<<

Establishing a Tephrochronologic Framework for the Middle () Type Area and Adjacent Portions of the and Northwestern Shelf, West and Southeastern New Mexico, USA

A dissertation submitted to the

Graduate School of the University of Cincinnati

In partial fulfillment of the requirements for the degree of

DOCTOR OF PHILOSOPHY (Ph.D.)

In the Department of Geology Of the McMicken College of Arts and Sciences

2011

by

Brian Lee Nicklen B.S. University of Nebraska-Lincoln, 2001 M.S. University of Cincinnati, 2003

Committee Members: Prof. Warren D. Huff (Ph.D.), Chair Prof. Carlton E. Brett (Ph.D.) Prof. Attila I. Kilinc (Ph.D.) Prof. J. Barry Maynard (Ph.D.) Dr. Gorden L. Bell Jr. Prof. Scott D. Samson (Ph.D.)

Abstract

Despite being recognized for many , bentonites in the Middle Permian

(Guadalupian ) type area of west Texas and southeastern New Mexico, have received little research attention. As these important deposits act as geologic timelines, they can be used as tools for long distance stratigraphic correlation and high-precision radioisotopic age dating. An important problem that these bentonites can address is the lack of temporal control for the Guadalupian. This is an important time interval for significant changes in

Earth’s climate and biodiversity that include events leading up to, and potentially including, the first pulse of a double-phase mass extinction at the end of the . Also needed is better temporal constraint for key Guadalupian global chemostratigraphic and geomagnetic markers. In light of this, the duration of the Guadalupian stages and boundary age estimations need to be updated in order to assess cause and effect of these significant events. Refinement of the regional stratigraphic framework of units that comprise the

Guadalupian type area is also necessary. Ambiguities are present regarding the shelf-to- basin relationships of units within the Capitan depositional system. Additionally, correlation of key sections of biostratigraphic and chronostratigraphic significance is needed. To address these problems, a tephrochronologic framework was established using apatite-based geochemical bentonite correlations in conjunction with high-precision zircon

U-Pb radioisotopic dating.

Electron microprobe analyses of apatite phenocrysts indicate that a geochemical fingerprint may be established for Guadalupian type area bentonties, allowing correlation of samples between localities. These correlations have linked several sections to the

i

location of the GSSP for the base of the at Nipple Hill in

National Park. This fingerprinting has also led to a new shelf-to-basin timeline based on the interpreted correlation of a sample from the shelfal to the basinal

Rader Member of the , providing biostratigraphic control for the bentonite within the Rader by extension.

The calculation of new U-Pb ID-TIMS dates in the Guadalupian type area indicate the need for changes to the and the temporal relationships of global events.

These new data suggest that current estimates of the -Capitanian boundary and probably the -Wordian boundary are too old. This extends the duration of the

Wordian to c. 3 myr and decreases the duration of the Capitanian to c. 4 myr. Results here also provide age estimates for the Illawarra geomagnetic reversal (c. 266.5 Ma) and the onset of the Kamura cooling event (c. 262.5 Ma). The latter provides constraint on the proposed within Guadalupian (Capitanian) mass extinction, placing it temporally closer to the main extinction pulse at the Permo- boundary than previously understood.

ii

iii

Acknowledgements

This dissertation would not have been possible without the help of several individuals and organizations. Each chapter has is own acknowledgements section, but I will use this space to elaborate and include some people who have supported me personally.

Warren Huff has been a tremendous advisor and I am thankful that I have been able to work so closely with him over the course of two graduate degrees. I am especially grateful for how understanding he has been regarding my work-life balance.

Committee members Carl Brett, Attila Kilinc, Barry Maynard, and Scott Samson are thanked for guidance over the past five years. They have also provided very useful and speedy feedback on drafts of the chapters that comprise this dissertation. Scott Samson is also acknowledged for allowing me to learn U-Pb radioisotopic techniques in his lab at

Syracuse University.

Gorden Bell is a member of my committee, but also a close collaborator. Any publication resulting from this work will have his name as a co-author. He went above and beyond his duties with the National Park Service and as a committee member, and I am greatly appreciative of his enthusiasm, encouragement and for many fruitful discussions of the various aspects of this study.

Bryan Sell is greatly appreciated for providing lodging and fine cooking in Syracuse,

New York and Saint-Cergue, Switzerland. He also graciously shared the unpublished (at the time) electron microprobe apatite data and interpretations from his dissertation. Bryan

iv

also provided advice on sample preparation, electron microprobe analysis, and U-Pb dating, for which I will be forever grateful.

I would also like to thank several individuals for providing encouragement and useful discussions along the way: Willis Tyrrell, Pete Holterhoff, Mitch Harris, Lance

Lambert and his graduate students.

The University of Cincinnati Department of Geology is a special place and I’m happy to have been a part of it for several years. I am thankful to my fellow graduate students, the faculty, and staff. A special thanks goes to Professor Emeritus Paul Potter for encouragement and advice over the course of my time at UC.

Urs Schaltegger graciously allowed me to use his laboratory at the University of

Geneva.

Funding for the project was provided by the American Association of Petroleum

Geologists (Ohio Geological Society Named Grant), the Clay Minerals Society, the Geological

Society of America, and the Graduate Student Governance Association and Department of

Geology at the University of Cincinnati. Each of these awards were specifically designated for graduate student research and I would like to express my gratitude to these organizations.

Finally, I would like to thank my family for all of their love and support; my parents, my in-laws, and my wife Keri. I hope this work is a good contribution to science, but my children Evelyn and William are the most important things I’ve created over the past 5 years.

v

Preface

In 1997, the Subcommission on Permian Stratigraphy published a time scale compilation to “encourage more rigorous studies on radiometric dating within biostratigraphically well constrained sections” (Permophiles Issue 34, p. 2). When the A

Global Time Scale 2004 was published (under the auspices of the International Commission on Stratigraphy), only one radioisotopic age date was available to provide internal constraint for the Middle Permian (Guadalupian Series) portion of the time scale. This date of 265.3 ± 0.2 Ma was reported to be from a bentonite 20 m below the base of the

Capitanian at its GSSP and is considered to be a maximum age estimate for the base of the (Bowring et al. 1998). While often being treated as occurring very near the base of the Capitanian, reports of its stratigraphic position have been inconsistent, casting doubt on how it should influence boundary age estimates for the three component stages of the

Guadalupian. As of this writing, no additional radioisotopic dates have been published, leaving this critical juncture in Earth’s history one of the least temporally constrained portions of the geologic time scale.

To address this lack of time control, a tephrochronologic framework was established using apatite-based geochemical bentonite correlations in conjunction with high-precision zircon U-Pb radioisotopic dating. In the process of establishing the framework issues relating to the lithostratigraphic, biostratigraphic, and sequence stratigraphic makeup of Guadalupian type were addressed.

Chapter 1 begins to establish this tephrochronologic framework by identifying and geochemically correlating bentonites in the regionally extensive Manzanita Member of the

vi

Cherry Canyon Formation. This framework links several sections to the Capitanian GSSP and constrains the lithostratigraphic position (Manzanita) and chronostratigraphic position (Wordian) for the type specimen of the cyclolobid ammonoid Newellites richardsoni. Also addressed is the proper stratigraphic placement of the radioisotopic date of Bowring and others (1998). Based on a review of published reports and a new measured section at the Capitanian GSSP presented here, the most accurate placement is in the Manzanita Limestone. This adjusts the only radioisotopic age constraint for the base of the Capitanian lower in the Wordian Stage than previously reported.

Chapter 2 addresses longstanding ambiguities in the shelf-to-basin correlation of the Rader Limestone Member of the Bell Canyon Formation. Results of apatite minor, trace, and rare earth element chemistry are interpreted to indicate a correlation between the lower portion of the Rader Limestone with the Yates Formation on the shelf. Specifically, the bentonite in the Rader is correlated to a bentonite in the Y3 high frequency sequence of

Osleger and Tinker (1999), establishing a new shelf-to-basin timeline for the Capitan depositional system. The bentonite in the Yates Formation was sampled from the PDB-04 research core that, through the new correlation, provides biostratigraphic constraint for the bentonite in the Rader that is radioisotopically dated in Chapter 3.

Chapter 3 presents the results of U-Pb ID-TIMS radioisotopic dating and new estimates for stage boundary ages in the Guadalupian. Also discussed are implications of this new temporal constraint on global events of the Guadalupian. Data presented here suggest that current estimates of the Wordian-Capitanian boundary and probably the

Roadian-Wordian boundary are too old. These new estimates extend the duration of the

Wordian to c. 3 myr and decrease the duration of the Capitanian to c. 4 myr. Results here

vii

also provide age estimates for the Illawarra geomagnetic reversal (c. 266.5 Ma) and the onset of the Kamura cooling event (c. 262.5 Ma). The latter provides constraint on the proposed within Guadalupian (Capitanian) mass extinction, placing it temporally closer to the main extinction pulse at the Permo-Triassic boundary than previously understood.

viii

Table of Contents

Abstract...... i Acknowledgements...... iv Preface...... vi Table of Contents ...... ix List of Figures...... xi List of Tables ...... xii List of Appendices...... xiii CHAPTER 1-TEPHROCHRONOLOGY OF THE MANZANITA MEMBER OF THE CHERRY CANYON FORMATION IN THE MIDDLE PERMAIN (GUADALUPIAN) TYPE AREA, WEST TEXAS AND SOUTHEASTERN NEW MEXICO, USA...... 1 ABSTRACT ...... 1 INTRODUCTION ...... 2 Geologic Background...... 7 Previous Work on Paleozoic Tephrochronology...... 10 METHODOLOGY ...... 12 Sample Collection and Preparation...... 12 Analytical Protocol ...... 14 RESULTS...... 15 Standards and Correction Factor...... 15 Manzanita Bentonites ...... 18 DISCUSSION ...... 22 Correlation of Bentonites...... 22 Complexity in the Data ...... 25 Implications ...... 26 Absolute Age of the Manzanita Bentonites...... 26 Stratigraphy of Nipple Hill...... 28 Back Ridge ...... 34 PDB-04 Research Core ...... 35 Constraining the Newellites richardoni Locus Typicus and Type Specimen...... 36 SUMMARY AND CONCLUSIONS...... 40 ACKNOWLEDGEMENTS...... 42 REFERENCES...... 43 CHAPTER 2-A NEW SHELF-TO-BASIN TIMELINE FOR THE MIDDLE PERMIAN (GUADALUPIAN) CAPITAN DEPOSITIONAL SYSTEM, WEST TEXAS AND SOUTHEASTERN NEW MEXICO, USA...... 49 ABSTRACT ...... 49 INTRODUCTION ...... 50 Background and Previous Work...... 54 Previous Shelf-to-Basin Bentonite Correlations...... 55 METHODS...... 56 Sample Collection and Prepartation...... 56 RESULTS...... 58 DISCUSSION ...... 60 Shelf-to-Basin Correlation...... 60 Biostratigraphic Control for Sample GM-20...... 63 CONCLUSIONS...... 66 ACKNOWLEDGEMENTS...... 67

ix

REFERENCES...... 68 CHAPTER 3-NEW ID-TIMS ZIRCON AGES FROM THE MIDDLE PERMIAN (GUADALUPIAN) TYPE AREA: IMPLICATIONS FOR THE GEOLOGIC TIME SCALE AND THE TIMING OF GLOBAL EVENTS...... 71 ABSTRACT ...... 71 INTRODUCTION ...... 72 Background...... 73 Global Events and Correlation Tools...... 74 METHODS...... 78 RESULTS...... 80 GM-20...... 80 GM-29...... 80 DISCUSSION ...... 82 Sample Age Interpretations...... 82 GM-20...... 82 GM-29...... 82 Stage Boundary Ages...... 84 Timing of the Kamura Event ...... 92 Timing of the Illawarra Reversal...... 93 CONCLUSIONS...... 94 ACKNOWLEDGEMENTS...... 95 REFERENCES...... 95 APPENDIX A ...... 101 APPENDIX B ...... 114 APPENDIX C ...... 118

x

List of Figures

CHAPTER 1 Figure 1.1 Regional geologic setting of the study area 3 Figure 1.2 Stratigraphic units in the Guadalupian type area 4 Figure 1.3 A) Locations of the PDB-04 research well. B) Outcrop and road cut sample localities. 5 Figure 1.4 Bivariate plots showing the uncorrected and corrected EMPA data for the Haldane K-bentonite collected at UKY with the Haldane EMPA data collect at SU 17 Figure 1.5 Mg-Mn-Cl and Mg-Mn-Ce/Y diagrams of the EMPA apatite data from the five bentonites at the Patterson Hills road cut. 20 Figure 1.6 A) Mg-Mn diagram for Patterson Hills road cut and several samples from other Guadalupian type area localities and the PDB-04 research core. B) Same diagram showing interpreted groups of correlative samples. 23 Figure 1.7 The three Manzanita bentonite correlation groups as Mg-Mn-Cl and Mg-Mn- Ce/Y diagrams 24 Figure 1.8 Measured sections showing bentonite sample positions and interpreted correlations. 29 Figure 1.9 A) Location of the locus typicus for the ammonoid Newellites richardsoni. The solid line near the locus typicus shows the outline of the small spur ridge, or point that juts from a southwestward trending questa. B) Measured section of the locus typicus. 38

CHAPTER 2 Figure 2.1 Schematic profile of the Capitan Reef depositional system showing traditional timelines used in shelf-to-basin correlations 51 Figure 2.2 Sequence stratigraphic framework of Osleger and Tinker (1999) 52 Figure 2.3 A) Location of the PDB-04 research well. B) Outcrop sample localities 53 Figure 2.4 Bivariate plots of the apatite minor, trace, REE EMPA data. 59 Figure 2.5 Gamma-ray log for the PDB-04 research well showing the placement of the Yates HFSs, informal Yates members, and the position of the two bentonites from this study. 61 Figure 2.6 Biostratigraphic control for GM-20 based on the correlation with PDB-04 1956 65

CHAPTER 3 Figure 3.1 Map of sample localities and Guadalupian GSSPs. 73 Figure 3.2 U-Pb Concordia and 206Pb/238U weighted mean diagrams for GM-20. 81 Figure 3.3 U-Pb Concordia and 206Pb/238U weighted mean diagrams for GM-29. 83 Figure 3.4 New stage boundaries age estimates and placement of key correlation events. 86 Figure 3.5 a,b Recalculated per-capita extinction and origination rates from Clapham and others (2009) for the Wordian and Capitanian stages 91

xi

List of Tables CHAPTER 1 Table 1.1 Sample Registry 13 Table 1.2 EMPA Analytical Protocol 15 Table 1.3 Interlaboratory correction factors. Based on measurements on apatite from the Late Deicke K-bentonite from Union Furnace, Pennsylvania. SU data from Sell (2010). 16 Table 1.4 Summary of stratigraphic positions published for 265.3 ± 0.2 Ma date of Bowring and others (1998). 28

CHAPTER 2 Table 2.1 Sample Register 57

xii

List of Appendices

Appendix A Apatite minor, trace, and REE EMPA data for Manzanita Limestone bentonite samples as weight percents. Smithsonian apatite standard (NMNH 104021) averages are from the SU sessions. 101 Appendix B Apatite minor, trace, and REE EMPA data for Yates and Bell Canyon formation bentonites as weight percents. Smithsonian apatite standard (NMNH 104021) averages are from the UKY sessions. 114 Appendix C U-Pb ID-TIMS isotopic data 118

xiii

Chapter 1

Tephrochronology of the Manzanita Member of the Cherry Canyon Formation in the Middle Permain (Guadalupian) Type Area, West Texas and Southeastern New Mexico, USA

Note: A modified version of this chapter will be submitted for peer-reviewed publication with co-authors G.L. Bell Jr. (United States National Park Service), L.L. Lambert (University of Texas at San Antonio) and W.D. Huff (University of Cincinnati)

ABSTRACT

Despite being recognized for many years, bentonites in the Middle Permian

(Guadalupian) type area of west Texas and southeastern New Mexico, have received little research attention. These important deposits offer opportunities for long distance stratigraphic correlation and high-precision radioisotopic age dating. In this study, apatite phenocryst chemistry is used to establish a tephrochronologic framework for the

Manzanita Member of the Cherry Canyon Formation. This framework is used to address current biostratigraphic, chronostratigraphic, and lithostratigraphic ambiguities and will serve as a basis for future studies on the stratigraphy and timing of the Guadalupian type area.

Samples were collected from bentonites at five field localities and one core, including Nipple Hill, which is the site of the Capitanian GSSP. Apatite phenocrysts from these samples were analyzed for minor, trace, and rare earth element chemistry using electron microprobe techniques. Results indicate the presence of three patterns or trends of data that are repeated at multiple localities. These groups of data are interpreted to represent coeval deposits and are correlated between several localities to form a

1

tephrochronologic framework. This framework links Nipple Hill with several other

Guadalupian type area localities.

Time control for the Guadalupian is poor and the precise stratigraphic position of the existing date from Nipple Hill of Bowring and others (1998) is uncertain. Based on a review of published reports and a new measured section of Nipple Hill presented here, the most accurate placement is in the Manzanita Limestone. This adjusts the only radioisotopic age constraint for the base of the Capitanian lower in the Wordian Stage than previously reported. The bentonite correlations proposed here constrain this position to the Mz-3 cycle of the Manzanita High Frequency Sequence and allow it to be traced to several outcrop localities and across the Delaware Basin in the subsurface.

The litho-and chronostratigraphic positions of the ammonoid Newellites richardsoni have been inconsistently reported in the literature and museum records. Based on the correlation of a bentonite from its relocated locus typicus near Casey’s Last Chance Well in

Culberson County, Texas, the type specimen is now constrained to the Wordian Manzanita

Limestone. A radioisotopic age has not been determined for this bentonite, but the tephrochronologic framework established here indicates it should coincide approximately with the date from Nipple Hill.

INTRODUCTION

Due to the geologically instantaneous nature of their deposition, widespread deposits of highly altered volcanic tephra (bentonites) can represent timelines, or isochrons, in the stratigraphic record. When coupled with geochemical information derived from their parent magma, bentonites are invaluable tools in geologic studies and have been used to address a diverse range of geologic problems, including those involving stratigraphic

2

correlation, absolute timing, and tectonomagmatic processes. Bentonites occur in several stratigraphic units of the Middle Permian (Guadalupian) type area in the Delaware Basin and Northwestern Shelf of west Texas and southeastern New Mexico (FIGURE 1.1), though little work has been reported on these deposits. In this chapter, bentonites are correlated between several localities to establish a tephrochronologic framework for the regionally extensive Middle Guadalupian (Wordian) Manzanita Limestone Member of the Cherry

Canyon Formation (FIGURE 1.2). This framework is based on geochemical data from apatite phenocrysts separated from the bentonites and links key sections of biostratigraphic and chronostratigraphic significance, including the Capitanian (Late

3

Guadalupian) Global Boundary Stratotype Section and Point (GSSP). Here this framework is used to address specific problems regarding ambiguities in Guadalupian ammonite biostratigraphy and existing temporal control for this portion of the geologic time scale. It also serves as a basis for ongoing work on the stratigraphy and timing of events in the

Guadalupian type area and the constraining of the GSSP boundary ages.

Guadalupe Mountains National Park (GMNP) contains the GSSPs for the three component stages of the Guadalupian Series (Glenister et al. 1999; Wardlaw, Davydov and

Gradstein 2004; FIGURE 1.3). Nipple Hill is the location of the GSSP for the Capitanian and also serves as the type locality for the Manzanita Limestone. The section at this locality extends only slightly above (0.6 m) the base of the stage- defining postserrata zone in the Pinery Limestone Member of the Bell Canyon Formation.

The only high-precision radioisotopic age date published for the Guadalupian comes from a sample collected at Nipple Hill, although the stratigraphic position has been inconsistently reported. There are also inconsistencies regarding the interpretation of stratigraphic intervals present at this locality. Establishing a tephrochronologic framework provides a

4

5

means to compare this section with others that extend further above the Capitanian boundary and as well as those that can be compared to Nipple Hill to help evaluate the stratigraphic succession at that locality.

In the southern portion of GMNP, a locality in the Patterson Hills referred to as Back

Ridge by Fall and Olszewski (2010) (FIGURE 1.3) has the only known section with bentonites above and below the base of the Capitanian. Linking this section to the

Capitanian GSSP via the Manzanita bentonites would provide an opportunity to extend the section at Nipple Hill further and evaluate the age assignment for the base of this stage by determining absolute dates above and below the boundary.

