.07 165.6 0.07 5.3 HC92017 Wht. Grain 14.9 18.66 0.19 5.80 162.5 0.22 11.7 HC92017 Brn. Grain 15.5 24.10 0.31 6.03 111.5 0.05 15.4 HC92017 Wht. Grain 15.5 5.31 0.25 9.29 413.3 0.24 8.9 HC92017 Bulk 15.8 27.77 0.24 6.88 89.6 BDL 35.4 HC92017 Bulk 16.5 6.12 0.31 8.92 113.7 0.08 48.7

40

Table 3: Continued

Core: Type: Depth (m) P2O5 Al Ca Mn Fe Y wt.% wt.% wt.% (ppm) wt.% (ppm) HC92017 Bulk 17.7 13.17 0.55 13.03 127.0 0.23 90.6 HC92017 Bulk 17.7 12.31 0.48 11.88 181.2 0.07 57.0 HC92017 Bulk 19.5 4.53 0.18 6.75 161.2 0.02 4.0 HC92017 Bulk 19.5 11.21 0.25 11.69 112.6 BDL 14.1 HC92017 Bulk/red 10.4 2.98 1.03 1.39 296.4 1.79 32.7 HC92017 Bulk/red 11.3 5.06 1.08 2.84 49.5 0.46 50.4 HC02017 Grain 4.3 33.57 0.46 20.30 92.9 0.40 350.5 HC02017 Grain 4.9 31.38 0.47 26.33 140.2 1.41 64.1 HC02017 Grain 4.9 27.03 0.33 16.78 105.0 0.93 46.8 HC02017 Grain 5.8 25.61 0.26 17.33 153.4 0.71 46.7 HC02017 Grain 6.4 26.57 0.21 21.39 128.8 0.88 88.0 HC02017 Grain 6.4 27.32 0.20 20.65 122.6 0.92 82.2 HC02017 Wht. Grain 8.2 31.21 0.84 26.61 141.1 0.60 181.9 HC02017 Grain 8.8 28.99 0.34 24.27 173.2 0.82 89.6 HC02017 Wht. Grain 9.8 27.17 0.17 23.82 59.2 0.23 88.2 HC02017 Wht. Grain 10.7 18.99 0.17 19.55 42.8 1.48 56.6 HC02017 Wht. Grain 11.6 21.00 0.36 21.72 104.2 0.72 64.4 HC02017 Wht. Grain 12.5 18.81 0.32 25.75 99.7 0.66 69.8 HC02017 Wht. Grain 13.1 24.47 0.30 23.99 110.6 0.47 67.5 HC02017 Wht. Grain 14.3 29.25 0.38 23.66 83.4 0.37 78.9 HC02017 Blk. Grain 14.9 26.56 0.19 24.55 27.6 0.28 63.9 HC02017 Blk. Grain 15.5 27.39 0.31 23.08 105.3 0.38 152.0 HC02017 Blk. Grain 16.8 25.23 0.25 22.34 112.6 0.67 53.6

41

Table 3: Continued

Core: Type: Depth (m) P2O5 Al Ca Mn Fe Y wt.% wt.% wt.% (ppm) wt.% (ppm) HC02017 Blk. Grain 18.0 26.25 0.15 22.21 29.8 0.20 75.6 HC02017 Blk. Grain 18.9 27.27 0.26 27.42 31.7 0.37 88.6 HC02017 Blk. Grain 19.8 24.67 0.26 26.97 23.7 0.22 87.7 HC02017 Blk. Grain 20.7 19.55 0.36 19.67 98.1 0.21 273.0 HC02017 Blk. Grain 21.9 26.05 0.26 51.63 90.5 0.37 68.9 HC02017 Blk. Grain 23.5 26.26 0.42 26.04 210.7 0.92 83.1 HC02017 Blk. Grain 24.1 23.72 0.15 29.18 86.0 0.21 65.8 HC02017 Blk. Grain 25.0 26.07 0.14 28.10 52.6 0.39 92.3 HC02017 Blk. Grain 25.6 22.50 0.12 23.18 17.5 0.32 44.5 HC02017 Blk. Grain 25.9 25.64 0.14 29.39 63.4 0.33 66.2 HC02017 Blk. Grain 26.2 28.47 0.08 34.45 BDL 0.14 124.3 HC02017 Tooth 26.2 30.78 0.05 34.10 83.0 0.11 43.3 HC02017 Blk. Grain 26.5 23.44 0.23 54.23 26.0 0.44 70.1

