MALLETT-THESIS-2021.Pdf

On the Origin and Paleoclimate Implications of Paleosols from the Blackwater Draw Formation at Bushland Playa Near Amarillo, Tx

by Colton Mallett, B.S. A Thesis In Geosciences Submitted to the Graduate Faculty of Texas Tech University in Partial Fulfillment of the Requirements for the Degree of

MASTER OF SCIENCE

Approved

Dr. Branimir Šegvić Chair of the Committee

Dr. Dustin Sweet

Dr. Neo McAdams

Mark Sheridan Dean of the Graduate School

May, 2021

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ACKNOWLEDGMENTS

The realization of this project is the results of many hours of work, contributed by many people. I would like to acknowledge and articulate my utmost appreciation for the assistance of these individuals.

I would like to extend my greatest appreciation to my advisor Dr. Branimir

Šegvić, who supplied unmatched academic and personal support throughout my time as his student. His unparalleled understanding of the intricacies of the project were instrumental in its success. I am beyond thankful for his understanding of my personal circumstances and the motivation he provided me. Truly, the success of this project would not have been possible without his unwavering support.

I would like to extend my appreciation and admiration to Giovanni Zanoni for his support in this project. I am incredibly appreciative for his willingness to provide hours of instruction and guidance. He has been an amazing friend and mentor.

I would like to thank my committee members, Dr. Neo McAdams and Dr. Dustin

Sweet for their valuable guidance and advice. Dr. Dustin Sweet’s willingness to share his expertise and understanding of the Blackwater Draw Formation and Southern High Plains was invaluable.

I would like to acknowledge Dr. Melanie Barnes and Dr. Kevin Werts for their instructions and assistance with my geochemical analysis conducted in the Texas Tech

University Geoanalytical Lab.

I would like to acknowledge and thank the Desk and Derrick Educational Trust Fund for their generous grant. I would also like to recognize and thank Mr. Michael B. Portnoy

Texas Tech University, Colton Mallett, May 2021 for his generous funding. Lastly, I would like to acknowledge the Texas Tech University

Department of Geosciences Scholarship for their generous financial assistance. These funds were instrumental in financing this project.

Lastly, I would like to thank my friends and family for their unwavering support

during this endeavor.

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TABLE OF CONTENTS

ACKNOWLEDGMENTS ...... ii ABSTRACT ...... vi LIST OF TABLES ...... vii LIST OF FIGURES ...... viii 1. INTRODUCTION ...... 1 2. BACKGROUND ...... 6 2.1 Current Landscape and Climate of the Southern High Plains ...... 6 2.1.1 Modern Geomorphology ...... 6 2.1.2 Modern Climate ...... 9 2.2 Geomorphic Evolution and Pertinent Strata of the Southern High Plains ...... 10 2.2.1 Development of the Southern High Plains (SHP) ...... 10 2.2.2 Ogallala Formation ...... 13 2.2.3 Lacustrine Strata ...... 15 2.3 Blackwater Draw Formation ...... 16 2.3.1 Physical Characteristics of the Blackwater Draw Formation ...... 16 2.3.2 Depositional Models ...... 18 2.3.3 Age Estimates ...... 19 2.3.4 Recent Observations ...... 21 3. MATERIALS AND METHODS ...... 24 3.1 Core Procurement ...... 24 3.2 Core Preparation ...... 25 3.3 Sampling ...... 26 3.4 Optical Microscopy ...... 28 3.5 Loss on Ignition ...... 28 3.6 Glass Disc Fusions ...... 29 3.7 X-Ray Fluorescence ...... 29 3.8 Laser Ablation Inductively Coupled Plasma Mass Spectrometry ...... 30 3.9 X-Ray Powder Diffraction ...... 31 3.10 Interpretation and modeling of XRD patterns ...... 31 3.11 Scanning Electron Microscopy – Energy Dispersive X-Ray Spectroscopy ...... 32 3.12 K-Ar Dating of Illite ...... 33 4. RESULTS ...... 35