Though bentonites are a notable feature in the Manzanita, the number present at any given locality varies (King 1948, Hampton 1983). One locality along US Highway

62/180 has at least five bentonites in the Manzanita and is here referred to as the

Patterson Hills road cut. The bentonites occur here in stratigraphic succession and are used to test whether apatite phenocryst chemistry from samples known to be different are distinct. This is a key step in geochemical fingerprinting and needs to be demonstrated before the bentonites from multiple localities can be compared. This locality is also important because it appears to share some of the interval that has been reported variously at Nipple Hill. It was also studied in detail by Diemer and others (2006) and Tyrrell and others (2004, 2006), who placed it in a sequence stratigraphic framework. Bentonite correlations extending from this locality offer an opportunity to place additional sections into the framework.

The tephrochronologic framework will also be used to constrain the chrono- and lithostratigraphic position and the locus typicus for the type specimen of the ammonoid

6

Newellites richardsoni. Literature and museum record descriptions of the geographic location of the locus typicus for the holotype of N. richardsoni have been inconsistent, creating uncertainty regarding the exact position. Doubt is present with respect to the litho- and chronostratigraphic positions of the holotype. A bentonite in what appears to be the Manzanita Limestone has been sampled from what is interpreted to be the correct locus typicus of the ammonoid. A successful correlation of this sample to one from Nipple Hill, could constrain this important type specimen to the Manzanita Limestone and the Wordian

Stage.

Geologic Background

The Delaware Basin and Northwestern Shelf are located in the foreland of the

Marathon-Ouachita orogenic belt, with the former being the western sub-component of the

Permian Basin (FIGURE 1.1). During the Wordian, the Delaware Basin was encircled by a carbonate margin that was transitioning from a ramp to a rimmed shelf platform (Kerans and Tinker 1999). Located roughly 5o north of the paleoequator, the study area is thought to have been in an arid tropical (Walker et al. 1995) or subtropical climate zone (Ziegler,

Hulber, and Rowley 1997).

The Manzanita Limestone Member is the uppermost of three named carbonate units within the basinal Cherry Canyon Formation (FIGURE 1.2). Carbonate portions of the member are dominated by lithologies ranging from mudstone to fine-grained packstone

(Hampton 1983, 1989; Diemer et al. 2006). A transition from dolostone to limestone occurs roughly 20 km from the basin margin (King 1948; Newell et al. 1953; Hampton

1983), though unaltered limestone remains present at the top of the member in some sections. Siliciclastic intervals are present and consist of very fine-grained quartz

7

sandstones and siltstones (Hampton 1983, 1989). Due to the paucity of index , direct biostratigraphic data from the Manzanita are difficult to obtain, although this member is well constrained to the Wordian based on its position below the GSSP of the Capitanian at

Nipple Hill.

The correlation of the Manzanita to shelf strata is controversial (Harris and Saller

1999; Tyrrell et al. 2004, 2006). Some workers have considered it to be equivalent to the

Shattuck Member of the (e.g. Newell et al. 1953; Kerans and Tinker 1999;

Sarg, Markello and Weber 1999) and others have correlated it to the lowest Seven Rivers

Formation (e.g Tinker 1998; Rush and Kerans 2010). Determining the relationship between the Manzanita and shelf margin strata is also difficult, with most having suggested it is either just older than (e.g. Newell et al. 1953; Tyrrell et al. 2004, 2006) or equivalent to the lowermost Capitan Formation (e.g. Beaubouef et al. 1999; Kerans and Tinker 1999).

Deposition of the member is thought to have coincided with a sea level lowstand (Hampton

1983, 1989; Sarg 1989; Kerans and Tinker 1999).

In their subsurface and outcrop studies, Tyrrell and others (2004, 2006) placed the

Manzanita in one high frequency sequence (HFS) that they name after the member

(Manzanita HFS). This interpretation differs from Kerans and Tinker (1999) who placed the thick underlying sandstone and first limestone of the Manzanita in a separate HFS (G

15) from the rest of the member (G 16). The Manzanita HFS was divided by Tyrrell and others (2004, 2006) into five cycles (Mz-1 to Mz-5), extending from a thick (c. 30-60 m) basal sand to the top of the uppermost carbonate of the member. Each cycle is composed of siliciclastic deposits overlain by carbonates. The siliciclastics are thought to have been transported to the shelf edge by eolian processes and were subsequently deposited in the

8

basin by suspension fallout or turbidity currents during the lowstand portions of each cycle

(Fischer and Sarnthein 1988; Tyrrell et al. 2004, 2006). During highstand portions of the cycle, carbonates were deposited and are also interpreted to result from turbidity currents

(Hampton 1983, 1989; Loftin 1996; Tyrrell et al. 2004, 2006). In areas proximal to the western margin of the Delaware Basin, the Mz-3 and Mz-4 cycles are frequently amalgamated, with no lowstand siliciclastics at the base of the latter cycle (Tyrrell et al.

2004, 2006). Tyrrell and others (2004, 2006) showed that in the eastern Delaware Basin, the Manzanita HFS thickens and tends to have better developed cycles and they suggested that the amalgamation of cycles in the west may result from decreased sediment supply and/or positive structural relief.

Bentonites have been reported in several Guadalupian units of the Delaware Basin and adjacent shelf strata, including the Manzanita, where they have long been recognized in outcrop (e.g. King 1948; Dahl 1965; Newell et al. 1953; Terrell 1960; and Hampton 1983,

1989) and in the subsurface (e.g. Adams 1935; Wilde 1975; Todd 1976; Garber et al. 1989;

Tyrrell et al. 2004, 2006). Bentonites in the Manzanita can be recognized at several outcrops and road cuts in and around Guadalupe Mountains National Park, by their green color, which stands in contrast to the buff color of the adjacent carbonates. Although common in the member, both King (1948) and Hampton (1983; 1989) noted that the bentonites are not always present at a given locality.

The most notable bentonite occurrence is at the type locality of the Manzanita

Limestone at Nipple Hill in GMNP, where at least two bentonites are present (Nicklen

2003). This is significant because these two bentonites occur well below the GSSP for the base of the Capitanian Stage (Upper Guadalupian) (FIGURE 1.3). Another key Manzanita

9

bentonite locality is a road cut along US Highway 62/180 near the junction with State

Highway 54 (FIGURE 1.3). Nicklen (2003) interpreted at least four discrete beds in the

Manzanita Limestone at the road cut, although the current study identifies five. In their detailed measured section of this road cut, Diemer and others (2006) placed the lower four bentonites in the carbonate portion of the Mz-3 cycle of Tyrrell and others (2004, 2006).

With no major faults running through the exposed section (Diemer et al. 2006), it represents an ideal place to examine multiple bentonites occurring in stratigraphic succession. Also exposed at this locality are the uppermost sandstone of the Cherry Canyon

Formation and what is probably the Hegler Member of the overlying Bell Canyon

Formation. Based on subsurface gamma ray log data, Tyrrell and others (2004, 2006) also interpreted a gamma-ray spike in a thin carbonate in the basal Manzanita (top of their MZ-

1 cycle) as a bentonite, although they noted that this portion of the section is typically amalgamated or covered in the western side of the Delaware Basin.

King (1948) noted that the Manzanita bentonites were useful correlation tools and included them in his cross sections. Other workers have used these bentonites to assist in making correlations by combining biostratigraphic data with stratigraphic position (Wilde

1975) or by tying spikes in subsurface gamma-ray logs to bentonites observed in outcrop

(Tyrrell et al. 2004, 2006). The current study is the first to employ phenocryst-based geochemical data to establish Manzanita bentonite correlations.

Previous Work on Paleozoic Tephrochronology

Most previous work on Paleozoic tephrochronology has focused on the Ordovician of North America and Scandinavia. In a study of the Late Ordovician Deicke and Millbrig K- bentonites, Samson, Kyle and Alexander (1988) demonstrated the effectiveness of primary

10

apatite phenocrysts as correlation tools, by finding unique REE (rare earth element) patterns determined by inductively coupled plasma mass spectrometry (ICP-MS). This initial work was followed up by Samson and others (1995), who compared the ability of trace element chemistry determined by electron microprobe analysis (EMPA), REE chemistry, and 87Sr/86Sr ratios in apatites to discriminate between Ordovician K- bentonites from the Taconic Basin in central New York. The results of their study showed that the K-bentonites have characteristic Cl, Fe, Mg, and Mn concentrations and 87Sr/86Sr ratios, while the REE concentrations are less useful in chemical correlation. Subsequent investigations (Emerson et al. 2004; Mitchell et al. 2004; Carey, Samson and Sell 2009; Sell and Samson 2011) have further demonstrated the ability of apatite trace element chemistry collected by EMPA to discriminate between Ordovician K-bentonite beds at various scales. Using both EMPA and ICP-MS data from apatite from several Ordovician K- bentonites, Sell (2010) found that while some REEs where useful in discriminating between samples, the results did not conflict with the EMPA data.

Biotite phenocryst chemistry has also been used in studies Ordovician K-bentonites

(e.g. Haynes, Melson and Kunk 1995; Huff 2008), with the results in general agreement with apatite-based studies on the same deposits (Sell and Samson 2011), though interpretations of the biotite data have varied. Major and trace element analysis of volcanic glass is frequently used in studies of tephra (Lowe 2011 and references therein) and to a lesser extent in studies of Ordovician K-bentonites, where it occurs as inclusions in quartz (e.g. Delano et al. 1994; Mitchell et al. 2004). In a study using both melt-inclusion-bearing quartz and apatite phenocrysts, Adhya (2009) found that neither phase contradicted the other in the ability to chemically discriminate 36 discrete

11

Ordovician K-bentonites. While glass inclusions in quartz are shielded from alteration processes, apatite crystals in a tephra are exposed. The uncertainty regarding the effect of the alteration process on the geochemistry of apatite has recently been addressed by Sell and Samson (in press), who showed that trace element concentrations of apatite crystals from several Ordovician K-bentonites had similar concentration ranges to those of

Quaternary tephra deposits, suggesting a minimal effect on the trace element chemistry of apatites from highly altered tephras.

Regardless of the method used for correlating bentonites, the ability to distinguish between discrete deposits must be demonstrated for each study on each new set of beds. It is also important that samples from different localities are of the same age, either through direct evidence (e.g. biostratigraphic) or inference. This study focuses on apatite phenocryst chemistry, determined by EMPA, from several samples either from or suspected to be from the Manzanita Limestone. Although additional phenocrysts (e.g. biotite) are also present in some samples, clear and euhedral apatite occurs in relative abundance in all samples encountered.

METHODOLOGY

Sample Collection and Preparation

Bulk bentonite samples were collected from three outcrop localities and the PDB-04 core (FIGURE 1.3) using a pick and small trowel (outcrop) or a stainless steel spoon (core) that was cleaned between uses. Samples were placed in one-gallon resealable plastic bags and labeled. Care was taken to ensure that only material that was suspected to be bentonite was included. TABLE 1.1 summarizes the locality information and the sample

12

names. The bulk samples were allowed to soak in beakers for at least 48 hours, prior to partial disaggregation by stirring. In some cases, the use of a mortar and pestle was required to disaggregate, however, care was taken to crush, rather than grind the samples.

The disaggregated samples were then flushed of clay-sized particles and sieved with a dedicated 178 µm mesh screen following a slightly modified version of the methods described by Sell (2010). Heavy mineral (> 2.85 g/cm3) separation was performed using lithium heteropolytungstate. Grains for analysis were hand picked using a binocular microscope to ensure that the most euhedral and clear apatites were chosen. The apatites were mounted in one-inch epoxy rounds and exposed and polished using 4000 grit sand paper and 0.05 µm aluminum oxide. Crystals were mounted with their c-axis parallel to the polishing surface in an attempt to avoid issues with Cl analysis noted by Stormer and others (1993). Grain mounts were coated with carbon prior EMPA.

Table 1.1- Sample Registry

13

Analytical Protocol

Chemical analysis of apatite phenocrysts was conducted on a JEOL JXA-8600

Superprobe in the Department of Earth Sciences at Syracuse University (SU) and a Cameca

SX-50 in the Department of Earth and Environmental Sciences at the University of

Kentucky (UKY). Each point was analyzed for Cl, Fe, Mg, Mn, Ce, and Y using 15 kV accelerating voltage and a beam current of 60 nA, with a beam diameter of 1-2 µm. There is currently no internationally recognized standard for calibrating these elements for EMPA of apatite. Counting times and calibration standards were different for the two probes and are summarized in TABLE 1.2. Testing on Manzanita apatites demonstrated variable concentrations for Cl based on the duration and order it was analyzed in each run. Also observed during testing were changes in Mn concentration based on whether it came first or second during the analytical run for each spot. To avoid these issues, Cl and Mn were measured first at each spot (on separate spectrometers), with Cl having shorter counting times (TABLE 1.2). The Smithsonian apatite standard (NMNH 104021) was analyzed following each calibration to assess precision and as a check on relative accuracy.

Previously analyzed apatites from the Ordovician Deicke K-bentonite collected at Union

Furnace, Pennsylvania (B. Sell pers. comm. 2010), were used as an additional secondary standard during the UKY sessions and analyzed after each calibration.

14

Table 1.2- EMPA Analytical Protocol

RESULTS

Standards and Correction Factor

As expected, the inter-laboratory comparison shows different absolute values for the secondary standards, but the precision achieved in the two labs was similar and shapes and trends of the Deicke data were the same. In order to unify the two datasets, a correction factor was determined by comparing data on the apatite crystals from the

Deicke K-bentonite that were analyzed on both probes (TABLE 1.3). Apatites from the

Deicke have been shown to have a relatively narrow geochemical composition over a wide geographic and diagenetic range (Emerson et al. 2004; Carey, Samson and Sell 2009; Sell

2010). The Deicke apatite data used to calculate the correction factors were also collected using the JEOL JXA-8600 Superprobe at SU, following the same analytical protocol as shown in TABLE 1.2 (B. Sell pers. comm 2010).

The correction factors determined from the Deicke were tested on apatite from the widespread Ordovician Haldane K-bentonite (Sell 2010). These apatites were also previously analyzed at SU and have similar trace element and REE (i.e. Ce & Y)

15

Table 1.3- Interlaboratory EMPA correction factors

concentrations as the Manzanita samples (B. Sell pers. comm. 2010). After applying the

Deicke correction factor to the UKY Haldane data, they align well with the SU Haldane data

(FIGURE 1.4). The correction factor was then applied to samples B-5, PDB-04 4008, PDB-

04 4024, PDB-04 4081, and GM-13A, which were collected after the SU sessions and only analyzed at UKY. Samples GM-2 and GM-18 from UKY were also corrected and combined with data from these samples collected at SU. For both samples, data points from the SU sessions suggested a wider range of Cl concentration than indicated by the main cluster.

This was confirmed by the corrected UKY data, which are included to better illustrate the observed range of data.

The results here also show the heterogeneous nature of Cl in the Smithsonian apatite standard as reported by Adhya (2009) and Sell (2010). In calculations of precision for Cl, values that were conservatively deemed to represent a different zone (e.g. < 0.3 wt.

%) were omitted. This feature, along with the relatively small concentrations of some trace elements (i.e. very near minimum detection limits), particularly Mn, seem to render this standard inappropriate for determining inter-laboratory correction factors for EMPA, at least for the samples studied here. The Deicke apatites have trace element concentrations that are much closer to those from the Manzanita and compositional zoning appears to be minimal, though detailed imaging was not performed. Due to the geographically wide distribution of this bentonite, apatites from the Deicke are not necessarily difficult to

16

obtain, but are likely not readily available to most electron microprobe laboratories. To avoid some of the issues mentioned above regarding the Smithsonian apatite standard, Sell

(2010) recommended the Fish Canyon Tuff as a possible secondary standard. Though it has not been analyzed here, apatites from the Fish Canyon Tuff have similar concentration

17

ranges as the Deicke and are probably more readily available, as it is already used as a fission-track standard.

Manzanita Bentonites

The trace element assemblages of apatite phenocrysts from eight bentonite samples are reported in APPENDIX A. Most data points listed represent the average of three separate spots analyzed per grain, with two spots being averaged for some very small grains. Intra-crystal chemical variation was minimal, such that the calculated averages represent the range of values encountered for each sample. Each spot was analyzed for certain minor, trace and rare earth elements (Cl, Mg, Fe, Mn, Ce, and Y) and in select grains for major elements (Ca, P, and F). The minor and trace elements chosen for analysis have proven useful in discriminating between Paleozoic apatite phenocrysts from discrete bentonites by previous workers (Samson et al. 1995; Emerson et al. 2004; Mitchell et al.

2004; Adhya 2009; Carey, Samson and Sell 2009; Sell and Samson 2011). The other two non-major elements were analyzed because they represent a light REE (Ce) and a heavy

REE (Y) in the absence of an entire REE analytical suite. Ce and Y concentrations of apatites from granites show covariation (Sha and Chappell 1999) and Sell and Samson (2011) have demonstrated the usefulness of this in characterizing Ordovician K-bentonites. The range of concentrations for these samples is similar to that of apatite from relatively unaltered

Cenozoic volcanic rocks using the same microprobe and analytical protocol as this study

(Sell and Samson in press).

FIGURE 1.5 presents apatite phenocryst chemistry from the five bentonites sampled at the Patterson Hills road cut, as Mg-Mn-Ce/Y and Mg-Mn-Cl trivariate diagrams. These elements were chosen because they proved to be the best for discriminating the apatites

18

from these bentonites in a series of bivariate plots. In both plots, each bentonite appears to have a unique grouping of data points, with samples B-4 and B-5 having noticeably higher

Cl weight % and Ce/Y ratios. The remaining three samples are differentiated primarily by their Mg and Mn concentrations, although subtle variations in Cl content and Ce/Y ratios are present. There is some overlap among the samples, but each has what appears to be a unique cluster or trend of data. The two samples that show the most similarity in the various bivariate and trivariate diagrams examined are B-4 and B-5. While FIGURE 1.5 seems to clearly show that the two samples are not completely indistinguishable, there is enough overlap in data to suggest that they may share some components. The Mg-Mn-Cl diagram shows what appear to be 2 groups for B-4, with one having higher Cl values that plot with B-5. Student’s t-test was used to determine whether the overlapping B-4 subgroup is statistically distinct from the B-5 cluster. This type of basic statistical analysis tests the null hypothesis that the means of two or more sets of data are not statistically different. To do this, it is assumed that the sub-grouping of data for B-4 is geologically driven. B-4 was split into two subgroups based on the observed gap in Cl values. The difference between the means of these subgroups is greater than that of the precision of the

Cl analysis reported at the 2σ level, suggesting the observed sub-grouping is real. Placing the B-4 data in these subgroups also splits them into normally distributed components.

This is important, as normal distribution is an assumption for Student’s t-test. The normally distributed B-5 Cl data contains no obvious gaps and was directly compared with the overlapping B-4 subgroup. Results of the test yield a high F-value and a low p-value.

Therefore the null hypothesis that the B-4 subgroup and B-5 cannot be distinguished with

19

20

respect to their Cl concentrations can be rejected. As it only takes one element to discriminate between samples, it can be said that despite the overlap seen in bivariate and trivariate diagrams, the B-4 and B-5 apatites have chemical compositions that are statistically different.

Though the data plotted in FIGURE 1.5 are primarily from the SU sessions, separate aliquots of apatite were analyzed from these samples at UKY and indicate that the patterns observed here are representative of the within sample variability. One way to assess this is simply by noting any distributional changes in the pattern of a given sample as more data are added. When plotted on their own, data from the UKY samples mimic the patterns from the same samples, but different apatites, analyzed at SU. After correction, the UKY data can be added to the SU based plots, and each sample can be evaluated for changes to the pattern. This can also be checked by calculating the variance for plotted elements. A fully characterized sample should not have substantial increases in variance with the addition of new data. Except where noted above, these tests did not yield any new sample information and thus imply that the samples have been fully characterized.

Given that these five bentonites occur in stratigraphic succession at the same locality, these data show that apatite phenocryst chemistry from Manzanita samples known to be different, display as unique clusters or trending patterns. With this established, samples from other localities can be compared with those from the Patterson Hills road cut to test for the presence of correlative bentonites.