Average: 24.79 0.48 14.18 102.8 0.48 63.2 Min: 2.43 0.05 0.13 17.5 0.00 0.1 Max: 38.42 2.69 54.23 413.3 3.36 350.5 Standards:

SCO-2 Measured 0.19 6.50 2.10 343.3 3.42 23.4 SCO-2 Known 0.18 6.94 2.81 333.1 3.25 23.1 % Error 5.26 6.80 25.27 2.97 4.97 1.29

42

Table 3: Continued Core: Type: Depth (m) La Ce Pr Nd Sm Eu (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) MC02017 Tooth 8.5 6.5 14.6 2.0 10.9 2.9 0.9 MC02017 Tooth 13.7 3.6 4.1 0.5 2.5 0.4 0.1 MC02017 Bone 15.8 2.4 2.0 0.2 1.1 0.2 0.1

MC02017 PO4 Clay 13.4 42.9 67.6 8.0 43.8 6.1 1.4

MC02017 PO4 Clay 13.3 40.2 64.5 7.3 40.1 5.4 1.2

MC02017 PO4 Clay 13.1 42.5 67.7 7.5 41.7 5.4 1.2 MC02017 Blk. Grain 9.8 55.7 88.6 11.6 64.0 9.8 2.5 MC02017 Blk. Grain 8.7 46.6 79.5 9.5 48.2 8.7 2.2 MC02017 Blk. Grain 12.2 57.6 95.8 10.4 60.0 8.3 2.1 MC02017 Blk. Grain 17.7 0.0 0.0 0.0 0.0 0.0 0.0 MC02017 Blk. Grain 8.8 52.2 87.6 11.6 60.0 9.7 2.4 MC02017 Blk. Grain 14.6 67.3 103.5 11.1 63.8 8.5 2.0 MC02017 Blk. Grain 14.6 59.5 96.2 12.1 61.7 8.8 2.1 MC02017 Bulk 16.2 112.1 178.1 21.9 113.8 16.3 3.8

MC02017 PO4 Clay 13.4 20.9 40.0 4.6 23.5 4.0 1.0 MC02017 Bulk 17.7 53.6 91.1 11.8 61.0 9.3 2.3 HC92017 Wht. Grain 5.5 5.3 6.0 1.0 4.5 0.7 0.2 HC92017 Wht. Grain 6.4 23.0 31.2 4.2 20.3 3.3 0.8 HC92017 Wht. Grain 6.4 3.1 3.2 0.6 2.6 0.4 0.1 HC92017 Wht. Grain 7.3 18.7 25.5 3.2 15.5 2.4 0.6 HC92017 Wht. Grain 7.9 75.5 116.0 15.1 75.5 10.8 2.7 HC92017 Wht. Grain 7.9 7.1 10.5 1.4 6.8 1.0 0.3 HC92017 Wht. Grain 9.1 30.8 47.0 5.2 29.7 4.1 1.0