21

DISCUSSION

Correlation of Bentonites

When added to the bivariate plots used to assess the Patterson Hills road cut, samples from the other localities, thought to be the same age, can be evaluated for potential correlations. FIGURE 1.6a plots apatite data for these new samples, along with the B-1 through B-5, as an Mg vs Mn bivariate diagram. Five groupings of data are interpreted in

FIGURE 1.6b, with the general boundaries of the each drawn to illustrate the overall respective patterns. Three of these groups (B-1, B-2, and B-3) show multiple samples sharing the same overall shape. A good example of this is seen in the B-1 group, with each of the four component samples (B-1, GM-8, GM-19, PDB-04 4024) represented in all portions of the trend, most notably the apparent subgroups with relatively high Mn and low Mg concentrations. Data on additional apatites from these samples analyzed at UKY show the same subgroups and “dogleg” trend, which points toward this being a distinct feature of this group. This is also true of the B-2 and B-3 groups, where distinct features of the data are replicated between samples and are also reproduced in the UKY dataset.

These repeated patterns indicate chemical similarity between samples and this is interpreted to result from the samples originating from a coeval tephra deposit. When combined and evaluated together, these interpretations do not violate the Law of

Superposition and correlations between localities can be made.

Although the Mg vs Mn plot is displayed here, it is important to note that these data must be evaluated using all possible discriminating elements. It requires only differences in composition for one element to reject what otherwise appears to be a correlation.

FIGURE 1.7 illustrates data from certain samples in trivariate space, with Mg vs Mn as the

22

23

24

basal plane and Cl and Ce/Y used to define the z-axes. While these diagrams again are only representative of the complete dataset, they are useful in that they further highlight the chemical similarities between samples and demonstrate that the addition of a third variable does not pull any of the samples in the group away from the main pattern.

Complexity in the Data

Patterns such as this those from the diagrams in FIGURES 1.6 and 1.7 show that, for the most part, these samples do not all plot as tight clusters, but rather show complexity and evidence for multiple populations of apatite. This has been seen in other tephrochronologic studies and has been explained in various ways, including amalgamation of multiple tephra deposits, chemically zoned magma chambers, magma mixing, or the presence of antecrysts or xenocrysts (Emerson et al. 2004; Carey, Samson and Sell 2009; Lowe 2011; Sell and Samson 2011).

The “dogleg” trend of B-1 is comparable to that of the Ordovician Millbrig K- bentonite reported by Carey, Samson and Sell (2009), who note its similarity to that of apatite data from the Bishop Tuff (Hildreth 1979). They suggest that the observed trend results from either a chemically zoned magma chamber or the amalgamation of material from multiple eruptions. The presence of ashes representing multiple eruptions in the B-2 group, as well as the other Manzanita deposits, cannot be discounted and could potentially be resolved by collecting samples at different horizons within the bentonite bed. Physical features that may be indicative of composite ash falls or post depositional reworking that have been identified in some Paleozoic tephra deposits (e.g. Haynes 1994, Ver Straeten

2004) were generally not observed in these samples, though B-1 is underlain by very hard, chert-like material that may represent a separate tephra deposit.

25

It is also not possible with the current dataset, nor an objective of this study, to definitively determine whether processes affecting magma chamber chemistry influenced these samples, though isotopic evidence and additional REE information from apatite and other phenocrysts such as zircon, may potentially be used to address this issue (e.g.

Davidson, Tepley and Knesel 1998; Dempster et al. 2003; Schoene et al. 2010). What can be stated is the complexity seen here appears to be reproducible among different aliquots of apatite, confirming the uniqueness of the pattern. This implies greater confidence in the interpreted correlations, as it would be unlikely for a unique and complex pattern to be repeated in several non-correlative samples.

Implications

Absolute Age of the Manzanita Bentonites

At present, only one radioisotopic age date has been published for the Guadalupian interval (Bowring et al., 1998; Wardlaw, Davydov and Gradstein 2004; Ogg, Ogg and

Gradstein 2008). In an effort to improve time constraint on this interval, several

Guadalupian type area bentonites, including three at Nipple Hill, were sampled during this study and have been the focus of single zircon U-Pb isotope dilution thermal ionization mass spectrometry (ID-TIMS) dating. Unfortunately, results thus far have shown that the zircon population of these samples are either highly xenocrystic or yield unreliable dating information due to very low amounts of radiogenic lead.

The existing date of 265.3 ± 0.2 Ma by Bowring and others (1998) was reported to be from a bentonite at Nipple Hill in a position 2 m above the top of the Hegler Member of the Bell Canyon Formation and 20 m below the base of the J. postserrata conodont zone

26

(base of the Capitanian). It is considered to be a maximum age estimate for the base of the

Capitanian Stage and has been used in construction of the most recent versions of the

Geologic Time Scale by the International Commission on Stratigraphy (Wardlaw, Davydov and Gradstein 2004; Ogg, Ogg and Gradstein 2008). The calculated age is significant because it is nearly 10 myr older than previous estimates for the base of the Capitanian

(Ross & Ross 1987).

Although the date is apparently precise, the stratigraphic position of the bentonite bed has been inconsistently reported in the literature. TABLE 1.4 summarizes the various reports of the measured position for the dated bentonite. The measurement of 37.2 m below the base of the J. postserrata conodont zone seems to be the most consistent as it is reported in both Glenister and others (1999) and Wardlaw, Davydon and Gradstein (2004), with the former authors noting that their data supersedes that of the original description. If the 37.2 m distance is accepted, this positioning suggests that one of the two bentonites in the Manzanita at Nipple Hill may have been the source for the 265.3 ± 0.2 Ma date (FIGURE

1.8). This is notable because the date has frequently been treated as being at or very near the base of the Capitanian (e.g. Wardlaw, Davydov and Gradstein 2004; Korte et al. 2005;

Menning et al. 2006; Słowakiewicz, Kiersnowski and Wagner 2009). The original position given as 20 m below the base of the Capitanian is closer to the position of a bentonite in the first carbonate interval below the base of the Pinery Limestone, though still approximately

5 m below that point in the section (FIGURE 1.8).

Despite the lack of success thus far in using single zircon crystals, it is possible that one of the Manzanita bentonites at Nipple Hill could have yielded a precise ID-TIMS date.

The use of a multigrain technique with a large number of zircons, as described by Bowring

27

and Erwin (1998), may have overcome the problems encountered in single crystal analysis.

If this were the case, the geochemically-based bentonite correlations interpreted here would carry this precise date to several additional sections and to the PDB-04 research core. The date could further be carried across the Delaware Basin by the subsurface correlations of Tyrrell and others (2004, 2006).

Table 1.4 - Summary of stratigraphic positions published for 265.3 ± 0.2 Ma date of Bowring and others (1998)

Stratigraphy of Nipple Hill

The bentonite correlations in FIGURE 1.8 confirm the Patterson Hills road cut as being Manzanita by providing a high-resolution link to the type locality of the member at

Nipple Hill. They also provide a basis for direct comparison of the Manzanita HFS between the two sections. The correlations suggest that the Mz-1 through Mz-4 cycles are present at

Nipple Hill, with the thick sand at the base of the HFS present beneath the measured interval. As is typical in outcrop and subsurface sections in the western part of the

Delaware Basin (Tyrrell et al. 2004, 2006), the siliciclastic portions of cycles Mz-2 and Mz-4 are absent. Although it is not readily identifiable, the Mz-5 cycle may also be present, but obscured by the poorly exposed interval. The nodular limestone in the Mz-4 cycle at the

Patterson Hills road cut is in the same position relative to the bentonites at Nipple Hill and may be a good marker for the top of that cycle, though its presence at this position in additional sections should be demonstrated. Care should also be taken as the Hegler

28

Limestone has a similar appearance in some areas (King 1948). The absence of the three uppermost bentonites from the Patterson Hill road cut may be real and related to differential preservation. A variation in the number of bentonites in the Manzanita between localities would be consistent with the findings of previous workers (King 1948;

Newell et al. 1953; Hampton 1983). If this discrepancy is indeed the case, the correlations presented here are important because the relative stratigraphic positions of Bentonites #3-

5 from the road cut are well constrained between the basal bentonites and the interval of

29

nodular limestone in the upper portion of the Mz-4 cycle. The additional bentonites at the

Patterson Hills road cut may be more conducive to single-crystal U-Pb dating than those examined thus far (Chapter 2), offering another opportunity for high-precision age determinations in the Manzanita. It is also possible that B-3 through B-5 are present at

Nipple Hill, but recessed to a degree that their identification in a weathered outcrop profile is difficult. Additional inspection of this interval may reveal the upper bentonites. For example, Hampton (1983) identified one bentonite in the Manzanita at Nipple Hill, while the section measured for this work in FIGURE 1.8 demonstrates at least two. In either case, the results here show that a potential remains for determining a high-precision radioisotopic date from bentonites in the Manzanita that can be clearly related to the

Capitanian GSSP.

These bentonite correlations also provide a datum from which the upper portions of the sections can be evaluated. The measured section of Nipple Hill (FIGURE 1.8) has been compared with those of Hampton (1983) and Beck (1967). Despite some differences in the thickness of the first carbonate interval, there is good enough agreement between the sections to allow direct comparison of specific intervals. One notable feature in each section is a carbonate interval present between the top of the nodular limestone of the

Manzanita and the base of the Pinery Limestone. This interval was interpreted by

Hampton (1983) and Newell and others (1953) to be the Hegler Limestone Member of the

Bell Canyon Formation, but was considered to be Manzanita by Beck (1967). Though this interval was not specifically commented on, King (1948) noted the lack of Hegler at Nipple

Hill and other localities in the Guadalupe Mountains. This discrepancy is noteworthy because the 265.3 ± 0.2 Ma date of Bowring and others (1998) is frequently reported as

30

being in or very near the Hegler (e.g. Jin, Shang and Cao 2000; Menning et al. 2006; Shen et al. 2010).

The type locality of the Hegler is on private land and was not visited in the current study, but the member is described by King (1948) as being dark-gray fine-grained limestone, with poorly preserved ammonoids and trace fossils on some bedding plans.

What has been called “Hegler” at Nipple Hill does not resemble Hegler at more accessible localities in the area, such as the Pinery type locality near Pine Spring, where many features typical of the member are present. The interval in question at Nipple Hill appears, at least in outcrop, much more like the Manzanita. It has the same surficial color (orange-buff), lithology (dolostone), general lack of fossils, and common calcite-spar lined vugs. Hampton

(1983) considered the interval in question to be dolostone and noted that a dolomitized interval resembling the Manzanita is present in this position in several sections and suggests it is part of the Hegler based on King (1948). This similarity described by King

(1948), however, refers to the “dark-gray, lumpy limestone” that is typical of the Hegler at its type locality and other areas. King (1948) did make mention of a change from dark- to light-gray limestone in the Hegler, but does not identify anything that resembles the interval in question. Newell and others (1953) showed the Hegler as dolostone in sections roughly 6 km east of Nipple Hill, though they mapped it as limestone nearer to the basin margin.

The tephrochronologic correlations in FIGURE 1.8 demonstrate that the Hegler from the Patterson Hills road cut is in the same approximate position as the “Hegler” at Nipple

Hill. The thickness of the latter is also roughly the same as that reported by King (1948) for the typical Hegler. These observations suggest that the interval in question between the

31

Pinery and Manzanita members at Nipple Hill is either a dolomitized portion of the Hegler or an equivalent unit in the same relative stratigraphic position. If the Hegler is missing at

Nipple Hill, as King (1948) and Beck (1967) indicated, the interval could either represent a carbonate with Manzanita characteristics, but in the typical position of the Hegler or, potentially, the carbonate portion of the Mz-5 cycle. The bentonite correlations demonstrate close alignment between the top of the nodular at Nipple Hill and the Patterson Hills road cut. This would seem to rule out the “Hegler” interval as being the missing carbonate for the Mz-5 cycle, though the bentonite correlations also show some variability between the sections, with the Mz-3 siliciclastics having different thicknesses.

Tyrrell and others (2004, 2006) have shown that the thickness of the siliciclastic portions of the Mz-5 cycle also varies between sections, with the differences being on the order of what is seen between the Patterson Hills road cut and the poorly exposed interval at Nipple

Hill.

Although physical appearance and relative stratigraphic position offer a basis for interpreting the identity of the “Hegler” interval, biostratigraphic data or a direct link to a known Hegler section could provide positive identification. The FAD of the fusulinacean

Polydiexodina occurs in the Hegler and identification of such would strongly suggest the presence of the member at Nipple Hill. The current study has not identified Polydiexodina at this locality and is not aware of such identification by previous workers. Clifton (1944) mentioned collecting fusulinid samples from the Manzanita at Nipple Hill, though no species name(s) or measured section were included. Although it would allow for a more definitive stratigraphic assignment, it may not be possible as fusulinids are noted as being

32

rare in both the Manzanita and Hegler throughout the study area (King 1948; Hampton

1983).

King (1948) noted that seams of thin (< 2.5 cm) white clay are interbedded with the

Hegler at the junction of Cherry and Lamar Canyons and interpreted these to be volcanic tuff. The sample collected for this study from a recessive interval in the “Hegler” has volcanic phenocrysts, but does not match the description of King (1948). Despite this difference in physical appearance, a correlation should not be ruled out in the absence of geochemical data. If samples of the clay seams from known Hegler localities can be obtained, it may be possible to confirm the identity of this interval through tephrochronologic correlation.

Without direct biostratigraphic or tephrochronologic confirmation, the identity of the carbonate interval at Nipple Hill between the top of the Mz-4 cycle and the base of the

Pinery remains unclear. The bentonite correlations between the Patterson Hills road cut and Nipple Hill suggest that the interval is either Hegler, a roughly Hegler equivalent interval with characteristics more similar to the underlying Manzanita, or possibly the carbonate portion of the Mz-5 cycle. In light of this, future workers should note this uncertainty and properly justify the use of any specific lithostratigraphic or sequence stratigraphic terminology. Most models have placed the Manzanita and Hegler in different

HFSs (Tinker 1998; Kerans and Tinker 1999; Tyrrell et al. 2004, 2006). These HFSs are used in interpreting shelf-to-basin relationships and when correlating an event on the shelf to the Capitanian GSSP. It is therefore desirable to know which units are present, or to at least be aware that there is uncertainty regarding the succession.

33

As with all GSSPs, it is important to identify the relationship between the boundary and key markers used in correlation. An example of this is geomagnetic polarity. It is possible that the magnetic carriers in the basinal Guadalupian stratotype beds were dissolved during times of heavy oil hydrocarbon saturation, making the equivalent shelfal strata the only option for obtaining good paleomagnetic data (Steiner 2006). With no conodont data available on the shelf, the relationship between the Capitanian GSSP and geomagnetic polarity is achieved through shelf-to-basin correlations. Any events or distinctive intervals in the Hegler equivalent portions of the shelf would need to be related to the GSSP with caution.

Back Ridge

Another notable locality has the only known section with bentonites both above and below the base of the Capitanian. This locality was named Back Ridge by Fall and

Olszewski (2010) and is in the Patterson Hills in the southern portion of GMNP (FIGURE

1.3). Samples from two bentonites in this section have been correlated based on apatite phenocryst chemistry (FIGURES 1.6 and 1.7). These correlations place the bentonites in the Mz-3 cycle and provide direct links to Nipple Hill and the Capitanian GSSP (FIGURE

1.8). A third bentonite is in the Rader Limestone Member of the Bell Canyon Formation at the top of the section and is the focus of shelf-to-basin correlations and high-precision U-Pb zircon dating (Chapters 2 & 3). Though not shown in FIGURE 1.8, a siliciclastic interval beneath the lowest bentonite may be present as it is in other localities, but obscured here by poor exposure. The measured section present here (FIGURE 1.8) is partial, but provides a datum from which to investigate the rest of the section.

34

The stratigraphy at Back Ridge appears to be atypical in that the Pinery Limestone seems to directly overly the Manzanita, though this is consistent with some measured sections in King (1948, plate 6) and could indicate a location more proximal to the shelf margin during deposition. A detailed biostratigraphic framework for this section is currently being established (L.L. Lambert pers comm. 2011) and may shed additional light on the stratigraphic relationships of this locality. While this locality may not be representative of a complete basinal section, high-precision radioisotopic dates above and below the base of the Capitanian would still provide an important constraint on the age of that boundary.

PDB-04 Research Core

The correlations here provide a tephrochronologic link between outcrop sections and the PDB-04 research core from Eddy County, New Mexico (FIGURE 1.3a). The core was studied in detail by Garber, Grover and Harris (1989), who identified bentonites in the

Manzanita and Yates formations. For this study, suspected bentonites were sampled at

4008 ft, 4024 ft, and 4081 ft core depths. The correlations shown in FIGURE 1.6 identify the samples from 4008 ft and 4024 ft as the B-3 and B-1 bentonites in the Mz-3 cycle at the

Patterson Hills road cut (FIGURE 1.8), demonstrating that at least one additional bentonite from the Patterson Hills road cut can be identified in more than one section. If the 265.3 ±

0.2 Ma date of Bowring and others (1998) is indeed from the Manzanita as discussed above, the correlations proposed here would carry that high-precision date to the research core and provide an absolute age datum at a position at or near 4024 ft core depth. These correlations also place the sample from 4081 ft lower in the Manzanita. It probably represents the bentonite interpreted at the top of the Mz-1 cycle by Tyrrell and others

35

(2004, 2006) in several subsurface sections. The apatite trace element chemistry for this sample is similar to that of the B-1 cluster, but the Ce/Y ratios distinguish it from the group

(APPENDIX A).

Constraining the Newellites richardoni Locus Typicus and Type Specimen

N. richarsoni is a monotypic cyclolobid ammonoid, that was originally named

Waagenoceras richardsoni by Plummer and Scott (1937). The new genus Newellites was proposed by Furnish and Glenister in Davis, Furnish and Glenister (1969), with W. richardsoni as its type specimen. Thought to be an advanced variety of the ammonoid

Waagenoceras dieneri, Miller and Furnish (1940) suggested that the genus is the least derived in a phylogenetic sequence with Timorites and Cyclolobus. Since assignment to a new genus, N. richardsoni has been considered to be an advanced offshoot of Waagenoceras

(Lambert and Glenister 2002; Leonova 2010).

Uncertainty exists regarding the exact position of the locus typicus for the holotype of N. richardsoni, as descriptions of the geographic location have been inconsistent. Though all records include Casey’s Last Chance well as a reference point, the direction of the locus typicus from the well varies. The original description by Plummer and Scott (1937, p. 160), stated that the locus typicus is southeast of the well near the head of Chico Draw. In contrast, the text of Miller and Furnish (1940, p. 173), placed the locus typicus south of the well, while a figure caption on the preceding page states the direction is northeast.

Inconsistencies are also present in the two record cards for this locality (BEG 55-T-3) from the Texas Memorial Museum, with one placing the locus typicus south of the well and the other southeast (A. Molineux pers. comm. 2007). In their designation of the genus

Newellites, Furnish and Miller (1969) placed the locus typicus southeast of the well.

36

In addition to inconsistencies regarding the exact location of the N. richardsoni locus typicus, some doubt is present with respect to the litho- and chronostratigraphic positions of the holotype. Though the majority of workers have placed the N. richardsoni holotype in the Manzanita Member of the Cherry Canyon Formation (e.g. Furnish and Glenister in

Davis, Furnish and Glenister 1969; Glenister, Furnish and Zuren 2009), the museum catalog card for the specimen infers it is from the Hegler Member of the Bell Canyon Formation.

This is may result from the presence of small normal faults in the area that place the otherwise overlying Hegler into close proximity to the Manzanita. It is also possible that this is related to the suggestion by Miller and Furnish (1940, table 6) that specimens similar to the holotype from the mouth of True Canyon, may be from the Hegler. The early studies by Plummer and Scott (1937) and Miller and Furnish (1940) predate the formal designation of the Manzanita Member by King (1942) and give the position of the holotype as the Delaware Mountain Formation, though later authors suggests it is from the “lower

Capitan horizon”. In the most recent placement of the holotype, Leonova (2010) stated that it is from the Capitan Formation and gives the age as Capitanian, rather than Wordian as it is elsewhere.

Using all available data, an attempt has been made to relocate the locus typicus and determine its correct litho- and chronostratigraphic positions. Based on several days of field investigations in the vicinity of the published localities, the most accurate position for the N. richarsoni locus typicus is now identified as a small spur ridge, or point, jutting north of a southwestward trending cuesta, 400 m southeast of Casey’s Last Chance Well (FIGURE

1.9). Extrapolation of the geologic map of King (1948, plate 3) suggests the cuesta that holds up this location should be Manzanita Limestone. Near the top of the spur, a 11.5 cm

37

38

layer of green bentonite forms a recessive notch in a sequence of mudstone/wackestone and calcareous siltstone (FIGURES 1.8 and 1.9). Below the bentonite, several cyclolobid ammonoids are present.

A sample from the bentonite (GM-25) proved to be a member of the B-2 correlation group (FIGURE 1.6b); it thus correlates to the younger of the two bentonites at Nipple Hill.

With Nipple Hill being the site of both the Capitanian GSSP and the type locality of the

Manzanita, this correlation constrains the lithostratigraphic position (Manzanita) and chronostratigraphic position (Wordian) of the N. richardsoni type specimen (FIGURE 1.8).

This correlation also places the top of the locus typicus in the MZ-3 cycle. The exact position of the type specimen within the measured section is not known. Cyclolobid ammonoids are recognized in the intervals below the bentonite (FIGURE 1.9) and the type specimen is likely from the Mz-2 or lower Mz-3 cycle, although all that can be said for certain is that the type specimen is not from the Mz-4 or Mz-5 cycles. The Mz-1 carbonate is commonly thin (Tyrrell et al. 2004, 2006) and might be represented at the base of the section. It is difficult to make a precise determination without knowing the total thickness of the sand at the base of the exposure.

A notable difference between this locality and the others studied is the prevalence of limestone. This is not surprising given the basinward transition from limestone to dolostone in the Manzanita as mentioned above. The other samples that fall into the B-2 group were collected from dolostone intervals, including a sample from the Nickel Creek road cut. This suggests that the dolomitization process had little to no effect on the chemistry of the apatite crystals of these samples. As bentonite correlations expand to the regional and even continental or intercontinental scale, the need to reliably compare

39

samples from different lithologies and diagenetic environments becomes increasingly important. Although not definitive, results here are encouraging and further support the robust nature of apatite chemistry as a reliable fingerprint in highly altered deposits.

SUMMARY AND CONCLUSIONS

This study demonstrates the utility of apatite phenocryst chemistry in correlating bentonites from the Manzanita Limestone in the Guadalupian type area. Results from the

Patterson Hills road cut indicate that samples from each of the five bentonites occurring in stratigraphic succession plot as unique clusters or trends in bivariate and trivariate diagrams. Discriminating between most samples is relatively straightforward, although, in some cases, the use of basic statistical methods is useful in elucidating less conspicuous differences. Despite complexity and relatively small differences in minor, trace, and REE concentrations, the patterns and trends of data are repeated in several samples from other localities. These repetitions are interpreted to represent groups of coeval deposits, which are treated as isochrons and correlated between localities to provide a high-resolution tephrochronologic framework. The collection of additional geochemical datasets from apatite. It might also provide insight into the complexity in the data observed in some of the bivariate and trivariate diagrams.

The results presented here show that despite having different calibration standards and analytical protocols, apatite phenocryst data acquired from two separate microprobes can be compared after applying a correction factor based on apatite from the Deicke K- bentonite analyzed at each laboratory. The Smithsonian apatite standard was used to assess accuracy and precision, but may not be appropriate for determining correction factors. Other, more widely available, secondary standards, such as apatite from the Fish

40

Canyon Tuff may also be useful, though this has not been evaluated here. Care must be taken to test the calculated correction factors on any standard, but these results are encouraging as they suggest that, even at the trace levels encountered here, apatite data from different studies can be compared if the appropriate corrections are determined.

A review of published reports on the 265.3 ± 0.2 Ma date of Bowring and others

(1998), indicates the most accurate position is probably in the Manzanita Limestone, rather than a point at or very near the base of the Capitanian as has been suggested. This position roughly coincides with two of the Manzanita bentonites that have been correlated to several other localities. These correlations place the age date in the Mz-3 cycle. This cycle is easily identified in wireline logs and can be traced across the Delaware Basin, providing a subsurface absolute age datum. Correlation of the Patterson Hills road cut to Nipple Hill links three additional bentonites to the Capitanian GSSP. These could potentially be dated and would provide a check on the position of the Bowring and others (1998) age and could allow for additional constraint of the timing of the Wordian and Capitanian stages. The correlations between these two localities also provide some insight into whether the

Hegler Limestone is present at Nipple Hill, but conclusive evidence is not currently available. Further age constraint on the base of Capitanian might be possible at Back Ridge in the Patterson Hills, where bentonites samples that correlate to the Manzanita at Nipple

Hill occur with a bentonite above the boundary in the Rader Member of Bell Canyon

Formation.

The bentonite at the N. richardsoni type locality correlates to a bentonite in the

Manzanita Limestone at the Patterson Hills road cut and the type locality of the Manzanita at Nipple Hill. This confirms the litho-(Manzanita) and chronostratigraphic (Wordian)

41

positions of type specimen from Glenister, Furnish, and Zuren (2009). In light of this confirmation, the phylogenetic reconstruction of the family Cyclolobidae of Leonova (2010) should be adjusted. This correlation also suggests that by linking together sections that are comprised of both limestone and dolostone, the dolomitization processes did not affect the apatite phenocryst chemistry.

ACKNOWLEDGEMENTS

I would like to thank the National Park Service for granting permission to collect samples from within Guadalupe Mountains National Park under permit # GUMO-2007-SCI-

0010. V. Moyers and C. Breen are thanked for securing access to the N. richardsoni locus typicus and for guiding G. Bell to previously scouted ammonoid localities and identifying

Casey’s Last Chance Well. B. Sell provided much appreciated assistance with EMPA at

Syracuse University and many stimulating discussions on apatite phenocryst chemistry. He is also thanked for providing mounted apatites from the Deicke and Haldane K-bentonites.

D. Moecher and S. Chakraborty provided access to the Cameca SX-50 at the University of

Kentucky. W. Tyrrell shared his knowledge of the regional subsurface stratigraphy of the

Delaware Mountain Group. T. Phillips helped with drafting some of the figures. N. Bose, C.

Ferguson, and J. Wilkins assisted with sample collection. M. Harris graciously shared information regarding the PDB-04 core and P. O’Neill (Louisiana Geological Survey) and J.

Donnelly (Austin Core Research Center) are thanked for providing access. Graduate student research funding was provided by the American Association of Petroleum

Geologists (Ohio Geological Society Named Grant), the Clay Minerals Society, the Geological

Society of America, and the Department of Geology at the University of Cincinnati.

42

REFERENCES Adams, J.E., 1935. Upper Permian stratigraphy of west Texas Permian Basin. American Association of Petroleum Geologists Bulletin, 19:1010-1022.

Adhya, S., 2009, “Geochemical fingerprinting of volcanic airfall deposits, a tool in stratigraphic correlation” [unpublished Ph.D. dissertation]: The State University of New York, Albany, New York, 532 p.

Beaubouef, R.T., Rossen, C., Zelt, F.B., Sullivan, M.D., Mohrig, D.C. and Jennette, D.C., 1999. Deep-water sandstones, Brushy Canyon Formation, west Texas: American Association of Petroleum Geologists Continuing Education Course Notes Series #40, 62p.

Beck, R.H., 1967, Depositional mechanics of the Cherry Canyon Formation, Delaware Basin, Texas [unpublished Master’s thesis], Lubbock, Texas, Texas Technical College, 107p.

Bowring, S.A. and Erwin, D.H., 1998. A new look at evolutionary rates in deep time: uniting paleontology and high-precision geochronology. GSA Today, 8:1-8.

Bowring, S.A., Davidek. K., Erwin, D.H., Jin, Y.G., Martin, M.W. and Wang, W., 1998. U-Pb zircon geochronology and tempo of the end-Permian mass extinction. Science, 280:1039-1045.

Carey, A., Samson, S.D. and Sell, B, 2009. Utility and limitations of apatite phenocryst chemistry for continent-scale correlation of Ordovician K-bentonites. Journal of Geology, 117:1-14.

Clifton, E.L., 1944. Ammonoids from upper Cherry Canyon Formation of the Delaware Mountain Group in Texas: American Association of Petroleum Geologists Bulletin, 28:1644-1646.

Dahl, H.M., 1965, Clay mineralogy of some Permian bentonites from the Delaware Basin area, Texas. The American Mineralogist, 50:1637-1646.

Davidson, J., Tepley, F. and Knesel, K., 1998. Isotopic fingerprinting may provide insights into evolution of magmatic systems. Eos Transactions of the American Geophysical Union, 79:185-193.

Davis, R.A., Furnish, W.M. and Glenister, B.F., 1969. Mature modification and dimorphism in Late Paleozoic ammonoids. In: Ernst, G., Westermann, G., Eds. Sexual dimorphism in metazoa and taxonomic implications. Symposium organized by the International Palaeontological Union, Committee on Evolution, Prague 1968. Stuttgart: Schweizerbart. International Union of Geological Sciences, no. 1.

Delano, J. W., Tice, J. W., Mitchell, C. E. and Goldman, D., 1994. Rhyolitic glass in Ordovician K-bentonites: A new stratigraphic tool. Geology, 22:115–118.

43

Dempster, T.J., Jolivet, M., Tubrett, M.N. and Braithwaite, C.J.R., 2003. Magmatic zoning in apatite: A monitor of porosity and permeability. Contributions to Mineralogy and Petrology, 145:568–577.

Diemer, J.A., Tyrrell, W.W. Jr., Bell, G.L. and Griffing, D.H., 2006. A Patterson Hills Section of the Bentonite-Bearing Manzanita Limestone of the Cherry Canyon Formationa, Culberson County, Texas, in Hinterlong, G., ed. Permian Basin Section SEPM, Publication 2006-46.

Emerson, N.R., Simo, J.A., Byers, C.W. and Fournelle, J., 2004. Correlation of (Ordovician, Mohawkian) K-bentonites in the upper Mississippi valley using apatite chemistry: implications for stratigraphic interpretation of the mixed carbonate-siliciclastic Decorah Formation. Palaeogeography, Palaeoclimatology, Palaeoecology, 210:215- 233.

Fall, L.M. and Olszewski, T.D., 2010. Environmental disruptions influence taxonomic composition of paleocommunities in the Middle Permian Bell Canyon Formation (Delaware Basin, west Texas). Palaois, 25:247-259.

Fischer, A.G. and Sarnthein, M., 1988. Airborne silts and dune-derived sands in the Permian of the Delaware Basin. Journal of Sedimentary Petrology, 58:637-643.

Garber, R.A., Grover, G.A. and Harris, P.M., 1989. Geology of the Capitan Shelf Margin- subsurface data from the northern Delaware Basin, in Harris, P.M. and Grover, G.A., eds., Subsurface and outcrop examination of the Capitan Shelf Margin-a Core Workshop, 3-269. Tulsa: Society of Economic Paleontologists and Mineralogists Core Workshop 13.

Glenister, B.F., Boyd, D.W., Furnish, W.M., Grant, R.E., Harris, M.T., Kozur, H., Lambert, L.L., Nassichuk, W.W., Newell, N.D., Pray, L.C., Spinosa, C., Wardlaw, B.R., Wilde, G.L. and Yancey, T.E., 1992. The Guadalupian: proposed International Standard for a Middle Permian Series. International Geological Review, 34:857-888.

Glenister, B.F., Furnish, W.M. and Zhou, Z., 2009. Cycloloboidea. In: Selden, P. Ed.,. Treatise on Invertebrate Paleontology: Part L (Revised) , 145-157. Boulder- Lawrence: Geological Society of America and University of Kansas Press.

Hampton, B.D., 1983. Carbonate sedimentology of the Manzanita Member of the Cherry Canyon Formation (Middle Guadalupian, Permian) Guadalupe Mountains, west Texas: [unpublished Master’s thesis], Madison, Wisconsin, University of Wisconsin, 178p.

Hampton, B.D., 1989. Carbonate sedimentology of the Manzanita Member of the Cherry Canyon Formation in Harris, P.M., and Grover, G.A., eds., Subsurface and outcrop examination of the Capitan Shelf Margin-a Core Workshop, 431-439. Tulsa: Society of Economic Paleontologists and Mineralogists Core Workshop 13.

44

Harris, P.M. and Saller, A.H., 1999. Subsurface expression of the Capitan depositional system, in Saller et al., eds., Geologic Framework of the Capitan Reef, 37-49. Tulsa: SEPM Special Publication No. 65.

Haynes, J. T. 1994. The Ordovician Deicke and Millbir K-bentonite beds of the Cincinnati Arch and the southern Valley and Ridge Province. Boulder: Geological Society of America Special Paper 290, 80 pp.

Haynes, J.T., Melson, W.G. and Kunk, M.J., 1995. Composition of biotite phenocrysts in Ordovician tephras casts doubt on the proposed trans- correlation of the Millbrig (United States) and the Kinnekulle K-bentonite (Sweden). Geology, 23: 847-850.

Hildreth, W. 1979. The Bishop Tuff: evidence for the origin of compositional zonation in silicic magma chambers. Geological Society of America Special Paper 180:43–75.

Huff, W.D., 2008. Ordovician K-bentonites: Issues in interpreting and correlating ancient tephras. Quaternary International, 178:276–287.

Jin, Y., Shang, Q. and Cao, C., 2000. Late Permian magnetostratigraphy and its global correlation. Chinese Science Bulletin, 45:698-705.

Kerans, C.,and Tinker, S.W., 1999. Extrinsic controls on development of the Capitan Reef complex, in Saller et al., eds., Geologic Framework of the Capitan Reef, 15-36. Tulsa: Society of Economic Paleontologists and Mineralogists Special Publication No. 65.

King, P.B., 1948. Geology of the southern Guadalupe Mountains, Texas. U.S. Geological Survey Professional Paper 215, 183pp.

Korte, C., Jasper, T., Kozur, H.W. and Veizer, J., 2005. δ18O and δ13Ccarb of Permian : a record of seawater evolution and continental glaciation. Palaeogeography, Paleoclimatplogy, Palaeoecology, 224:333–351.

Lambert, L.L. and Glenister, B.F., 2002. Ammonoid zonation of the type Guadalupian (Middle Permian Series)—Analyses of zonal boundaries versus chronostratigraphic boundaries and their historical contexts. Geological Society of America Abstracts with Programs. 34(3):30.

Leonova, T.B., 2010. Revision of the Permian Ammonoid Family Cyclolobidae: Paleontological Journal, 44:267-274.

Loftin, T.R. Jr., 1996. Depositional stacking patterns within the Cherry Canyon Formation, Delaware Basin, west Texas, in W.D. De Mis and A.B. Cole, eds. The Brushy Canyon Play in Outcrop and Subsurface: Concepts and Examples, 137-145. Permian Basin Section Society of Economic Paleontologists and Mineralogists, Publication 96-38.

45

Lowe, D.J., 2011. Tephrochronology and its application: A review. Quaternary Geochronology, 6:107-153.

Menning, M., Alekseev, A.S., Chuvashov, B.I., Davydov, V.I., Devuyst, F.-X., Forke, H.C., Grunt, T.A., Hance, L., Heckel, P.H., Izokh, N.G., Jin, Y.-G., Jones, P.J., Kotlyar, G.V., Kozur, H.W., Nemyrovska, T.I., Schneider, J.W., Wang, X.-D., Weddige, K., Weyer, D. and Work, D.M., 2006. Global time scale and regional stratigraphic reference scales of Central and West Europe, East Europe, Tethys, South , and North America as used in the --Permian Correlation Chart 2003 (DCP 2003). Palaeogeography, Palaeoclimatology, Palaeoecology, 240:318-372.

Miller, A.K. and Furnish, W.M., 1940, Permian ammonoids of the Guadalupe Mountains region and adjacent areas: Geological Society of America Special Paper 26, 242pp.

Mitchell, C.E., Adhya, S., Bergström, S.M., Joy, M.P. and Delano, J.W., 2004. Discovery of the Ordovician Millbrig K-bentonite Bed in the Trenton Group of New York State: implications for regional correlation and sequence stratigraphy in eastern North America. Palaeogeography, Palaeoclimatology, Palaeoecology, 210:331-346.

Newell, N.D., Rigby, J.K., Fischer, A.G., Whiteman, A.J., Hickox, J.E. and Bradley, J.S., 1953. The Permian Reef Complex of Guadalupe Mountains Region, Texas and New Mexico. San Francisco: W.H. Freeman & Company, 236 pp.

Nicklen, B.L., 2003. Middle Guadalupian (Permian) Bentonite Beds, Manzanita Member, Cherry Canyon Formation, West Texas: Stratigraphic and Tectonomagmatic Applications: unpublished M.S. Thesis University of Cincinnati, 66 p.

Plummer, F.B. and Scott, G., 1937. Upper Paleozoic ammonites in Texas; The geology of Texas, Volume III. Texas University Bulletin, 2132:1-237.

Rush, J. and Kerans, C., 2010. Stratigraphic response across a structurally dynamic shelf: the latest Guadalupian composite sequence at Walnut Canyon, New Mexico, U.S.A.. Journal of Sedimentary Research, 80:808-828.

Samson, S.D., Kyle, P.R. and Alexander, E.C. Jr., 1988. Correlation of North American Ordovician bentonites by using apatite chemistry. Geology, 16:444-447.

Samson, S.D., Matthews, S., Mitchell, C.E. and Goldman, D., 1995. Tephrochronology of highly altered ash beds: the use of trace element and strontium isotope geochemistry apatite phenocrysts to correlate K-bentonites. Geochemica et Cosmochimica Acta, 59:2527-2536.

Sarg, J.F., Markello, J.R. and Weber, L.J., 1999. The second-order cycle, carbonate-platform growth, and reservoir, source, and trap prediction, in Harris, P.M. et al., eds., Advances in Carbonate Sequence Stratigraphy: Application to Reservoirs, Outcrops

46

and Models, 11-34. Tulsa: Society of Economic Paleontologists and Mineralogists Special Publication No. 63.

Schoene, B., Latkoczy, C., Schaltegger, U. and Günther, D., 2010. A new method integrating high-precision U-Pb geochronology with zircon trace element analysis (U-Pb TIMS- TEA). Geochimica et Cosmochimica Acta, 74:7144-7159.

Sell, B.K., 2010, “Apatite trace element tephrochronology” [unpublished Ph.D. dissertation]: Syracuse University, Syracuse, New York, 306 p.

Sell, B.K. and Samson, S.D., 2011. Apatite phenocryst compositions demonstrate a miscorrelation between the Millbrig and Kinnekulle K-bentonites of North America and Scandinavia. Geology, 39:303-306.

Sell, B.K. and Samson, S.D., in press. A tephrochronologic method based on apatite trace element chemistry. Quaternary Research.

Sha, L.-K. and Chappell, B.W., 1999. Apatite chemical composition, determined by electron microprobe and laser-ablation inductively coupled plasma spectrometry, as a probe into granite petrogenesis. Geochimica et Cosmochimica Acta 63:3861-3881.

Shen, S.-Z., Henderson, C.M., Bowring, S.A., Cao, C.-Q., Wang, Y., Wang, W., Zhang, H., Zhang, Y.-C. and Mu, L., 2010. High-resolution (Late Permian) timescale of South China. Geological Journal, 45:122-134.

Słowakiewicz, M., Kiersnowski, H. and Wagner, R., 2009. Correlation of the Middle and Upper Permian marine and terrestrial sedimentary sequences in Polish, German, and USA Western Interior Basins with reference to global time markers. Paleoworld, 18:193-211.

Steiner, M.B., 2006. The magnetic polarity time scale across the Permian–Triassic boundary. In: Lucas, S.G., Cassinis, G., Schneider, J.W. (Eds.), Non-marine Permian biostratigraphy and biochronology, 15–38. Geological Society of London Special Publication No. 265.

Stormer, J., Pierson, M. and Tacker, R., 1993. Variation of F and Cl X-ray intensity due to anisotropic diffusion in apatite during electron microprobe analysis. American Mineralogist, 78:641-648.

Terrell, J.H., 1960. Separation of zircon and biotite from bentonite for absolute dating and possibilities for dating certain west Texas bentonites: [unpublished Master’s thesis] Houston, Texas, Rice University.