43

Table 3: Continued Core: Type: Depth (m) La Ce Pr Nd Sm Eu (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) HC92017 Blk. Grain 9.1 2.3 3.5 0.5 2.3 0.4 0.1 HC92017 Wht. Grain 9.4 45.4 79.1 9.2 50.8 7.3 1.8 HC92017 Wht. Grain 9.8 19.8 30.6 3.5 17.6 2.7 0.7 HC92017 Wht. Grain 9.8 6.9 10.1 1.4 6.9 1.1 0.3 HC92017 Brn. Grain 10.4 7.8 4.2 1.6 6.2 1.1 0.3 HC92017 Brn. Grain 11.3 33.2 39.1 5.9 31.0 4.7 1.2 HC92017 Brn. Grain 11.9 43.9 67.5 8.3 46.4 7.2 1.9 HC92017 Brn. Grain 11.9 22.5 32.2 4.6 24.9 4.1 1.1 HC92017 Brn. Grain 12.8 41.0 44.2 7.7 37.9 6.1 1.5 HC92017 Brn. Grain 13.4 46.9 63.9 9.5 46.2 6.9 1.7 HC92017 Brn. Grain 13.4 46.3 49.8 10.4 46.7 8.3 2.1 HC92017 Wht. Grain 14.3 12.4 17.3 2.2 11.0 1.8 0.4 HC92017 Wht. Grain 14.3 31.8 37.2 5.8 28.9 4.6 1.1 HC92017 Brn. Grain 14.6 7.7 6.2 1.3 5.6 0.9 0.2 HC92017 Wht. Grain 14.6 1.2 1.1 0.3 1.2 0.2 0.0 HC92017 Wht. Grain 14.6 7.1 15.2 1.7 8.9 1.6 0.4 HC92017 Brn. Grain 14.6 30.0 46.1 5.7 31.3 4.7 0.9 HC92017 Wht. Grain 14.6 6.7 10.6 1.5 7.5 1.6 0.4 HC92017 Wht. Grain 14.9 12.0 14.4 1.9 9.3 1.5 0.4 HC92017 Brn. Grain 15.5 22.4 33.9 4.2 22.8 3.2 0.7 HC92017 Wht. Grain 15.5 7.2 13.8 1.5 7.7 1.2 0.3 HC92017 Bulk 15.8 43.9 29.3 7.4 33.1 6.1 1.5 HC92017 Bulk 16.5 33.3 45.2 6.2 33.7 5.5 1.3

44

Table 3: Continued Core: Type: Depth (m) La Ce Pr Nd Sm Eu (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) HC92017 Bulk 17.7 57.3 79.0 11.9 57.9 9.2 2.2 HC92017 Bulk 17.7 29.7 38.8 5.7 30.4 5.4 1.4 HC92017 Bulk 19.5 8.5 13.8 2.0 9.5 1.7 0.3 HC92017 Bulk 19.5 7.9 7.0 1.7 7.4 1.6 0.4 HC92017 Bulk/red 10.4 25.9 46.4 5.4 30.3 4.5 1.2 HC92017 Bulk/red 11.3 36.6 63.8 8.3 42.3 6.3 1.6 HC02017 Grain 4.3 57.1 104.3 14.9 76.7 15.3 4.0 HC02017 Grain 4.9 33.0 46.0 6.0 31.0 4.6 1.1 HC02017 Grain 4.9 33.2 53.9 6.2 34.8 4.9 1.0 HC02017 Grain 5.8 38.9 51.3 7.5 37.2 5.6 1.3 HC02017 Grain 6.4 57.4 73.9 11.0 52.2 8.1 1.9 HC02017 Grain 6.4 50.2 61.1 9.3 44.9 7.0 1.7 HC02017 Wht. Grain 8.2 46.3 71.2 9.4 47.3 7.1 1.8 HC02017 Grain 8.8 50.8 63.2 8.9 44.2 6.5 1.6 HC02017 Wht. Grain 9.8 48.8 51.4 8.2 38.6 5.9 1.5 HC02017 Wht. Grain 10.7 32.3 34.3 5.0 26.2 4.1 1.0 HC02017 Wht. Grain 11.6 35.8 49.5 7.0 34.2 5.1 1.2 HC02017 Wht. Grain 12.5 37.4 45.1 6.0 32.2 4.9 1.2 HC02017 Wht. Grain 13.1 41.8 50.0 6.5 35.3 5.1 1.2 HC02017 Wht. Grain 14.3 35.7 46.5 5.9 32.2 4.8 1.2 HC02017 Blk. Grain 14.9 30.7 30.4 4.7 24.2 3.9 1.0 HC02017 Blk. Grain 15.5 64.3 76.4 11.4 55.3 8.4 2.1 HC02017 Blk. Grain 16.8 30.5 38.3 4.8 26.3 3.7 0.9