Tinker, S.W., 1998. Shelf-to-basin facies distributions and sequence stratigraphy of a steep- rimmed carbonate margin: Capitan depositional system, McKittrick Canyon, New Mexico and Texas. Journal of Sedimentary Research, 68:1146-1174.

47

Todd, R.G., 1976. Oolite-bar progradation, San Andres Formation, Midland Basin, Texas. American Association of Petroleum Geologists Bulletin, 60:907-925.

Tyrrell, W.W., Diemer, J.A., Bell, G.L. and Griffing, D.H., 2004. Subsurface stratigraphy of the Manzanita Limestone Member, Cherry Canyon Formation, northern Delaware Basin, New Mexico and West Texas, in Trentham, R.C., ed. West Texas Geological Society Publication No. 04-112, p. 125-155.

Tyrrell, W.W., Diemer, J.A., Bell, G.L. and Griffing, D.H., 2006. The Manzanita Limestone Member, Cherry Canyon Formation, northern Delaware Basin, New Mexico and West Texas, in Hinterlong, G., ed. Permian Basin Section SEPM, Publication No. 2006-46.

Ver Straeten, C.A., 2004. K-bentonites, volcanic ash preservation, and implications for Early to Middle Devonian volcanism in the Acadain orogen, eastern North America. Geological Society of America Bulletin, 116:474-489.

Walker, D.A., Golanka, J., Reid, A. and Reid, S., 1995. The effects of paleolatitude and paleogeography on carbonate sedimentation in the late Paleozoic, in Huc, A., ed., Paleogeography, Paleoclimate, and Source Rocks, 133-155: Tulsa, American Association of Petroleum Geologists Studies in Geology 40.

Wardlaw, B.R., Davydov, V. and Gradstein, F.M., 2004. The Permian Period, in Gradstein, F.M., Ogg, J.G., and Smith, A.G., eds., A Geologic Time Scale 2004. 249-270. Cambridge: Cambridge University Press.

Wilde, G.L., 1975. Fusulinid-defined Permian Stages, in Permian Exploration, Boundaries, and Stratigraphy, 67-83. Tulsa: Tulsa Society of Economic Paleontologists and Mineralogists and West Texas Geological Society, Permian Basin Section, Publication 75-65.

Young, E., Myer, A., Munson, E. and Conklin, N., 1969. Mineralogy and Geochemistry of Fluorapatite from Cerro do Mercado, Durango, Mexico, D84-D93. United States Geological Survey Professional Paper, Report P 650-D.

Ziegler, A.M., Hulver, M.L. and Rowley, D.B., 1997. Permian world topography and climate, in Martini, I.P., ed., Late Glacial and Postglacial Environmental Changes: Quaternary, Carboniferous-Permian, and , 111-146. Oxford: Oxford University Press.

48

Chapter 2

A New Shelf-to-Basin Timeline for the Middle Permian (Guadalupian) Capitan Depositional System, West Texas and Southeastern New Mexico, USA

Notes: (1) A modified version of this chapter will be submitted for peer-reviewed publication with co-authors G.L. Bell Jr. (United States National Park Service) and W.D. Huff (University of Cincinnati). (2) Depth references for the PDB-04 research well are in feet to be consistent with previous studies and core storage records. All depths reported are core depths, with the corresponding log depth being 12 feet greater (Garber et al. 1989).

ABSTRACT

The correlation of the Rader Limestone Member of the Bell Canyon Formation to its shelf equivalent has varied over the years. Some models have equated it with the Seven

Rivers Formation, while other to the Yates Formation, or a position straddling both units.

More recent work has refined this shelf-to-basin correlation using existing biostratigraphic tielines and sequence stratigraphic models, and places the Rader in the Y-3 high frequency sequence of the Yates Formation. Despite the presence of bentonites in both the Rader

Limestone and the Yates Formation, and the potential to add an absolute age datum to the sequence stratigraphic hierarchy, geochemical tephrochronologic techniques have not yet been applied to the deposits.

To address these issues, apatite phenocryst trace element chemistry was determined by electron microprobe analysis for samples from three bentonites collected from outcrop localities in Guadalupe Mountains National Park and two from the PDB-04 research well. Results are interpreted to indicate a correlation between a bentonite 20 m above the base of the Rader in the Patterson Hills of Guadalupe Mountains National Park,

49

and one at 1956 ft core depth in the PDB-04 core. Careful examination of the gamma-ray log for the well indicates that the 1956 ft interval is in the Y-3 high frequency sequence and confirms this proposed shelf-to-basin correlation. Further refinement can be made to the third cycle set of this high frequency sequence.

The correlation interpreted here also provides biostratigraphic constraint for the sample from the Rader (GM-20). This sample is the focus of U-Pb radioisotopic dating

(Chapter 3), but key index fossils have not been reported for its sampling locality. Data from the core indicate that GM-20 is in the upper portion of the Polydiexodina fusulinid zone, which corresponds to the Jinogondolella postserrata conodont zone. This provides biostratigraphic constraint for the sample and places it in the Capitanian Stage of the

Guadalupian Series. It also demonstrates the utility of tephrochronologic techniques in correlating key sections that may lack certain data with those that contain it.

INTRODUCTION

Shelf-to-basin correlations of the Capitan Reef margin have long been a source of debate and controversy. Despite tremendous exposures of strata, the presence of massive reef and fore-reef deposits have made tracing individual units difficult, if not impossible.

As with any attempt at correlating units from different depositional environments, common correlation tools, such as index fossils have been used to establish timelines that link deposits on the shelf, with their equivalent basinal deposits (FIGURE 2.1). One point of debate has been the correlation of the basinal Rader Limestone Member of the Bell Canyon

Formation to the shelf succession, with some workers equating it to the Seven Rivers

Formation (e.g. Hull 1957; Silver and Todd 1969; Garber et al. 1989; Beaubouef et al. 1999), others to the Yates Formation (e.g. Jacka 1979; Pray 1988), or a combination of the two

50

(e.g. Newell et al. 1953). By combining well-established fusulinid timelines with a sequence stratigraphic framework, Tinker (1996) provided a more refined correlation and placed the lower Rader Limestone in the Yates Y-3 high frequency sequence (HFS) and the upper Rader Limestone in his Yates Y-4 high frequency sequence (FIGURE 2.2). This correlation has been followed by recent workers (e.g Rush and Kerans 2010), though the difficulty in determining the exact part of the sequence stratigraphic framework the Rader represents, due to lack of recognizable surfaces, has been noted (Fall and Olszewski 2010).

Despite knowledge of the presence of highly altered volcanic tephra (bentonites) in the

Yates Formation and the Rader Limestone, teprochronologic techniques have not been used to test or further refine these alternative shelf-to-basin correlations.

51

The objective of this study is to test the correlation between the Yates Formation and the Rader Limestone by comparing the apatite trace element chemistry of bentonites from both units. Samples have been collected within and below the Rader Limestone in

Guadalupe Mountains National Park (GMNP) and within the Yates from the PDB-04 research well and in outcrop at McKittrick Ridge, in GMNP (FIGURE 2.3). The sample from within the Rader has been the focus of U-Pb ID-TIMS zircon dating (Chapter 3). If this sample can be correlated to the shelf, it would provide a temporal constraint for this portion of the sequence stratigraphic framework by creating the first “absolute” shelf-to- basin timeline for the Capitan depositional system. One problem, however, is that biostratigraphic control is currently not available at its sampling locality, although it can be inferred from the lithostratigraphic position of the sample. This can also be addressed using tephrochronologic techniques, as the PDB-04 core does have relevant data on index fossils. A positive correlation between the core and dated sample would provide

52

53

biostratigraphic constraint and allow radioisotopic data from that sample to inform estimates of stage boundaries and the timing of key global events.

Background and Previous Work

The Yates Formation is a carbonate and siliciclastic shelf succession that was deposited landward of the Capitan Reef (FIGURE 2.1). Based on exposures in McKittrick

Canyon (Tinker 1998) and Slaughter Canyon (Osleger 1998), Osleger and Tinker (1999) recognize five high frequency sequences (HFS) in the Yates Formation that they correlate through the Capitan Formation and into the carbonate tongues of the Bell Canyon

Formation (FIGURE 2.2). One of the shelf-to-basin correlations Osleger and Tinker (1999) have proposed is the Y-3 HFS to the lower Rader Limestone (see also Tinker 1996). Four facies tracts (shelf crest, outer shelf, reef margin, and continental siliciclastics) were used by Osleger and Tinker (1999) to identify periods of aggradation, retrogradation, and progradation through the succession. This model follows the depositional profile of

Dunham (1972), where a bathymetric high is present at the shelf crest, followed by an uninterrupted basinward slope to the shelf margin.

The Rader Limestone is one of several carbonate tongues that extend from the

Capitan fore-reef into the Delaware Basin (FIGURE 2.1). Relatively little work has been conducted on the Rader Limestone, though its megaconglomerate facies exposed along US

Highway 62/180 (“Rader Slide”) is frequently visited on field trips to the region. One of the few detailed works was by Lawson (1989), who divided the Rader into lower, middle, and upper units. Like King (1948), Lawson (1989) recognized an apple-green bentonite within the Rader at several localities and interpreted it to be a single, correlatable deposit at the base of her middle unit. In the current study, a bentonite of the same color and description

54

was sampled 20 m above the base of the Rader at Back Ridge (FIGURE 2.3) and may represent the same bed. Unfortunately, the measured sections with bentonites described by Lawson (1989) are on private land (D-Ranch) that is not currently accessible to researchers. Based on the thickness of the Rader at its type locality (65m) (Lawson 1989), the sample from Back Ridge (GM-20), is probably from the lower portion of the member.

Previous Shelf-to-Basin Bentonite Correlations

In 1984 the Gulf PDB-04 research well was drilled at the northern end of the

Delaware Basin in Eddy County, New Mexico, approximately 90 kilometers northeast of

GMNP (FIGURE 2.3). In their study of the core from this well, Garber and others (1989) describe two claystone beds, which they interpreted as bentonites, in the lower Yates

Formation. These beds are separated by 8 meters of dolostone and are significant because they represent the only known bentonite samples from the shelfal strata of the Capitan

Reef depositional system. At present, the only work to be done on these samples was that of Garber and others (1989) who examined the physical properties of the beds and made interpretations of shelf-to-basin correlations based on their stratigraphic positions. Two potential correlations for the Yates bentonites suggested by the authors are the BC bentonite of Wilde (1975), and the bentonite noted in the Rader Limestone by King (1948).

Additional reports of bentonite beds in the Yates Formation come from Todd (1976), who recognized them in the subsurface on the west side of the Central Basin Platform, though he did not report specific wells.

The BC bentonite of Wilde (1975) is placed in the upper part of the Bell Canyon

Formation. This is up section from the Rader Limestone (FIGURE 2.3), where King (1948) recognized a bentonite bed at several localities. King (1948) also noted the occurrence of

55

two bentonite beds in the upper part of the Bell Canyon Formation in Niehaus et al.

Caldwell No. 1 well, 56 kilometers southeast of the Guadalupe Mountains. Though King

(1948) made no attempt to correlate these bentonites, Garber and others (1989) suggested they were equivalent to the BC bentonite of Wilde (1975) based on their similar stratigraphic position. The position of the BC bentonite was placed above the McCombs

Member by Garber and others (1989, fig 5) but as noted above, they also suggested the possibility that one of the Yates bentonites correlates with the bentonite noted by King

(1948) in the Rader.

METHODS

Sample Collection and Prepartation

Bulk bentonite samples were collected from three outcrop localities and the PDB-04 core (FIGURE 2.3) using a pick and small trowel (outcrop) or a stainless steel spoon (core) that was cleaned between uses. Samples were placed in one-gallon plastic bags and labeled. Care was taken to insure that only material that was suspected to be bentonite was included. TABLE 2.1 summarizes the locality information and the sample names. The bulk samples were allowed to soak in beakers for at least 48 hours, prior to partial disaggregation by stirring. In some cases, the degree of silicification required the use of a mortar and pestle to disaggregate, however, care was taken to crush, rather than grind the samples. The disaggregated samples were then flushed of clay-sized particles and sieved with a 178 µm mesh screen following a slightly modified version of the methods described by Sell (2010). Heavy mineral (> 2.85 g/cm3) separation was performed using lithium heteropolytungstate. Crystals for analysis were hand picked using a binocular microscope

56

to ensure that the most euhedral and clear apatites were chosen. The apatites were

mounted in one-inch epoxy rounds and exposed and polished using 4000 grit sand paper

and 0.05 µm aluminum oxide. Crystals were mounted with their c-axis parallel to the

polishing surface in an attempt to avoid issues with Cl analysis noted by Stormer and

others (1993). Grain mounts were coated with carbon prior EMPA (electron microprobe

analysis).

Table 2.1- Sample Register

Electron Microprobe Analysis

Chemical analysis of apatite phenocrysts was conducted on a Cameca SX-50 electron

microprobe in the Department of Earth and Environmental Sciences at the University of

Kentucky. Each spot was analyzed for Cl, Fe, Mg, Mn, Ce, & Y using a 15 kV accelerating

voltage and a beam current of 60 nA (faraday cup current), with a beam diameter of 1-2

µm. Counting times were 30 seconds on peak for Cl; 40 seconds for Fe, Mg, and Mn; 60

seconds for Ce and Y. For background, counting times were 15 seconds for Cl; 40 seconds

for Fe, Mg, and Mn; 60 seconds for Ce and Y. Testing on apatite from the Guadalupian

Manzanita Limestone demonstrated variable concentrations for Cl based on the duration

and order it was analyzed in each run. Also observed during testing were changes in Mn

concentration based on whether it came first or second during the analytical run for each

spot. To avoid these issues, Cl and Mn were measured first at each spot (on separate

57

spectrometers), with Cl having shorter counting times. The Smithsonian apatite standard

(NMNH 104021) was analyzed following each calibration to assess precision and as a check on relative accuracy. Previously analyzed apatites from the Ordovician Deicke K-bentonite collected at Union Furnace, Pennsylvania (B. Sell pers. comm. 2010), were used as an additional secondary standard and analyzed after each calibration.

RESULTS

The trace element assemblages of apatite phenocrysts from five bentonite samples are reported in APPENDIX B, along with averages from repeated analyses of the

Smithsonian apatite standard. Most data points listed represent the average of three separate spots analyzed per grain, with two spots being averaged for some very small grains. Intra-crystal chemical variation was minimal, such that the calculated averages represent the range of values encountered for each sample. Each spot was analyzed for minor, trace and rare earth elements (Cl, Mg, Fe, Mn, Ce, and Y). The minor and trace elements chosen for analysis have proven useful in discriminating between Paleozoic apatite phenocrysts from discrete bentonites by previous workers (Samson et al. 1995;

Emerson et al. 2004; Mitchell et al. 2004; Adhya 2009; Carey, Samson and Sell 2009; Sell and Samson 2011) and, with the exception of Fe, were useful in discriminating between closely spaced bentonite samples in the Wordian Manzanita Limestone (Chapter 1). The other two non-major elements where analyzed because they represent a light REE (Ce) and a heavy REE (Y) in the absence of an entire REE analytical suite. Ce and Y concentrations of apatites from granites have shown covariation (Sha and Chappell 1999). When used as a ratio, Ce/Y was also useful in discriminating between and characterizing Manzanita

58

bentonites. Recently, Sell and Samson (2011) have demonstrated the usefulness of this ratio in characterizing Ordovician K-bentonites.

FIGURE 2.4 presents apatite phenocryst chemistry from the five outcrop and core samples (TABLE 2.1). The diagrams demonstrate that the samples from the same lithostratigraphic interval (i.e. GM-20 [Rader] and GM-21[immediately sub-Rader] and the

PDB-04 samples [Yates]) have a unique grouping of data points and are clearly distinguishable. As with the Manzanita samples, Mg, Mn, Cl, and Ce/Y are the best discriminators, based on all possible bivariate diagrams. Fe appears to also be a good discriminator, though the analyses are not as precise as the other elements (APPENDIX B).

59

DISCUSSION

The plots of the EMPA data in FIGURE 2.4 show that some samples exhibit the same tight clustering or complex data trends. These samples are interpreted to indicate the unique geochemistry of coeval tephras. Samples GM-20 and PDB-04 1956 both have relatively complex patterns of data that suggest the presence of at least two apatite populations. In each diagram, they consistently share a main cluster of data points and a second grouping of more widely dispersed data. The cause of this is unclear, though some potential explanations are discussed in Chapter 1. That these two samples share not only one tight cluster of data, but also a complex pattern, strongly suggests that they are from a coeval tephra deposit. These samples are interpreted to represent deposits from the same tephra and are therefore correlative. This interpretation establishes a timeline between the basinal Rader Limestone and the Yates Formation on the shelf.

Shelf-to-Basin Correlation

The apatite data interpreted here support the correlation between the Yates

Formation and the Rader Limestone. Furthermore, they confirm the conclusion of Tinker

(1998), that the lower portion of the Rader Limestone is equivalent to the Y-3 cycle of the

Yates Formation. FIGURE 2.5 shows the gamma-ray log for the PDB-04 core and the positions of the Yates HFSs from Osleger and Tinker (1999, fig 12a). The two bentonites at

1932 and 1956 ft are identified as sharp spikes in the log. From this, it is clear the Yates bentonites sampled here occur in the Y-3 cycle. When coupled with the fusulinid biostratigraphy, the Capitanian shelf-to-basin correlations are now supported by multiple high-resolution tielines.

60

61

In their study of the three-dimensional architecture of the Yates HFS in McKittrick and Slaughter canyons, Osleger and Tinker (1999, fig 12) compare profiles from those two localities with a composite cross section from the Northwestern Shelf and Central Basin

Platform by Borer and Harris (1991). In this comparison, they interpret the placement of the Yates HFS boundaries on the gamma-ray log of the PDB-04 research well. Additionally, they note that, for the Y-3 HFS, four cycle sets (shallowing upward successions of meter- scale cycles), are present at both canyons and suggest that they may be directly correlative.

The base of three of the four Y-3 cycle sets in Slaughter Canyon are marked by siliciclastics that were reworked in shallow marine environments during transgression (Borer and

Harris 1991). FIGURE 2.5 shows the interpreted cycle sets for the Y-3 cycle in the PDB-04 gamma-ray log by the current study and the positions of the Yates bentonites. Based on these interpretations, the position of the two bentonites in the Yates Formation and, by direct correlation, the bentonite (GM-20) in the Rader Limestone, can be further refined to the third cycle set in the Y-3 HFS.

The interpretation of these data suggest that a bentonite correlating with PDB-04

1932 should be present in the basin above the level of GM-20. It is possible that accumulation of ash was wiped away by the debris flow deposits of the upper portions of the Rader. Alternatively, a bentonite may be present further up in the Rader or even higher, but has eluded discovery. Experience has shown that, while obvious at some localities, bentonites can often be highly inconspicuous, and easy to miss.

Despite the implication that there should be a shelf equivalent to GM-21, no bentonites were reported below the PDB-04 1956 sample in the core until the Manzanita

Limestone at 4008 ft (Garber et al. 1989). This is not surprising as GM-21 was collected

62

from fine-grained siliciclastics. Many models for the Captanian depositional system follow the reciprocal sedimentation concept of Wilson (1967)(e.g. Borer and Harris; Tinker 1998) and equate siliciclastics in the basin with sea-level lowstands and bypass on the shelf. A period of lowstand and siliciclastic bypass on the shelf would create poor conditions for the preservation of accumulated volcanic ash. While no shelf equivalent may exist for GM-21, it probably does correlate to an immediately sub-Rader bentonite identified by King (1948, plate 6) near the junction of Lamar and Bell canyons on the D Ranch, though this has not yet been demonstrated geochemically.

Sample 442-1-1 was collected from the Yates Formation at McKittrick Ridge

(FIGURE 2.3), in a position downdip (i.e. reefward) of the shelf crest facies tract of Osleger and Tinker (1998), but its sequence stratigraphic position was not identified. The correlation interpreted here places this sample in the third cycle set of the Y-3 HFS. It also suggests that another bentonite should be present in the lower portion of the same cycle set. If either or both of these bentonites can be identified in the future, they should be useful markers for this part of the shelf sequence as there is considerable amount of facies and thickness variability up- or downdip with respect to the shelf margin. As with this example, they could also provide direct correlation to other outcrop localities or subsurface logs at the cycle set scale.