45

Table 3: Continued Core: Type: Depth (m) La Ce Pr Nd Sm Eu (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) HC02017 Blk. Grain 18.0 36.9 38.5 5.6 29.3 4.6 1.2 HC02017 Blk. Grain 18.9 46.1 51.5 7.2 37.9 5.8 1.4 HC02017 Blk. Grain 19.8 44.3 49.9 7.5 38.3 5.8 1.5 HC02017 Blk. Grain 20.7 132.4 144.4 26.0 117.7 19.7 4.8 HC02017 Blk. Grain 21.9 34.7 36.0 5.2 27.2 4.1 1.0 HC02017 Blk. Grain 23.5 42.9 55.2 7.2 39.1 5.9 1.4 HC02017 Blk. Grain 24.1 36.2 33.9 5.0 25.7 3.8 1.0 HC02017 Blk. Grain 25.0 58.0 60.1 9.3 43.2 5.9 1.5 HC02017 Blk. Grain 25.6 21.5 19.4 2.8 12.7 2.1 0.6 HC02017 Blk. Grain 25.9 32.0 30.0 4.5 23.2 3.5 0.9 HC02017 Blk. Grain 26.2 80.8 92.1 14.9 67.5 9.3 2.7 HC02017 Tooth 26.2 26.2 16.7 2.1 9.6 1.2 0.4 HC02017 Blk. Grain 26.5 32.1 33.1 4.9 25.7 3.8 1.0

Average: 35.0 47.2 6.5 33.1 5.1 1.3 Min: 0.0 0.0 0.0 0.0 0.0 0.0 Max: 132.4 178.1 26.0 117.7 19.7 4.8 Standards:

SCO-2 Measured 33.0 62.6 7.7 29.8 4.9 1.2 SCO-2 Known 26.4 54.5 6.6 25.0 4.9 1.1 % Error 19.91 12.99 14.29 16.11 0.00 9.58

46

Table 3: Continued Core: Type: Depth (m) Gd Tb Dy Ho Er Tm (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) MC02017 Tooth 8.5 3.9 0.7 4.7 1.2 4.2 0.6 MC02017 Tooth 13.7 0.5 0.1 0.5 0.1 0.5 0.1 MC02017 Bone 15.8 0.2 0.0 0.3 0.1 0.3 0.1

MC02017 PO4 Clay 13.4 7.2 1.1 6.3 1.4 4.3 0.6

MC02017 PO4 Clay 13.3 6.4 0.9 5.5 1.2 3.7 0.5

MC02017 PO4 Clay 13.1 6.3 0.9 5.3 1.2 3.7 0.5 MC02017 Blk. Grain 9.8 13.6 2.0 12.3 3.2 10.6 1.4 MC02017 Blk. Grain 8.7 10.5 1.6 9.7 2.2 7.0 1.0 MC02017 Blk. Grain 12.2 9.9 1.5 8.9 2.0 6.1 0.8 MC02017 Blk. Grain 17.7 0.0 0.0 0.0 0.0 0.0 0.0 MC02017 Blk. Grain 8.8 11.8 1.9 11.3 2.5 7.9 1.1 MC02017 Blk. Grain 14.6 10.4 1.5 9.3 2.1 6.6 0.9 MC02017 Blk. Grain 14.6 10.8 1.7 10.2 2.3 7.0 1.0 MC02017 Bulk 16.2 20.4 3.1 19.0 4.3 13.4 1.8

MC02017 PO4 Clay 13.4 4.4 0.7 3.9 0.8 2.4 0.3 MC02017 Bulk 17.7 11.2 1.7 10.2 2.3 6.8 0.9 HC92017 Wht. Grain 5.5 0.8 0.1 0.7 0.2 0.5 0.1 HC92017 Wht. Grain 6.4 4.1 0.6 3.7 0.8 2.5 0.3 HC92017 Wht. Grain 6.4 0.5 0.1 0.5 0.1 0.3 0.0 HC92017 Wht. Grain 7.3 3.0 0.5 2.9 0.7 2.1 0.3 HC92017 Wht. Grain 7.9 13.2 2.0 12.3 2.8 8.8 1.2 HC92017 Wht. Grain 7.9 1.3 0.2 1.2 0.3 0.9 0.1 HC92017 Wht. Grain 9.1 5.2 0.8 5.0 1.1 3.6 0.5

47

Table 3: Continued Core: Type: Depth (m) Gd Tb Dy Ho Er Tm (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) HC92017 Blk. Grain 9.1 0.5 0.1 0.4 0.1 0.3 0.0 HC92017 Wht. Grain 9.4 8.5 1.3 7.6 1.7 5.2 0.7 HC92017 Wht. Grain 9.8 3.5 0.5 3.3 0.8 2.4 0.3 HC92017 Wht. Grain 9.8 1.3 0.2 1.3 0.3 0.9 0.1 HC92017 Brn. Grain 10.4 1.3 0.2 1.2 0.3 0.9 0.1 HC92017 Brn. Grain 11.3 5.9 0.9 5.8 1.4 4.2 0.6 HC92017 Brn. Grain 11.9 9.2 1.4 9.0 2.1 6.5 0.9 HC92017 Brn. Grain 11.9 5.3 0.8 5.2 1.2 3.8 0.5 HC92017 Brn. Grain 12.8 7.6 1.2 7.2 1.7 5.2 0.7 HC92017 Brn. Grain 13.4 8.6 1.3 8.2 1.9 5.9 0.8 HC92017 Brn. Grain 13.4 10.1 1.5 9.4 2.2 6.6 0.8 HC92017 Wht. Grain 14.3 2.2 0.3 2.1 0.5 1.5 0.2 HC92017 Wht. Grain 14.3 5.5 0.8 5.1 1.2 3.7 0.5 HC92017 Brn. Grain 14.6 1.1 0.1 0.9 0.2 0.6 0.1 HC92017 Wht. Grain 14.6 0.2 0.0 0.2 0.0 0.1 0.0 HC92017 Wht. Grain 14.6 2.1 0.3 2.0 0.5 1.4 0.2 HC92017 Brn. Grain 14.6 5.3 0.8 4.2 0.9 2.8 0.4 HC92017 Wht. Grain 14.6 2.0 0.3 2.0 0.5 1.4 0.2 HC92017 Wht. Grain 14.9 2.0 0.3 1.9 0.4 1.4 0.2 HC92017 Brn. Grain 15.5 3.6 0.5 2.9 0.7 2.1 0.3 HC92017 Wht. Grain 15.5 1.5 0.2 1.3 0.3 0.9 0.1 HC92017 Bulk 15.8 7.7 1.2 7.5 1.8 5.5 0.7 HC92017 Bulk 16.5 6.7 1.0 6.2 1.4 4.2 0.6

48

Table 3: Continued Core: Type: Depth (m) Gd Tb Dy Ho Er Tm (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) HC92017 Bulk 17.7 11.4 1.7 10.7 2.4 7.4 1.0 HC92017 Bulk 17.7 6.9 1.1 7.2 1.7 5.3 0.7 HC92017 Bulk 19.5 2.0 0.3 1.5 0.3 1.0 0.1 HC92017 Bulk 19.5 2.4 0.4 2.6 0.7 2.1 0.3 HC92017 Bulk/red 10.4 5.2 0.8 4.4 0.9 2.8 0.4 HC92017 Bulk/red 11.3 7.3 1.1 6.4 1.4 4.2 0.6 HC02017 Grain 4.3 21.4 3.6 25.4 7.5 33.8 5.0 HC02017 Grain 4.9 5.8 0.9 5.7 1.5 5.3 0.8 HC02017 Grain 4.9 5.8 0.8 5.0 1.3 4.4 0.6 HC02017 Grain 5.8 7.2 1.1 6.6 1.6 5.0 0.7 HC02017 Grain 6.4 10.2 1.6 9.6 2.2 7.0 0.9 HC02017 Grain 6.4 8.9 1.4 8.7 2.1 6.5 0.9 HC02017 Wht. Grain 8.2 10.5 1.6 11.5 3.7 15.3 2.5 HC02017 Grain 8.8 8.5 1.3 8.5 2.1 7.1 1.0 HC02017 Wht. Grain 9.8 7.9 1.2 8.3 2.0 6.6 0.9 HC02017 Wht. Grain 10.7 5.4 0.8 5.5 1.3 4.3 0.6 HC02017 Wht. Grain 11.6 6.4 1.0 6.1 1.5 5.3 0.8 HC02017 Wht. Grain 12.5 6.4 1.0 6.4 1.6 5.4 0.8 HC02017 Wht. Grain 13.1 6.7 1.0 6.6 1.6 5.2 0.7 HC02017 Wht. Grain 14.3 6.4 1.0 6.5 1.7 6.3 0.9 HC02017 Blk. Grain 14.9 5.4 0.8 5.6 1.4 4.4 0.6 HC02017 Blk. Grain 15.5 11.5 1.8 11.6 3.1 11.1 1.6 HC02017 Blk. Grain 16.8 4.8 0.7 4.6 1.1 3.6 0.5