Biostratigraphic Control for Sample GM-20

In addition to establishing a new shelf-to-basin timeline, the correlation interpreted here also provides biostratigraphic control for the radioisotopic data for sample GM-20.

The sample locality has been the focus of a recent work on brachipods in the debris flows of the Pinery and Rader limestones (Fall and Olszewski 2010), but key index fossils like

63

fusulinids, and especially , have not been reported. Fortunately, fusulinids were systematically collected over a 2086 foot interval (636 m) through the PDB-04 core, beginning at 1918 ft (Garber et al. 1989). The uppermost sample is 38 ft (11.6 m) above the bentonite sample at 1956 and was identified as Polydiexodina, with its lowest sampled occurrence in the core at 3827 ft (FIGURE 2.6). Based on the correlation discussed above, this places GM-20 within the Polydiexodina fusulinid zone. This is useful information, but, as this zone overlaps the Wordian/Capitanian boundary, it does not provide chronostratigraphic refinement useful in addressing the lack of temporal constraint for the

Guadalupian. To do this, the position of the bentonite sample within the zone must be determined.

Although the uppermost fusulinid sampled is Polydiexodina, this does not necessarily represent its last-occurrence datum (LAD). The LAD of Polydiexodina has long been established to be in the Yates Formation, with Newell and others (1953) having placed it near the top of their “B” member (Y-5 HFS of Tinker 1996, 1998 and Osleger and

Tinker 1999). This placement has been confirmed by more recent workers (Kerans and

Harris 1993; Brown 1996). The top of the Yates “B” member has traditionally been equated to the top of the “Hairpin” dolomite (Brown 1996), which is one of four informal member names for the upper Yates Formation proposed by Pray and Esteban (1977).

These members are readily identifiable in many gamma-ray logs and are useful in correlating the upper portion of the Yates Formation in the subsurface. Examination of the

PDB-04 log shows the top of the “Hairpin” dolomite to occur at approximately 1720 ft

(FIGURE 2.6). As the top of this informal member corresponds to the Polydiexodina LAD, it places GM-20, through the geochemical correlation, 236 ft below. Given that the lowest

64

65

occurrence of Polydiexodina in the core is at 3827 ft, or 1871 ft below PDB-04 1956, sample

GM-20 can be considered to occur in the upper part of that fusulinid zone. The relatively thick limestone of the Capitan reef and fore-reef slope accounts for a large amount of the core between 1956 and 3827. This is important to note as a Capitan reef/fore-reef section of a given thickness probably represents less time than a section of similar thickness on the shelf or in the basin. Even if the Capitan thickness is removed, the distance from PDB-04

1956 to the Polydiexodina LAD is less than half that to the first occurrence.

With GM-20 firmly placed in the upper portion of the Polydiexodina zone, its position within the appropriate conodont zone can now be determined by comparing the ranges of each index fossil. FIGURE 2.6 shows the ranges of Polydiexodina and the conodont Jinogondolella postserrata from Lambert and others (2010). The upper half of the

Polydiexodina zone corresponds to the J. postserrata zone. This places sample GM-20 in the

J. postserrata zone and the Capitanian, as the first-appearance datum of this conodont defines the base of that stage.

CONCLUSIONS

This study adds a new high-precision shelf-to-basin tieline to complement the LAD of Polydiexodina (Newell et al. 1953; Brown 1996) and the FAD of Yabeina (Tyrrell 1969).

Based on apatite phenocryst geochemistry, a bentonite from the lower portion of the basinal Rader Limestone correlates with the Y-3 HFS in the Yates Formation on the shelf.

This confirms the correlation of Tinker (1996, 1998) and adds further refinement by placing the correlated bentonites in the third cycle set of the HFS. A key marker bed has been identified within the Rader Limestone that identifies its exact position within the shelf defined sequence stratigraphic framework of Osleger and Tinker (1999). Based on this

66

correlation, the biostratigraphic position of sample GM-20 is in the upper portion of the

Polydiexodina fusulinid zone, which overlaps with the J. postserrata conodont zone. This sample is the focus of radioisotopic dating and the data presented in Chapter 3 are now biostratigraphically constrained to the Capitanian.

ACKNOWLEDGEMENTS

I would like to thank the National Park Service for granting permission to collect samples from within Guadalupe Mountains National Park under permit # GUMO-2007-SCI-

0010. B.K. Sell is thanked for providing the previously analyzed Deicke apatites and for many stimulating discussions on apatite phenocryst chemistry. D. Moecher and S.

Chakraborty provided access to the Cameca SX-50 at the University of Kentucky. T. Phillips helped with drafting some of the figures. N. Bose, C. Ferguson, and J. Wilkins assisted with sample collection. L. Fall and S. Marcus located the bentonite in the Rader at Back Ridge and alerted me to its presence. P.M. Harris graciously shared information regarding the

PDB-04 core. P. O’Neill (Louisiana Geological Survey) and J. Donnelly (Austin Core

Research Center) are thanked for providing access to the PDB-04 core. Graduate student research funding was provided by the American Association of Petroleum Geologists (Ohio

Geological Society Named Grant), the Clay Minerals Society, the Geological Society of

America, and the Department of Geology at the University of Cincinnati.

67

REFERENCES

Adhya, S., 2009, “Geochemical fingerprinting of volcanic airfall deposits, a tool in stratigraphic correlation” [unpublished Ph.D. dissertation]: The State University of New York, Albany, New York, 532 p.

Beaubouef, R.T., Rossen, C., Zelt, F.B., Sullivan, M.D., Mohrig, D.C. and Jennette, D.C., 1999. Deep-water sandstones, Brushy Canyon Formation, west Texas: American Association of Petroleum Geologists Continuing Education Course Notes Series #40, 62p.

Borer, J.M. and Harris, P.M., 1991. Lithofacies and cyclicity of the Yates Formation, Permian Basin: Implications for reservoir heterogeneity. American Association of Petroleum Geologists Bulletin, 75:726-779.

Brown, A., 1996. Position of the Polydiexodina last-occurrence datum in Guadalupian Strata: revised correlations at McKittrick Canyon. In Demis, W.D., and Cole, A.G. (eds.), The Brushy Canyon Play in Outcrop and Subsurface: Concepts and Examples: Permian Basin Section of Society of Economic Paleotologists and Mineralogists Publication 96-38, p. 75-83.

Carey, A., Samson, S.D., and Sell, B, 2009. Utility and limitations of apatite phenocryst chemistry for continent-scale correlation of Ordovician K-bentonites. Journal of Geology, 117:1-14.

Dunham, R.J., 1972. Capitan reef, New Mexico and Texas—Facts and Questions to aid interpretation and group discussion: Midland, Society of Economic Paleontologists and Mineralogists, Permian Basin Section, Publication 72-14, 297 p.

Emerson, N.R., Simo, J.A., Byers, C.W. and Fournelle, J., 2004. Correlation of (Ordovician, Mohawkian) K-bentonites in the upper Mississippi valley using apatite chemistry: implications for stratigraphic interpretation of the mixed carbonate-siliciclastic Decorah Formation. Palaeogeography, Palaeoclimatology, Palaeoecology, 210:215- 233.

Garber, R.A., Grover, G.A. and Harris, P.M., 1989. Geology of the Capitan Shelf Margin- subsurface data from the northern Delaware Basin, in Harris, P.M., and Grover, G.A., eds., Subsurface and outcrop examination of the Capitan Shelf Margin-a Core Workshop: Tulsa, Society of Economic Paleontologists and Mineralogists Core Workshop 13, p. 3-269.

Hull, J.P.D. Jr., 1957. Petrogenesis of Permian Delaware Mountain Sandstone, Texas and New Mexico. American Association of Petroleum Geologists Bulletin, 41:278-307.

68

Jacka, A.D., 1979. Deposition and entrapment of hydrocarbons in Bell Canyon and Cherry Canyon deep-sea fans of the Delaware Basin. In Sullivan, N.M. ed., Guadalupian Delaware Mountain Group of West Texas and Southeast New Mexico: Society of Economic Paleontologists and Mineralogists, Permian Basin Section, 79-18:104-120.

Kerans, C. and Tinker, S.W., 1999. Extrinsic controls on development of the Capitan Reef complex, in Saller et al., eds., Geologic Framework of the Capitan Reef, Tulsa, OK, SEPM Special Publication No. 65, p. 15-36.

King, P.B., 1948. Geology of the southern Guadalupe Mountains, Texas, U.S. Geological Survey Professional Paper 215, 183p.

Mitchell, C.E., Adhya, S., Bergström, S.M., Joy, M.P. and Delano, J.W., 2004. Discovery of the Ordovician Millbrig K-bentonite Bed in the Trenton Group of New York State: implications for regional correlation and sequence stratigraphy in eastern North America. Palaeogeography, Palaeoclimatology, Palaeoecology, 210:331-346.

Newell, N.D., Rigby, J.K., Fischer, A.G., Whiteman, A.J., Hickox, J.E. and Bradley, J.S., 1953. The Permian Reef Complex of Guadalupe Mountains Region, Texas and New Mexico: W.H. Freeman & Company, San Francisco, 236 p.

Osleger, D.A., 1998. Sequence architecture and sea-level dynamics of Upper Permian shelfal facies, Guadalupe Mountains, southern New Mexico. Journal of Sedimentary Research, 68:327-346.

Osleger, D.A. and Tinker, S.W., 1999. Three-dimensional architecture of Upper Permian high-frequency sequences, Yates-Capitan shelf margin, Permian Basin, U.S.A., in Harris, P.M. et al., eds., Advances in Carbonate Sequence Stratigraphy: Application to Reservoirs, Outcrops and Models, Tulsa, OK, SEPM Special Publication No. 63, p. 169- 185.

Samson, S.D., Matthews, S., Mitchell, C.E. and Goldman, D., 1995. Tephrochronology of highly altered ash beds: the use of trace element and strontium isotope geochemistry apatite phenocrysts to correlate K-bentonites. Geochemica et Cosmochimica Acta, 59:2527-2536.

Sell, B.K., 2010, “Apatite trace element tephrochronology” [unpublished Ph.D. dissertation]: Syracuse University, Syracuse, New York, 306 p.

Sell, B.K. and Samson, S.D., 2011. Apatite phenocryst compositions demonstrate a miscorrelation between the Millbrig and Kinnekulle K-bentonites of North America and Scandinavia. Geology, 39:303-306.

Silver, B.A. and Todd, R.G., 1969. Permian cyclic strata, northern Midland and Delaware basins, west Texas and southeastern New Mexico: American Association of Petroleum Geologists Bulletin, 53:2223-2251.

69

Stormer, J., Pierson, M. and Tacker, R., 1993. Variation of F and Cl X-ray intensity due to anisotropic diffusion in apatite during electron microprobe analysis. American Mineralogist, 78: 641-648.

Tinker, S.W., 1996. “Reservoir-scale sequence stratigraphy: McKittrick Canyon and 3-D subsurface examples, west Texas and New Mexico” [unpublished Ph.D. thesis]: The University of Colorado, Boulder, Colorado, 245 p.

Tinker, S.W., 1998. Shelf-to-basin facies distributions and sequence stratigraphy of a steep- rimmed carbonate margin: Capitan depositional system, McKittrick Canyon, New Mexico and Texas. Journal of Sedimentary Research, 68:1146-1174.

Todd, R.G., 1976. Oolite-bar progradation, San Andres Formation, Midland Basin, Texas. American Association of Petroleum Geologists Bulletin, 60:907-925.

Walker, D.A., Golanka, J., Reid, A. and Reid, S., 1995. The effects of paleolatitude and paleogeography on carbonate sedimentation in the late Paleozoic, in Huc, A., ed., Paleogeography, Paleoclimate, and Source Rocks: Tulsa, American Association of Petroleum Geologists Studies in Geology 40, p. 133-155.

Wilde, G.L., 1975. Fusulinid-defined Permian Stages, in Permian Exploration, Boundaries, and Stratigraphy: Tulsa Society of Economic Paleontologists and Mineralogists and West Texas Geological Society, Permian Basin Section, Publication 75-65, p. 67-83.

Wilde, G.L., Rudine, S.F. and Lambert, L.L., 1999. Formal designation: Reef Trail Member, Bell Canyon Formation, and its significance for recognition of the Guadalupian- Lopingian boundary in Saller, A.H., Harris, P.M., Kirkland, B.L. and Mazzullo, S.J., eds., Geologic framework of the Capitan Reef, Tulsa, OK, SEPM Special Publication No. 65, p. 63-83.

Wilson, J.L., 1967. Cyclic and reciprocal sedimentation in Virgilian strata of southern New Mexico. Geological Society of America, Bulletin, 78:805-818.

Ziegler, A.M., Hulver, M.L. and Rowley, D.B., 1997. Permian world topography and climate, in Martini, I.P., ed., Late Glacial and Postglacial Environmental Changes: Quaternary, Carboniferous-Permian, and Proterozoic: Oxford University Press, Oxford, p. 111- 146.

70

CHAPTER 3

New ID-TIMS Zircon Ages from the Middle Permian (Guadalupian) Type Area: Implications for the Geologic Time Scale and the Timing of Global Events

Note: A modified version of this chapter will be submitted for peer-reviewed publication with co-authors B.K. Sell (University of Geneva), G.L. Bell Jr. (United States National Park Service), and W.D. Huff (University of Cincinnati)

ABSTRACT

The temporal positions for the bases of the three component stages of the Middle

Permian (Guadalupian) are poorly constrained by radioisotopic dates. The Guadalupian is an important time interval for significant changes in Earth’s climate and biodiversity that include events leading up to the proposed double-phased mass extinction at the end of the

Paleozoic. In light of this, the duration of the Guadalupian stages and boundary age estimations need to be updated in order to assess cause and effect of these significant events. Also needed is better temporal constraint for key Guadalupian global chemostratigraphic and geomagnetic markers. This problem was approached by locating bentonites in the Guadalupian type area with zircon crystals, which were dated via CA-ID-

TIMS. Results here provide age estimates for the Illawarra geomagnetic reversal (c. 266.5

Ma) and the Kamura cooling event (at most c. 262.5 Ma), both of which have been implicated in the end-Guadalupian mass extinction. The Illawarra Reversal also provides constraint for the extinction of the dinocephalian , which can now be given a radioisotopically-based age estimate. These new data also suggest that the current estimates of the Wordian-Capitanian boundary, and probably the Roadian-Wordian

71

boundary, are too old. This extends the duration of the Wordian to c. 3 myr and decreases the duration of the Capitanian to c. 4 myr.

INTRODUCTION

Despite being the focus of significant research attention, the Guadalupian type area is one of the least constrained portions of the Paleozoic time scale. The Global Boundary

Stratotype Sections and Points (GSSPs) for the three component stages of the Guadalupian

Series are located in Guadalupe Mountains National Park (GMNP) in west Texas (FIGURE

3.1). To date, only one radioisotopic date (Bowring et al. 1998) has been determined for the Guadalupian, and reports of its stratigraphic position (e.g. Bowring et al. 1998;

Glenister et al. 1999; Wardlaw et al. 2004) have been inconsistent. Current estimates for the boundaries of the Guadalupian stages are based on this single date. Temporal control on key correlation events, such as the Illawarra geomagnetic reversal (IR) and excursions in the global δ13C curve are poor and may be inaccurate if the existing date is referenced from an incorrect stratigraphic position. Improved time constraint is desirable, as there is currently considerable debate regarding events during and after the Guadalupian, including both terrestrial and marine mass extinctions and their associated causes. To address the lack of temporal control for this interval, new isotope dilution thermal ionization mass spectrometry (ID-TIMS) U-Pb ages were calculated for zircons from layers of highly altered tephra (bentonites) collected in the Guadalupian type area at GMNP

(FIGURE 3.1).

72

Background

The GSSPs for the three stages of the Guadalupian Series are located in GMNP, which is situated in the western Delaware Basin and southern Guadalupe Mountains. During the

Guadalupian, the Delaware Basin was encircled by a carbonate margin that was transitional between a distally steepened ramp and a reef-rimmed shelf platform (Kerans and Tinker

1999). Located roughly 5o north of the paleoequator, the study area is thought to have been in an arid tropical (Walker et al. 1995) or subtropical climate zone (Ziegler, Hulber, and Rowley 1997).

The stratotype sections are in the basinal units of the Cutoff Formation and the

Delaware Mountain Group and the base of each stage is defined by species transitions within the conodont genus Jinogondolella (Glenister et al. 1999; Wardlaw, Davydov and

73

Gradstein 2004). Current age estimates for the boundaries of the Guadalupian stages are poorly constrained, with a date of 265.3 ± 0.2 Ma by Bowring and others (1998) being the basis for most estimates (e.g. Wardlaw et al. 2004; Menning et al. 2006; Ogg et al. 2008).

As alluded to above, the exact stratigraphic position of this date has been inconsistently reported (see review in Chapter 1), though it is commonly treated as being at or very near the base of the Capitanian (e.g. Wardlaw, Davydov and Gradstein 2004; Korte et al. 2005;

Menning et al. 2006; Słowakiewicz, Kiersnowski and Wagner 2009).

Global Events and Correlation Tools

One aspect of the Guadalupian that makes time constraint so critical is the proposed double-phased mass extinction at the end of the Paleozoic (i.e. end-Guadalupian and end-

Permian) (Jin et al. 1994; Stanley and Yang 1994). Estimates for the end-Guadalupian extinction suggest that 58% of marine genera did not survive into the Late Permian

(Stanley and Yang 1994; Knoll et al. 1996). Terrestrial records have also been interpreted as recording a multi-phased end-Permian extinction, with an end-Guadalupian pulse marked primarily by the extinction of 18 genera of Dinocephalian therapsids in South

Africa (Retallack et al. 2006). This has been challenged by Lucas (2009), who recognizes the same elimination of genera, but places the “dinocephalia ” (DEE) in the early Capitanian or Wordian, depending on the placement of the Illawarra geomagnetic reversal.

Recent data from marine invertebrates in South China and have been interpreted to indicate that the first pulse of the double-phase mass extinction was not concentrated at the end of the Guadalupian, but earlier in the Capitanian (e.g. Isozaki et al.

2007b; Wignall et al. 2009; Bond et al. 2010). The record from South China appears to be

74

contemporaneous with flood basalt volcanism of the Emeishan Large Igneous Province

(LIP), which may have been a causal mechanism for the extinction (Wignall 2001; Ali et al.

2002; Zhou et al. 2002; Wignall et al. 2009). Based on C isotope chemostratigraphic correlation between South China and Japan, the extinction predates volcanism in the latter

(Bond et al. 2010) and might be related to a cooling event (Isozaki et al. 2007b). It might also be possible that rather than an abrupt end-Guadalupian extinction pulse, there was a gradual decrease in diversity from the Middle Guadalupian to the end of the Permian, controlled partially by reduced origination rates in the Capitanian and

(Clapham et al. 2009). If this were the case, a causal mechanism for an abrupt extinction would no longer be required, but rather something that could explain significant reductions in origination.

In a mid-Panthalassan paleo-atoll carbonate section in Japan, Isozaki and others

13 (2007a) identified a unique elevated plateau of δ C values and named it the Kamura event.

The paleo-atoll carbonates lack conodonts, but based on fusulinids, it was proposed that the Kamura event occurred during the early-middle portions of the Capitanian. It was interpreted to indicate a period of high bio-productivity that drew down atmospheric CO2 and caused cooling. The event is the first significant δ13C positive excursion since the early

Mississippian and marks the beginning of relatively high amplitude fluctuations across the

Permian-Triassic boundary that indicates a change to the post-Paleozoic global climate and oceanography (Isozaki et al. 2007b). The “end-Guadalupian” extinction is interpreted by

Isozaki and others (2007b) to occur in the middle of the Kamura event, and they cite it as a potential cause. As mentioned above, it is possible that there was no major extinction pulse in the Guadalupian, but cooling associated with the Kamura event could also be potential

75

cause of a reduction in origination rates in the Capitanian and Wuchiapingian (Clapham et al. 2009).