49

Table 3: Continued Core: Type: Depth (m) Gd Tb Dy Ho Er Tm (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) HC02017 Blk. Grain 18.0 6.1 1.0 6.3 1.5 4.9 0.7 HC02017 Blk. Grain 18.9 7.7 1.2 7.7 1.8 5.8 0.8 HC02017 Blk. Grain 19.8 7.5 1.2 7.4 1.8 5.4 0.7 HC02017 Blk. Grain 20.7 24.6 3.9 24.5 5.7 17.7 2.4 HC02017 Blk. Grain 21.9 5.4 0.8 5.4 1.3 4.1 0.6 HC02017 Blk. Grain 23.5 7.4 1.1 7.2 1.8 5.9 0.8 HC02017 Blk. Grain 24.1 5.0 0.8 5.0 1.2 3.8 0.5 HC02017 Blk. Grain 25.0 7.6 1.1 7.2 1.7 5.2 0.7 HC02017 Blk. Grain 25.6 2.8 0.4 3.0 0.7 2.3 0.3 HC02017 Blk. Grain 25.9 4.7 0.7 4.6 1.1 3.5 0.5 HC02017 Blk. Grain 26.2 11.7 1.7 10.5 2.4 6.9 0.9 HC02017 Tooth 26.2 2.0 0.3 2.3 0.6 2.1 0.3 HC02017 Blk. Grain 26.5 5.0 0.8 4.9 1.2 3.5 0.4

Average: 6.5 1.0 6.2 1.5 4.9 0.7 Min: 0.0 0.0 0.0 0.0 0.0 0.0 Max: 24.6 3.9 25.4 7.5 33.8 5.0 Standards:

SCO-2 Measured 4.9 0.7 3.9 0.7 2.1 0.3 SCO-2 Known 4.4 0.7 3.9 0.8 2.3 0.3 % Error 11.04 3.57 0.00 11.40 7.90 6.25

50

Table 3: Continued Core: Type: Depth (m) Yb Lu U T REE T REE (ppm) (ppm) (ppm) (ppm) wt.% MC02017 Tooth 8.5 4.6 0.8 114.4 127.6 0.01 MC02017 Tooth 13.7 0.8 0.2 82.6 26.1 0.00 MC02017 Bone 15.8 0.5 0.1 49.6 18.2 0.00

MC02017 PO4 Clay 13.4 3.7 0.6 133.5 271.3 0.03

MC02017 PO4 Clay 13.3 3.2 0.5 206.5 248.1 0.02

MC02017 PO4 Clay 13.1 3.3 0.5 120.1 252.3 0.03 MC02017 Blk. Grain 9.8 9.2 1.7 563.1 477.4 0.05 MC02017 Blk. Grain 8.7 6.3 1.0 337.0 359.2 0.04 MC02017 Blk. Grain 12.2 5.5 0.8 169.1 374.8 0.04 MC02017 Blk. Grain 17.7 0.0 0.0 0.0 6.8 0.00 MC02017 Blk. Grain 8.8 7.4 1.1 319.6 402.3 0.04 MC02017 Blk. Grain 14.6 5.8 0.9 99.9 412.1 0.04 MC02017 Blk. Grain 14.6 6.3 1.0 226.6 356.6 0.04 MC02017 Bulk 16.2 12.0 1.9 248.5 644.4 0.06

MC02017 PO4 Clay 13.4 2.2 0.3 33.0 124.5 0.01 MC02017 Bulk 17.7 5.9 0.9 56.2 328.0 0.03 HC92017 Wht. Grain 5.5 0.5 0.1 178.9 24.4 0.00 HC92017 Wht. Grain 6.4 2.2 0.3 205.9 119.7 0.01 HC92017 Wht. Grain 6.4 0.3 0.0 240.9 11.9 0.00 HC92017 Wht. Grain 7.3 1.9 0.3 164.3 92.1 0.01 HC92017 Wht. Grain 7.9 8.2 1.3 136.9 428.5 0.04 HC92017 Wht. Grain 7.9 0.8 0.1 180.1 36.6 0.00 HC92017 Wht. Grain 9.1 3.4 0.5 149.4 167.7 0.02

51

Table 3: Continued Core: Type: Depth (m) Yb Lu U T REE T REE (ppm) (ppm) (ppm) (ppm) wt.% HC92017 Blk. Grain 9.1 0.3 0.0 167.9 10.9 0.00 HC92017 Wht. Grain 9.4 4.6 0.7 66.0 263.9 0.03 HC92017 Wht. Grain 9.8 2.2 0.3 122.8 106.8 0.01 HC92017 Wht. Grain 9.8 0.8 0.1 153.0 33.8 0.00 HC92017 Brn. Grain 10.4 0.6 0.1 238.6 25.8 0.00 HC92017 Brn. Grain 11.3 3.5 0.5 144.3 158.5 0.02 HC92017 Brn. Grain 11.9 5.9 0.9 148.9 263.2 0.03 HC92017 Brn. Grain 11.9 3.0 0.5 342.1 125.4 0.01 HC92017 Brn. Grain 12.8 4.1 0.6 146.8 187.5 0.02 HC92017 Brn. Grain 13.4 4.9 0.8 151.2 247.4 0.02 HC92017 Brn. Grain 13.4 5.2 0.8 163.0 231.3 0.02 HC92017 Wht. Grain 14.3 1.3 0.2 115.2 64.2 0.01 HC92017 Wht. Grain 14.3 3.3 0.5 94.9 158.6 0.02 HC92017 Brn. Grain 14.6 0.4 0.1 161.1 25.4 0.00 HC92017 Wht. Grain 14.6 0.1 0.0 12.0 4.5 0.00 HC92017 Wht. Grain 14.6 1.2 0.2 35.3 55.6 0.01 HC92017 Brn. Grain 14.6 2.4 0.4 196.5 157.9 0.02 HC92017 Wht. Grain 14.6 1.1 0.2 45.1 41.2 0.00 HC92017 Wht. Grain 14.9 1.3 0.2 343.1 58.8 0.01 HC92017 Brn. Grain 15.5 1.8 0.3 264.9 114.5 0.01 HC92017 Wht. Grain 15.5 0.9 0.1 25.6 45.9 0.00 HC92017 Bulk 15.8 4.4 0.7 194.7 186.2 0.02 HC92017 Bulk 16.5 3.6 0.6 28.6 198.2 0.02