In South China, Bond and others (2010, fig. 11) recently identified a similar plateau in δ13C values and interpreted it to be the Kamura event. Unlike the paleo-atoll in Japan, conodonts have been recovered from the section in South China and the beginning of the

Kamura event is placed in the upper portion of the J. postserrata zone. The end of the event is marked by a negative excursion in δ13C in the J. altudaensis zone. Bond and others (2010) argued that the “end-Guadalupian” extinction event occurs in this negative excursion and that it is probably related to the Emeishan volcanics. They also recognized that cooling recorded by the Kamura event could be a cause. The extinction in Japan occurs within the elevated δ13C plateau (Isozaki et al. 2007b) and appears to predate the extinction in South

China and the Emeishan volcanics, though backsmearing (i.e. Signor-Lipps effect) may affect this interpretation (Bond et al. 2010).

A similar trend of elevated δ13C values has also been recognized in West Texas though it appears to begin in the Wordian rather than Capitanian (Korte et al. 2005, fig. 6).

It should be noted that these data are somewhat sporadic, with two elevated data points within the Wordian and two from very near the base of the Capitanian defining the upward trend. In addition to the work by Korte and others (2005), Noble and others (2011) have recently reported δ13C values for the Capitanian Lamar Limestone Member from a road cut near GMNP. Their results indicate elevated values in the Lamar, with a negative shift near the top of the studied section, though the authors note that the data are scattered. They interpret the shift to indicate the influx of meteoric water, but acknowledge that secular change is a possibility. If the latter is the case, this shift to lower δ13C values might indicate

76

the end of the Kamura event. It is also possible that the trend of elevated δ13C continues, as initial reports suggest the plateau of high values extends into the uppermost Capitanian

Reef Trail Member, before a large negative shift near the end of the stage (i.e. end

Guadalupian) (Noble et al. 2009). These new data place the West Texas curve in closer alignment to those from Japan (Isozaki et al. 2007b) and Croatia (Isozaki et al. in press), supporting the notion that the Kamura event is global and can be used as a correlation tool.

The Ilawarra reversal marks the end of the 40 myr Permo-Carboniferous Reversal

Superchron (Irving and Parry, 1963) and is considered to be the most pronounced magnetostratigraphic marker in the Paleozoic (Menning et al. 2006). The IR has frequently been employed as a correlation tool for Permian sections. Despite the significance of this feature, its age has yet to be constrained. One issue is that the stratigraphic position of the

IR was moved as new magnetostratigraphic studies were reported. Based on the work of

Peterson and Nairn (1971) and Menning (2000), the IR was thought to occur very near the base of the Capitanian and was placed in this position by Wardlaw and others (2004).

Based on comparing North American and Transcaucasion magnetostratigraphic profiles,

Steiner (2006) suggested that the IR is middle-upper Wordian. This interpretation was adopted by Ogg and others (2008, fig. 9.4).

Sample GM-29 from this study comes from very near the base of the Middle

Wordian South Wells Limestone. According to Steiner (2006), the oldest normal polarity in the Guadalupian type area occurs in the shelfal Queen Formation, though it may possibly be in the underlying . One issue in directly generating a magnetostratigraphic sequence from the basinal Guadalupian Stratotype sections is the dissolution of magnetic carriers while the rock units were heavily saturated with oil

77

(Steiner 2006). In their sequence stratigraphic framework, Kerans and Tinker (1999) correlate the South Wells with the Grayburg and Queen formations. This places the sampled bentonite at the approximate position of the IR and any radioisotopic date determined for this deposit would provide an estimate for the age of this important global correlation tool. As mentioned above, the IR is used to constrain the timing of the dinocephalian extinction event of Lucas (2009). Dinocephalians were the first group of therapsids to show explosive diversification and they also underwent an abrupt extinction

(loss of 18 genera).

The IR has also been cited by Isozaki and others (2009a,b) as a record of a mantle superplume that caused decay in Earth’s magnetic field, leading to increased exposure to cosmic radiation, increased cloud cover, the Kamura cooling event in the Capitanian, followed by violent plume volcanism at the end of the Guadalupian and the end of the

Permian. The Kamura cooling event and the plume volcanism are then linked by the authors to the end-Guadalupian extinction, with the end-Permian volcanism linked to the mass extinction at that boundary. This hypothesis has come under criticism (Ali 2010; see also Isozaki 2010), with much of the discussion involving the absolute timing of specific events, including the IR. Specifically, Ali (2010) argued that the time required for plume ascent from the core-mantle boundary to Earth’s surface is probably > 10 myr and that much of the plume volcanism cited by Isozaki and others (2009a,b) either predates the IR or, in the case of the Emeishan LIP, occurred < 3 myr after the polarity reversal.

METHODS

Bulk bentonite samples were collected using a pick and a small trowel that were cleaned between uses. Samples were placed in one-gallon plastic bags and labeled. Care

78

was taken to ensure that only material that was suspected to be bentonite was included.

The bulk samples were soaked in water for at least 48 hours, prior to partial disaggregation by stirring. The disaggregated samples were then flushed of clay-sized particles and sieved with a 178 µm mesh screen following a slightly modified version of the methods described by Sell (2010). Heavy mineral (> 2.85 g/cm3) separation was performed using lithium heteropolytungstate. Grains for analysis were hand picked using a binocular microscope with reflected light to ensure that the most euhedral and clear zircon crystals were chosen.

The morphology of all zircon crystals are acicular.

Prior to isotopic analysis, each zircon crystal from a single sample was annealed in a quartz crucible at 900 °C for 48 hours and chemically abraded following the methods of

Mattinson (2005), in an attempt to remove domains with post-crystallization lead loss.

Following cleaning with water, HNO3, and HCl, single zircon crystals were hand selected for dissolution and each was spiked with the EARTHTIME 202Pb–205Pb–233U–235U (ET2535) tracer solution. Zircon crystals were dissolved in 200 µl sallivex capsules in a solution of

0 ~70 µl 40% HF and trace amounts of HNO3 at 210 C for 48 hours. Samples were then dried down and redissolved in 6N HCl overnight. A second dry down and redissolution step was done in 3N HCl. Anion exchange column (50 µl) chemistry (Krogh 1973) was performed, with U and Pb being collected in the same beaker. This was then dried down with a drop of 0.05M H3PO4. A Si-gel emitter was then added (Gerstenberger and Haase

1997) and the sample was mounted on a single outgassed Re filament. Isotopic analysis was performed on a Thermo Scientific TRITON thermal ionization mass spectrometer at the University of Geneva. Statistical filtering of raw mass spectrometer data was

79

performed using the software program Tripoli (Bowring et al. in press). Error propagation and age calculations were made using the Isoplot/Ex v.3 program of Ludwig (2005).

RESULTS

GM-20 Sample GM-20 was collected from an 18 cm, apple-green bentonite in the Rader

Limestone in the Patterson Hills (N 31deg 49' 28.4" W 104deg 52' 32.1") (FIGURE 3.1).

Approximately 100 crystals were separated from a ~500 ml bulk bentonite sample. Zircon crystals averaged 110 µm in length and 28 µm in width. Due to the low amounts of radiogenic lead in individual zircons crystals that resulted in short mass spectrometer runs, it was difficult to obtain reliable analyses for this sample. Four concordant analyses yield a mean 206Pb/238U age of 262.58 ± 0.45 Ma (with decay constant uncertainties; MSWD=2.0;

APPENDIX C, FIGURE 3.2). Due to the low amount of radiogenic lead in the individual zircons (typically < 3 pg), two multi-crystal analyses (GM20-9 and GM20-10) were combined with two single crystal analyses (GM20-3 and GM20-4) to calculated the weighted mean. Additional single crystal (GM20-17) and multi-crystal (GM20-14) analyses yield 206Pb/238U ages of 262.51 ± 0.25 Ma and 261.99 ± 0.44 Ma respectively, but were not included in the weighted mean calculation because of high correlation coefficients

(APPENDIX C).

GM-29 Sample GM-29 was collected from a 5 cm, dark green bentonite below the South

Wells Limestone in “Monolith Canyon” (N31deg 54.740' W104deg 46.943 )(FIGURE 3.1).

Approximately 100 crystals were separated from a ~500 ml bulk bentonite sample.

Individual zircon crystals from GM-29 were larger in size (mean length= 274 µm; mean

80

81

width= 55 µm) and had more radiogenic lead than those from GM-20. Eight concordant single crystal analyses yield a mean 206Pb/238U age of 266.50 ± 0.24 Ma (with decay constant uncertainties; MSWD=1.5; TABLE 1, FIGURE 3.3). Crystal 11 was omitted from the calculation because it was clearly biased by an inherited lead component. The analysis of GM29-6 yielded an individual 206Pb/238U age of 268.46 Ma and was omitted because it is older than the other points and outside of their analytical uncertainty.

DISCUSSION

Sample Age Interpretations

GM-20

The calculated date of 262.58 ± 0.45 Ma is interpreted to be the age of crystallization for the zircon and an estimate for the depositional age of the tephra. Although only four analyses are used in the age calculation, the results of additional single and multi-crystal analysis support an age of c. 262.5 Ma. It is this estimate that is used in the following discussion on the implications of these new data. While this was not an ideal sample to analyze, it should be pointed out that it is more likely would not have been possible prior to the widespread adoption of chemical abrasion pretreatment in the mid-2000s. Small, acicular zircons would most likely not have survived air-abrasion (Ovtcharova et al. 2006), thus any domains with post-crystallization lead loss would have remained.

GM-29

The calculated date of 266.50 ± 0.24 Ma is interpreted to be the age of crystallization for the zircon and an estimate for the depositional age of the tephra. This sample is difficult to interpret, as the analyses form a spread of data with overlapping

82

83

uncertainties and few obvious points to omit. It is possible that more precise analyses would show multiple crystallization events prior to eruption. The calculation of the

206Pb/238U age uses the youngest population of zircons, with the exception of GM29-2, which is omitted because it does not intersect concordia and is interpreted to show lead loss. Unfortunately, there are no additional dates from this section for comparison. If the date from Bowring and others (1998) were from the Manzanita as proposed in Chapter 1, it would be consistent with the age estimate of 266.50 ± 0.24 for the position in the sub-South

Wells Cherry Canyon Formation.

Stage Boundary Ages

The new U-Pb ID-TIMS data reported here call for a reassessment of the age estimates for the Guadalupian stage boundaries. Although some uncertainty exists regarding the exact stratigraphic position of the Bowring and others (1998) date, the placement in the Manzanita that was made in Chapter 1 is used here as an additional absolute age reference point. The traditional placement of this date near the base of the

Capitanian, requires the South Wells, Manzanita, and Hegler limestones to have been deposited in < 1.5 myr, which seems considerably unlikely.

Sample GM-20 is from 20 meters above the base of the Rader Limestone in the

Patterson Hills (FIGURE 3.1) and is linked to the J. posterserra conodont zone through a bentonite correlation discussed in Chapter 2. The age of the Rader Limestone is estimated at c. 264 Ma by Menning and others (2006, Figure 4). In light of the results for GM-20, the position of the Rader Limestone should be shifted up to c. 262.5 Ma. This shift implies that the age of the Pinery Limestone, and the base of the Capitanian, should also be at a younger position. While no data are available to directly date the Pinery, results from GM-29 and

84

the likely position of the 265.3 Ma age of Bowring and others (1998) provide some additional constraint with which to make an estimate. Based on this information, it seems reasonable that the base of the Capitanian is <264 Ma. A tentative assignment of 263.5 Ma seems appropriate given the apparent trend in ages of the carbonate members of the

Delaware Mountain Group. This marks a shift to a position 1.5 myr younger than that of

Menning and others (2006) and 2.3 myr younger than that of Ogg and others (2008).

Current estimates on the age of the base of the Wordian might also be too old. In both Ogg and others (2008) and Menning and others (2006), the estimate is 268 Ma.

Sample GM-29 comes from a bentonite that occurs below the base of the South Wells

Limestone in a drainage referred to here as “Monolith Canyon” (FIGURE 3.1). This drainage is adjacent to Nipple Hill and the position of this sample was traced down from the

Capitanian GSSP. The sample occurs stratigraphically above the Getaway Limestone, and therefore above the base of the Wordian (FIGURE 3.4). As the sample is from a position below the base of the Capitanian and above the base of the Wordian, it is clearly within the latter stage. It is also within the J. asserrata conodont zone, which defines the base of the

Wordian and spans the entire stage (Glenister et al. 1999; Wardlaw et al. 2004).

The age estimate for GM-29 near the base of the South Wells, means the Roadian-

Wordian boundary is at least 266.50 ± 0.24 Ma, but likely older. An age estimate for the base of the Wordian, depends of how much time elapsed between the Getaway Limestone and the South Wells Limestone. As the date from GM-29 is the only one reported for this interval (including the Roadian) it is difficult to make a precise age assignment for the boundary. The Wordian GSSP is just below the top of the Getaway Limestone at Guadalupe

Pass (Glenister et al. 1999) (FIGURE 3.1). Given this, the base of the Wordian is tentatively

85

86

estimated to be 270 Ma (FIGURE 3.4). This places 500,000 kyr between the top of the

Getaway Limestone and the base of the South Wells Limestone. If the estimates of 268 Ma for the Wordian base are accepted, the new data reported here would mean that 1.5 myr elapsed between the two members. An older estimate for the Wordian base would mean an even longer stage duration than estimated in recent time scales (Menning et al. 2006;

Ogg et al. 2008). Until additional bentonites, near the Roadian-Wordian boundary, can be located and dated, estimates on its age will remain relatively imprecise. An additional bentonite is present in the sub-South Wells interval, but it is 7 cm below below GM-29 and dating this sample would not provide any additional constraint of the base of the Wordian, rather it could, at best, corroborate the age estimate for GM-29.

The base of the Wuchiapingian Stage (i.e. Guadalupian/Lopingian boundary) is also poorly constrained temporally. Its current age estimate is based, in part, on the placement of the 265.3 ± 0.2 Ma date of Bowring and others (1998) very near the base of the

Capitanian (Menning et al. 2006). If the date of Bowring and others (1998) is from the

Manzanita as suggested in Chapter 1, these data appear to form a trend, with the lower carbonate members of the Delaware Mountain Group occurring roughly every 1 myr (i.e.

South Wells=266.5 Ma, Manzanita=265.5 Ma, Hegler=? (264.5 Ma), Pinery=? (263.5 Ma),

Rader=262.5 Ma). Given the lack of age constraint for the base of the Wuchiapingian, the data here can be used to provide an estimate. If it assumed that the observed trend continues with the younger members, the uppermost Guadalupian unit (Reef Trail

Member) would be 259.5 Ma, and possibly younger if the informal McKittrick Canyon

Limestone of Wilde and others (1999) is included. Rush and Kerans (2010) recently placed the McKittrick Canyon Limestone-Reef Trail Member into four HFSs, with the Lamar

87

spanning two. An estimate of 259.5 Ma is consistent with Mundil and others (2004), who also conclude that this boundary is < 260 Ma, based on a series of ID-TIMS zircon dates from bentonites in the Shangsi site of central China. The estimate of c. 259.5 Ma is 1 myr younger than the estimate for the base of the Wuchiapingian by Menning and others

(2006). Their estimate is based in part on the 265.3 ±0.2 Ma age of Bowring and others

(1998) (at a position near the base of the Capitanian) and on radioisotopic ages from the

Late Changsingian and Permian-Triassic boundary interval. If the boundary is not younger, as the data here imply, then the carbonate members above the Rader Limestone were deposited at the faster rate than those of the lower Bell Canyon and Cherry Canyon formations.

Though no radioisotopic data from bentonites have been reported near this boundary, the basalts of the Late Guadalupian Emeishan (LIP) in South China have been dated by Zhong and Zhu (2006), who used ID-TIMS zircon data obtained from two mafic intrusions (259.3 ± 1.3 and 260.7 ± 0.8 Ma) to suggest an age estimate of 260 Ma for the

Emeishan flood basalts. This estimate is consistent with the SHRIMP age of 259.3 ± 3 Ma for the Emeishan LIP reported by Zhou and others (2002). Based on magnetobiostratigraphic correlation, the Emeishan basalts predate the base of the

Wuchiapingian by at least two conodont (Ali et al. 2002) and were active beginning in the J. altudaensis zone (Bond et al. 2010). As the Emeishan basalts are thought to have been deposited relatively rapidly (<1 myr) (Ali et al. 2002; Zheng et al. 2010), the

260.7 ± 0.8 Ma age would seem like a good minimum estimate, given it is from a cross cutting feature. This would support a younger age of the base of the Wuchiapingian. This is consistent with Mundil and others (2004), who also conclude that this boundary is < 260

88

Ma, based on a series of ID-TIMS zircon dates from bentonites in the Shangsi site of central

China.

It is difficult to quantify the how much younger this boundary should be based on the lack of constraint on the several conodont zones in the upper Capitanian. While there appears to be a strong correlation of conodonts zones between West Texas and South

China (Mei et al., 1994, 1998; Wardlaw, 2000; Wardlaw et al. 2004), it is possible that the

FAD for the uppermost Guadalupian conodont hongshuiensis is diachronous (Jin et al. 2006). Depending on how rapidly these species evolved, the base of the Wuchiapingian may yet be close to c. 260.5 Ma. The age of c. 259.5 Ma indicated in FIGURE 3.4, should be treated as a loose estimate. Altered tephras are present near the boundary at the

Wuchiapingian GSSP have been sampled for radioisotopic dating (Jin et al. 2006), but age estimates have yet to be reported. Additional work on this section, or one that has correlative events or conodont zones, may allow for a better-constrained estimate for the base of this stage.

With changes in the age estimates of stage boundaries come changes in their durations. The data here suggest that the duration of the Wordian (c. 3.5 myr) is probably closer to that of the Capitanian (c. 4 myr) than previously understood. The Guadalupian time scales in Wardlaw and others (2004) and Ogg and others (1998) both estimate that duration of the Capitanian to be more than twice as long as that of the Wordian (FIGURE

3.4). This can have a significant effect on rates of geologic and biologic phenomena calculated at the stage scale. Examples of this are per-capita rates of extinction, origination, and preservation, which included the duration of the interval being examined in the calculations (Foote 2000). In a study of the end-Guadaluian extinction in marine

89

invertebrates, Clapham and others (2009) calculated per-capita extinction and origination rates for the (Late Early Permian) through the (latest Permian) at the stage level and concluded that, rather than a sharp decrease in diversity at the end of the Guadalupian, there was a gradual decrease from the Wordian to the end of the

Chinghsingian, driven primarily by extrememly low origination rates in the Capitanian and

Wuchiapingian. Their calculations used the Guadalupian stage duration estimates from

Wardlaw and others (2004), where the Wordian is less than half a long as the Capitanian.

As discussed above, the current study suggests that the duration of the Wordian is probably closer to that of the Capitanian than previously understood. A longer duration for the Wordian and a shorter duration for the Capitanian, decrease extinctions rates for the former and increase them for the latter. After recalculating the per-capita extinction rates from Clapham and others (2009) (standardized data) using the new estimated stage durations, the data are plotted along with those from Clapham and others (2009) (FIGURE

3.5a). The recalculated data show a gradual increase to a peak in the Capitanian, followed by a decrease across the end-Guadalupian into the Wuchiapingian. This contrasts with the original trend of gradually decreasing rates from a Wordian peak to a minimum in the

Wuchiapingian. It also places the per-capita trend in closer agreement with that of the per- taxon trend that shows a moderately elevated extinction rate during the Capitanian.

Recalculations of per-capita origination likewise show a decrease in the Wordian and an increase in the Capitanian (FIGURE 3.5b). In this case, the Wordian value shifts down from a maximum position, but remains higher than both the Roadian and Capitanian values. The Capitanian also makes a shift, but it is a very slight increase and its value remains conspicuously low when compared to the other stages. Despite these changes, the

90

overall trend of the origination rates is more or less the same as the original. That the origination rate in the Capitanian remains low is notable, as it is implicated in artificially elevating the per-taxon extinction rate for this stage (Clapham et al. 2009).

Based on the new stage durations estimated here, the end-Guadalupian decrease in diversity might have been affected more by elevated extinction rates in the Capitanian than the original data set suggest, though origination rates remain very low and could still be the dominant factor. The shift in peak Guadalupian extinction from the Wordian to the

Capitanian aligns it with the low origination rate, and suggests that any causal mechanism

91

should either explain both phenomena, or at least be consistent with each. An extinction peak and origination low is more permissive of the suggestion that the Emeishan LIP as a potentially cause (e.g. Wignall 2001; Ali et al. 2002; Wignall et al. 2009) than had there been decreased origination and no elevated extinction. It is also consistent with models relating extinction with cooling during the Kamura event (Isozaki et al. 2007b), which might also have been the cause of reduced origination rates (Clapham et al. 2009). While new estimates for stage durations clearly have an effect on the per-capita extinction and origination trends in the Guadalupian, the overall signal is still dominated by peak extinction in the Changhsingian. Additional radioisotopic data should refine this interval further and provide better temporal constraint for these metrics.