52

Table 3: Continued Core: Type: Depth (m) Yb Lu U T REE T REE (ppm) (ppm) (ppm) (ppm) wt.% HC92017 Bulk 17.7 6.2 0.9 74.3 349.8 0.03 HC92017 Bulk 17.7 4.6 0.7 46.5 196.6 0.02 HC92017 Bulk 19.5 0.6 0.1 34.0 45.7 0.00 HC92017 Bulk 19.5 1.7 0.3 114.6 50.4 0.01 HC92017 Bulk/red 10.4 2.4 0.4 13.9 163.5 0.02 HC92017 Bulk/red 11.3 3.6 0.5 20.9 234.3 0.02 HC02017 Grain 4.3 41.8 6.6 217.3 767.8 0.08 HC02017 Grain 4.9 5.7 1.0 149.6 212.3 0.02 HC02017 Grain 4.9 4.5 0.8 244.5 204.0 0.02 HC02017 Grain 5.8 4.4 0.7 138.2 215.8 0.02 HC02017 Grain 6.4 6.2 1.0 125.5 331.3 0.03 HC02017 Grain 6.4 5.7 0.9 121.2 291.4 0.03 HC02017 Wht. Grain 8.2 18.5 3.4 176.7 432.0 0.04 HC02017 Grain 8.8 6.6 1.1 204.0 301.2 0.03 HC02017 Wht. Grain 9.8 6.0 1.0 164.8 276.5 0.03 HC02017 Wht. Grain 10.7 3.8 0.6 109.6 182.0 0.02 HC02017 Wht. Grain 11.6 5.5 0.9 150.9 224.8 0.02 HC02017 Wht. Grain 12.5 5.2 0.9 222.8 224.2 0.02 HC02017 Wht. Grain 13.1 4.7 0.7 161.0 234.6 0.02 HC02017 Wht. Grain 14.3 6.5 1.2 155.3 235.6 0.02 HC02017 Blk. Grain 14.9 3.8 0.6 346.6 181.6 0.02 HC02017 Blk. Grain 15.5 11.3 2.0 402.9 423.9 0.04 HC02017 Blk. Grain 16.8 3.3 0.5 121.0 177.3 0.02

53

Table 3: Continued Core: Type: Depth (m) Yb Lu U T REE T REE (ppm) (ppm) (ppm) (ppm) wt.% HC02017 Blk. Grain 18.0 4.2 0.7 262.4 217.0 0.02 HC02017 Blk. Grain 18.9 4.8 0.7 341.6 269.0 0.03 HC02017 Blk. Grain 19.8 4.4 0.7 267.7 264.1 0.03 HC02017 Blk. Grain 20.7 15.0 2.3 66.0 814.0 0.08 HC02017 Blk. Grain 21.9 3.6 0.6 106.8 198.8 0.02 HC02017 Blk. Grain 23.5 5.7 1.0 118.3 265.8 0.03 HC02017 Blk. Grain 24.1 3.4 0.5 203.5 191.4 0.02 HC02017 Blk. Grain 25.0 4.3 0.7 143.5 298.7 0.03 HC02017 Blk. Grain 25.6 2.0 0.3 127.9 115.5 0.01 HC02017 Blk. Grain 25.9 2.8 0.4 211.0 178.6 0.02 HC02017 Blk. Grain 26.2 5.3 0.7 145.8 431.7 0.04 HC02017 Tooth 26.2 2.1 0.3 44.0 109.4 0.01 HC02017 Blk. Grain 26.5 2.7 0.4 203.6 189.8 0.02

Average: 4.6 0.7 160.2 214.3 0.02 Min: 0.0 0.0 0.0 4.5 0.00 Max: 41.8 6.6 563.1 814.0 0.08 Standards:

SCO-2 Measured 2.1 0.3 3.8 SCO-2 Known 2.2 0.4 3.2 % Error 4.76 21.20 15.79

54 BIOGRAPHICAL SKETCH Education 2015 B.S. (Geology) Florida State University GPA 3.52 2012 A.A. (General Studies) Hillsborough Community College GPA 3.90

Laboratory Experience Florida State University Sample preparation and lab maintenance - mixed acid baths - preparing samples for ICP-MS analysis

National High Magnetic Field Laboratory Sample preparation in a trace metal clean lab - triple acid digestion - anion/cation column chemistry - isotope dilution chemistry - preparation of standard solutions

Instrument Operation Experience Agilent 7500cs Inductively Coupled Plasma Mass-Spectrometer Thermo Finnigan MAT 262 TIMS Thermo Scientific Neptune Multi-Collector ICP-MS

Presentations Turner K. M. and Owens J. D. (2017) A Geochemical Analysis of Rare Earth Elements Associated with Significant Sedimentary Phosphate Deposits of West- Central Florida. American Geophysical Union Fall 2017. New Orleans.

Turner K. M. and Owens J. D. (2018) A Geochemical Analysis of Rare Earth Elements Associated with Significant Sedimentary Phosphate Deposits of West- Central Florida. Southeastern Biogeochemistry Symposium 2018. Tallahassee.

55