Timing of the Kamura Event

Based on the correlations of Bond and others (2010), the date determined for GM-

20 provides a maximum age estimate for the beginning of the Kamura Event, as it likely occurs lower in the J. postserrata conodont zone than the elevated δ13C values. The age is younger than estimates based on the GTS 2004 in previous studies (e.g. Isozaki et al.

2007b; Isozaki et al. in press, fig 3). This is notable as it means that global environmental change (i.e. cooling), implied by the elevated δ13C, occurred closer to the “end-Guadalupian” mass extinction and the Permo-Triassic boundary than previously understood. It also places the onset of high amplitude fluctuations in δ13C closer to the Permo-Triassic boundary, indicating that the transition to a post-Paleozoic global climate and oceanography, as suggested by Isozaki and others (2007b), may have happened earlier.

The duration of the Kamura event was probably on the lower end of the 3-4 myr estimate of Isozaki and others (2007b) or even shorter, as it appears to begin after the c. 262.5 Ma

92

date from GM-20 (based on its placement in the late J. postserrata zone by Bond et al. 2010) and ends near the onset of Emeishan volcanism at c. 260 Ma (Bond et al. 2010). The addition of C isotopic measurement closer to the GM-20 sampling point at GMNP, could allow for further constraint on the timing and duration of this event.

Timing of the Illawarra Reversal

The age determination for sample GM-29 provides an estimate for the age of the IR at 266.5 Ma. This estimate follows the placement of the IR by Steiner (2006) in the

Grayburg or Queen formations on the shelf. These formations have been correlated to the

South Wells Member by several workers (e.g. Wilde 1975; Beaubouef et al. 1999; Kerans and Tinker 1999). Now the IR can be used to provide an age estimate for sections unable to be dated, either directly by bentonites, or indirectly by biostratigraphy.

The dinocephalian extinction event (DEE) precedes the IR (Lucas 2009) and is therefore older than 266.5 Ma. Based on current data, this places this terrestrial extinction at least 6.5 myr before the end of the Guadalupian, and 3 myr before the proposed marine extinctions in South China and Japan. The lack of coincidence between the DEE and either a mid-Capitanian or end Guadalupian extinction pulse, suggests that an abrupt mechanism to explain both a terrestrial and marine extinction is unnecessary. This new time constraint also places the DEE c. 3 myr earlier than the onset of the Kamura event and the beginning of the high amplitude fluctuations in δ13C values at the end of the Permian.

If the IR does mark a mantle superplume, as Isozaki (2009 a,b) suggests, the time constraint presented here places it c. 6 myr before the Emeishan LIP and c. 3 myr before the Kamura event. This would seem to preclude the IR as a fingerprint of a direct trigger of the Kamura event, in the form of geomagnetic instability and increased cosmic radiation, as

93

has been proposed (Isozaki et al. 2007a,b; Isozaki 2009 a,b). This new constraint also places the amount of time elasped between the superplume and the Emeishan LIP within the 2-20 myr estimated by Isozaki (2010), but still less than the >10 myr requirement argued by Ali (2010). Whether a mantle superplume did initiate the apparent clustering of major global events at the end of the Paleozoic cannot be definitively answered here, but the new data presented above provide important reference points that can be used to calibrate events in future studies.

CONCLUSIONS

The calculation of new U-Pb ID-TIMS dates in the Guadalupian type area indicate the need for changes to the geologic time scale and the temporal relationships of global events.

An age estimate of 263.5 Ma for the base of the Capitanian Stage is made based on a bentonite in the lower half of the stage defining J. postserrata conodont zone dated at 262.5

Ma. This means the age of the Wordian-Capitanian boundary is younger than estimated in recent time scales. It also decreases the duration of the Capitanian to c. 4 myr and provides a maximum age estimate for the globally correlatable Kamura cooling event. The mass extinction that has been interpreted to occur within this event is one conodont zone above the dated bentonite and therefore no older than 262.5 Ma.

The new radioisotopic date from below the South Wells Limestone provides an age estimate of 266.5 Ma for the globally correlatable Illawarra reversal. It also further supports the conclusion from Chapter 1, that the existing Guadalupian radioisotopic age of

265.3 ± 0.2 Ma, should be placed in the Manzanita Limestone, rather than a point nearer to the base of the Capitanian. Although it is less certain than the basal Capitanian shift, the

Roadian-Wordian boundary might also be younger than current estimates.

94

It has also been proposed that the Guadalupian lacks an abrupt extinction event, but rather shows a gradual decrease in diversity. The alterations to the Capitanian stage proposed here indicate that if a descrease in the diversity of taxa during the Guadalupian made them more susceptible to collapse, it happened closer to the main extinction pulse at the Permo-Triassic boundary than previously understood. The implied changes in durations for both the Wordian and Capitanian stages serve as reminders that calculated rates of geologic and biologic events are dependent on the quality of temporal constraint available for a given interval. Additional U-Pb ID-TIMS dates will provide a check on the interpretations here and allow for further temporal refinement for this important period in

Earth’s history.

ACKNOWLEDGEMENTS

U. Schaltegger is thanked for allowing access his laboratory at the University of

Geneva. N. Bose and J. Wilkins assisted with sample collection. L. Fall and S. Marcus located the bentonite in the Rader Limestone at Back Ridge and alerted me to its presence. T.

Phillips helped with drafting some of the figures. Graduate student research funding was provided by the American Association of Petroleum Geologists (Ohio Geological Society

Named Grant), the Clay Minerals Society, the Geological Society of America, and the

Graduate Student Governance Association and Department of Geology at the University of

Cincinnati.

REFERENCES Ali, J.R., 2010. Comment on “Illawarra Reversal: the fingerprint of a superplume that triggered Pangean breakup and the end-Guadalupian (Permian) mass extinction” by Yukio Isozaki. Gondwana Research, 17:715-717

95

Ali, J.R., Thompson, G.M., Song, X.-Y. and Wang, Y.-L., 2002. Emeishan Basalts (SW China) and the ‘end-Guadalupian’ crisis: magnetostratigraphic constraints. Journal of the Geological Society, 159:21–29.

Beaubouef, R.T., Rossen, C., Zelt, F.B., Sullivan, M.D., Mohrig, D.C., and Jennette, D.C., 1999. Deep-water sandstones, Brushy Canyon Formation, west Texas: American Association of Petroleum Geologists Continuing Education Course Notes Series #40, 62p.

Bond, D.P.G., Wignall, P.B., Wang, W., Izon, G., Jiang, H.-S., Lai, X. -L., Sun, Y. –D., Newton, R.J., Shao, L.-Y., Védrine, S. and Cope, H., 2010. The mid-Capitanian (Middle Permian) mass extinction and carbon isotope record of South China. Palaeogeography, Palaeoclimatology, Paleoecology, 292:282-294.

Bowring, J.F., McLean, N.M., and Bowring, S.A., 2011. Engineering cyber infrastructure for U-Pb geochronology: Tripoli and U-Pb_Redux. Geochemistry, Geophysics, Geosystems, in press.

Bowring, S.A., Davidek. K., Erwin, D.H., Jin, Y.G., Martin, M.W. and Wang, W., 1998. U-Pb zircon geochronology and tempo of the end-Permian mass extinction: Science, 280:1039-1045.

Clapham, M.E., Shen, S. and Bottjer, D.J., 2009. The double mass extinction revisited: reassessing the severity, selectivity, and causes of the end-Guadalupian biotic crisis (Late Permian). Paleobiology, 35:32-50.

Crowely, J.L., Schoene, B. and Bowring, S.A., 2007. U-Pb dating of zircon in the Bishop Tuff at the millennial scale. Geology, 35:1123-1126; doi:10.1130/G24017A.

Foote, M. 2000. Origination and extinction components of taxonomic diversity: general problems. In D. H. Erwin, and S. L. Wing, eds. Deep time: Paleobiology’s perspective. Paleobiology 26(Suppl. to No. 4):74–102.

Gerstenberger, H. and Haase, G., 1997. A highly effective emitter substance for mass spectrometric Pb isotopic ratio determinations. Chemical Geology, 136:309-312.

Glenister, B.F., Wardlaw, B.R., Lambert, L.L., Spinosa, C., Bowring, S.A., Erwin, D.H., Menning, M., and Wilde, G.L., 1999. Proposal of Guadalupian and Component Roadian, Wordian and Capitanian Stages as International Standards for the Middle Permian Series. Permophiles, 34:3-11.

Irving, E. and Parry, L.G., 1963. The magnetism of some Permian rocks from New South Wales. Geophysical Journal of the Royal Astronomical Society, 7:395–411.

Isozaki, Y., Kawahata, H. and Ota, A., 2007a. A unique carbon isotope record across the Guadalupian-Lopingian (Middle-Upper Permian) boundary in mid-oceanic

96

paleoatoll carbonates: the high-productivity “Kamura event” and its collapse in Panthalassa. Global Planetary Change, 55:21–38.

Isozaki, Y., Kawahata, H. and Minoshima, K., 2007b. The Capitanian (Permian) Kamura Cooling Event: the beginning of the Paleozoic– transition. Palaeoworld, 16:16–30.

Isozaki, Y., 2009a. Illawarra Reversal: the fingerprint of a mantle superplume triggered Pangean breakup and end-Permian mass extinction. Gondwana Research, 15:421– 432.

Isozaki, Y., 2009b. Integrated plume winter scenario for the double-phased extinction during the Paleozoic–Mesozoic transition: the G-LB and P-TB events from a Panthalassan perspective. Journal of Asian Earth Sciences, 36:459-480.

Isozaki, Y., 2010. Reply to the comment by J.R. Ali on “Illawarra Reversal: the fingerprint of a superplume that triggered Pangean breakup and the end-Guadalupian (Permian) mass extinction” by Yukio Isozaki. Gondwana Research, 17:718-720.

Isozaki, Y., Aljinović, D., and Kawahata, H., The Guadalupian (Permian) Kamura event in European Tethys. Palaeogeography, Palaeoclimatology, Palaeoecology (2010), doi:10.1016/j.palaeo.2010.09.034

Jaffey, A.H., Flynn, K.F., Glenenin, L.E., Bentley, W.C. and Essling, A.M., 1971. Precision measurement of half-lives and specific activities of 235U and 238U. Physical Review, C4:1889-1906.

Jin, Y.-G., Zhang, J. and Shang, Q.-H., 1994. Two phases of the end-Permian mass extinction. In: Embry, A.F., Beauchamp, B., Glass, D.J., (Eds.), Pangea: Global Environments and Resources, vol. 17. Memoir Canadian Society of Petroleum Geologists, pp. 813–822.

Jin, Y.-G., Shen, S.-Z., Henderson, C.M., Wang, X.-D., Wang, W., Wang, Y., Cao, C.-Q. and Shang, Q.-H., 2006. The Global Stratotype Section and Point (GSSP) for the boundary between the Capitanian and Wuchiapingian Stage (Permian). Episodes, 29:253–262.

Kerans, C. and Tinker, S.W., 1999. Extrinsic controls on development of the Capitan Reef complex, in Saller et al., eds., Geologic Framework of the Capitan Reef, Tulsa, OK, SEPM Special Publication No. 65, p. 15-36.

Korte, C., Jasper, T., Kozur, H.W., Veizer, J., 2005. δ18O and δ13Ccarb of Permian brachiopods: a record of seawater evolution and continental glaciation. Palaeogeography, Paleoclimatplogy, Palaeoecology, 224: 333–351.

Lambert, L. L., Wardlaw, B. R., Nestell, M. K. and Nestell, G. P., 2002. Latest Guadalupian (Middle Permian) conodonts and foraminifers from West Texas. Micropaleontology,

97

48:343-364.

Lambert, L.L., Bell, G.L. Jr., Fronimos, J.A., Wardlaw, B.R., and Yisa, M.O., 2010. of a more complete Reef Trail Member section near the type section, latest Guadalupian Series type region. Micropaleontology, 56:233-253.

Lucas, S.G., 2009. Timing and magnitude of tetrapod extinctions across the Permo-Triassic boundary. Journal of Asian Earth Sciences, 36:491-502.

Ludwig, K.R., Isoplot/Ex. V.3, USGS Open-File Repository (2005).

Mattinson, J.M., 2005, Zircon U-Pb chemical abrasion (“CA-TIMS”) method: Combined annealing and multi-step partial dissolution analysis for improved precision and accuracy of zircon ages. Chemical Geology, 220:47-66.

Mei, S.-L., Jin, Y.-G. and Wardlaw, B.R., 1994. Succession of conodont zones from the Permian “Kuhfeng” Formation, Xuanhan, Sichuan and its implication in global correlation. Acta Palaeontology Sinica, 33:1-23.

Mei, S.-L., Jin, Y.-G. and Wardlaw, B.R., 1998. Conodont succession of the Guadalupian– Lopingian boundary strata in Laibin of Guangxi, China and West Texas, USA. Palaeoworld, 9:53–76.

Menning, M., 2000. Magnetostratigraphic results from the Middle Permian Type Section, Guadalupe Mountains, west Texas. Permophiles, 37:16.

Menning, M., Alekseev, A.S., Chuvashov, B.I., Davydov, V.I., Devuyst, F.-X., Forke, H.C., Grunt, T.A., Hance, L., Heckel, P.H., Izokh, N.G., Jin, Y.-G., Jones, P.J., Kotlyar, G.V., Kozur, H.W., Nemyrovska, T.I., Schneider, J.W., Wang, X.-D., Weddige, K., Weyer, D. and Work, D.M., 2006. Global time scale and regional stratigraphic reference scales of Central and West Europe, East Europe, Tethys, South China, and North America as used in the Devonian-Carboniferous-Permian Correlation Chart 2003 (DCP 2003). Palaeogeography, Palaeoclimatology, Palaeoecology, 240:318-372.

Mundil, R., Ludwig, K.R., Metcalfe, I. and Renne, P.R., 2004. Age and timing of the Permian mass extinction: U/Pb dating of closed-system zircons. Science, 305:1760–1763.

Noble, P.J., Poulson, S.R., Maldonado, A.L., Lambert, L.L., Wardlaw, B.R., Mestell, M.K., and Bell, G.L. 2009. Latest Guadalupian carbon isotope record from the Reef Trail Member of the Bell Canyon Formation, West Texas, USA. Abstracts with Programs, Geological Society of America, 41(2):28.

Noble, P.J., Naraoka, H., Poulson, S.R., Fukui, E., Jin, Y. and O’Connor, S., 2011. Paleohydrographic influences on Permian Basin radiolarians in the Lamar Limestone, Guadalupe Mountains, West Texas, elucidated by organic biomarker and stable isotope geochemistry. Palaios, 26:180-186.

98

Ogg, J. G., Ogg, G., and Gradstein, F.M., 2008. The Concise Geologic Time Scale. Cambridge: Cambridge University Press, 184 pp.

Peterson, D.N., and Nairn, A.E.M., 1971. Palaeomagnetism of Permian red beds from the South-western United States. Geophysical Journal of the Royal Astronomical Society, 23:191-207.

Retallack, G.J., Metzger, C.A., Greaver, T., Jahren, A.H., Smith, R.M.H. and Sheldon, N.D., 2006. Middle-Late Permian mass extinction on land. Geological Society of America Bulletin, 188:1398-1411.

Rush, J. and Kerans, C., 2010. Stratigraphic response across a structurally dynamic shelf: the latest Guadalupian composite sequence at Walnut Canyon, New Mexico, U.S.A.. Journal of Sedimentary Research, 80:808-828.

Schmitz, M.D. and Schoene, B., 2007. Derivation of isotope ratios, errors, and error correlations for U-Pb geochronology using 205Pb-235U-(233U)-spike isotope dilution thermal ionization mass spectrometric data. Geochemistry, Geophysics, Geosystems, 8:Q08006, doi:10.1029/2006GC001492.

Sell, B.K., 2010, “Apatite trace element tephrochronology” [unpublished Ph.D. dissertation]: Syracuse University, Syracuse, New York, 306 p.

Słowakiewicz, M., Kiersnowski, H. and Wagner, R., 2009. Correlation of the Middle and Upper Permian marine and terrestrial sedimentary sequences in Polish, German, and USA Western Interior Basins with reference to global time markers. Paleoworld, 18:193-211.

Stanley, S.M. and Yang, X., 1994. A double mass extinction at the end of the Paleozoic era. Science, 266:1340-1344.

Steiner, M.B., 2006. The magnetic polarity time scale across the Permian–Triassic boundary. In: Lucas, S.G., Cassinis, G., Schneider, J.W. (Eds.), Non-marine Permian biostratigraphy and biochronology. Geological Society of London Special Publica- tion, 265, pp. 15–38.

Walker, D.A., Golanka, J., Reid, A. and Reid, S., 1995. The effects of paleolatitude and paleogeography on carbonate sedimentation in the late Paleozoic, in Huc, A., ed., Paleogeography, Paleoclimate, and Source Rocks: Tulsa, American Association of Petroleum Geologists Studies in Geology 40, p. 133-155.

Wardlaw, B.R., 2000. Guadalupian conodont biostratigraphy of the Glass and Del Norte Mountains, in Wardlaw, B.R., Grant, R.E., and Rohr, D.M., eds., The Guadalupian Symposium: Smithsonian Contributions to Earth Sciences no. 32, p. 37-88.

99

Wardlaw, B.R., Davydov, V. and Gradstein, F.M., 2004. The Permian Period, in Gradstein, F.M., Ogg, J.G., and Smith, A.G., eds., A Geologic Time Scale 2004. 249-270. Cambridge: Cambridge University Press.

Wignall, P.B., 2001. Large igneous provinces and mass extinctions. Earth-Science Reviews, 53:1–33.

Wignall, P.B., Sun, Y.-D., Bond, D.P.G., Izon, G., Newton, R.J., Védrine, S., Widdowson, M., Ali, J.R., Lai, X.-L., Jiang, H.-S., Cope, H. and Bottrell, S.H., 2009. Volcanism, mass extinction, and carbon isotope fluctuations in the Middle Permian of China. Science, 324:1179–1182.

Wilde, G.L., 1975. Fusulinid-defined Permian Stages, in Permian Exploration, Boundaries, and Stratigraphy: Tulsa Society of Economic Paleontologists and Mineralogists and West Texas Geological Society, Permian Basin Section, Publication 75-65, p. 67-83.

Zheng, L., Yang, Z., Tong, Y. and Yuan, W., 2010. Magnetostratigraphic constraints on two- stage eruptions of the Emeishan continental flood basalts. Geochemistry, Geophysics, Geosystems, 11: Q12014, doi:10.1029/2010GC003267.

Zhong, H. and Zhu, W.-G., 2006. Geochronology of layered mafic intrusions from the Pan-Xi area in the Emeishan large igneous province, SW China. Miner Deposita, 41:599- 606.

Zhou, M. F., Malpas, J., Song, X. Y., Robinson, P. T., Sun, M., Kennedy, A. K., Lesher, C. M. and R. R. Keays, 2002. A temporal link between the Emeishan large igneous province (SW China) and the end‐Guadalupian mass extinction. Earth Planetary Science Letters, 196:113–122.

Ziegler, A.M., Hulver, M.L. and Rowley, D.B., 1997. Permian world topography and climate, in Martini, I.P., ed., Late Glacial and Postglacial Environmental Changes: Quaternary, Carboniferous-Permian, and Proterozoic: Oxford University Press, Oxford, p. 111- 146.

100

Appendix A

Apatite minor, trace, and REE EMPA data for Manzanita Limestone bentonite samples as weight percents. Smithsonian apatite standard (NMNH 104021) averages are from the SU sessions.

101

102

103

104

105

106

107

108

109

110

111

112

113

Appendix B

Apatite minor, trace, and REE EMPA data for Yates and Bell Canyon formation bentonites as weight percents. Smithsonian apatite standard (NMNH 104021) averages are from the UKY sessions.

114

115

116

117

Appendix C

U-Pb ID-TIMS isotopic data

118

119