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4.1 Optical Petrography ...... 36 4.2 X-Ray Diffraction of Whole-rock Fraction ...... 38 4.3 X-Ray Diffraction of Clay Fraction ...... 45 4.3.1 Qualitative Identification of Phases in the Clay Fraction ...... 46 4.3.2 Grouping Clay Mineral Assemblages...... 48 4.3.3 Modeling and Qualitative Description of Clay Minerals ...... 51 4.4 Scanning Electron Microscopy – Energy Dispersive Spectroscopy ...... 56 4.5 X-Ray Fluorescence ...... 61 4.6 Laser Ablation Inductively Coupled Plasma Mass Spectrometry ...... 67 4.7 K-Ar Dating of Illite ...... 70 5. DISCUSSION ...... 72 5.1. Clay-Mineralogy-Based Pedogenetic Scheme ...... 72 5.1.1. Surface Soil...... 77 5.1.2. Paleosol I ...... 77 5.1.3. Paleosol II ...... 78 5.1.4. Paleosol III ...... 79 5.1.5. Paleosol IV ...... 79 5.1.6. Paleosol V ...... 80 5.1.7. General Remarks ...... 80 5.2. Paleosols Geochemistry ...... 85 5.2.1. Major elements signature...... 86 5.2.2. Minor elements signature ...... 86 5.2.3. General Remarks ...... 87 5.3. Paleosols and Inferences on Quaternary Climate Across the Southern High Plains ...... 88 5.4. Implications of K-Ar Ages of Detrital Illite and Whole-Rock Trace Element Chemistry ...... 91 5.5. Final Remarks ...... 96 6. CONCLUSIONS ...... 97 REFERENCES ...... 99 APPENDICES ...... 111

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ABSTRACT

The Blackwater Draw Formation (BDF) is composed of a series of Quaternary aged, stacked loess- paleosol couplets, that form a relatively thin mantle over the ~120,000 km2

Southern High Plains (SHP) in west Texas and eastern New Mexico. The current geomorphology of the Southern High Plains began to develop around approximately 1.6

Ma as large river systems established themselves, isolating the SHP from surrounding landforms, thus allowing eolian deposition to dominate. This study utilizes clay mineral speciation and variations in mineralogy and geochemistry to construct a pedogenic framework and extrapolate on prevailing paleoclimate conditions during the Quaternary

Period from a 14 m core acquired near Amarillo TX. Furthermore, K-Ar dating of detrital and authigenic illite throughout the extent of the core was utilized to determine changes or variations in provenance and paleoclimate.

Geochemistry, mineralogy, K-Ar ages of authigenic clay minerals and their speciation indicate the BDF at Bushland is composed of 5 buried paleosols and a surface soil. The formation itself at this locality was further divided into an upper and lower member. The paleosols in the upper member, Paleosol I through IV were deposited under more temperate climatic conditions, while the lower member, composed of Paleosol V, was deposited under more arid conditions. A difference in detrital illite ages and variations in geochemistry indicate a shift or change in sediment input from a secondary provenance.

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LIST OF TABLES

4.0.1. Summary of Samples and Analyses……………………………………………...35 4.2.1. Quantification of the Whole Rock X-Ray Diffraction Data……………………..44 4.3.1. Summarized Sybilla modelling parameters………………………………….…52 4.3.2. Quantification of the Clay Fraction X-Ray Diffraction Data……………………55 4.4.1. Idealized chemical composition of pertinent minerals………………………..…56 4.4.2. Representative EDS analyses of select minerals……………………………...…58 4.4.3. Representative EDS analyses of clay mineral phases from selected samples...... 60 4.5.1. Normalized XRF chemistry of paleosols from the Bushland Core……………...65 4.5.2. Normalized XRF chemistry of paleosols from the Bushland Core……………...66 4.7.1. Potassium- Argon data of the size fractions from the selected samples…………71

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LIST OF FIGURES

1.1. Extent of Southern High Plains………………………………………...……….…2 1.2. Location of the Bushland Playa Core Sample………………………...…………..5 2.1.1. Geographic Extent of the Southern High Plains……………………..…….…..….7 2.1.2. Elevation profiles of the Southern High Plains……………………………………8 2.1.3. Climate classifications of the US…………………………………………...... ….10 2.2.1. Simplified stratigraphy of the Southern High Plains……………………...….….12 2.3.1. Classification of surface sediments on the Southern High Plains………….…….17 2.3.2. Previously described stratigraphic columns of the Blackwater Draw Formation..23 3.1. Location of the bushland playa core sample…………………….…….…………25 3.2. Core sample locations…………………………….…...…………………………27 4.1.1. Photomicrographs of select samples……………………………………..………38 4.2.1. Diffractograms of all analyzed samples……………………………….…………40 4.2.2. Bar charts of quantified whole rock mineralogy…………………………………43 4.3.1. Representative XRD patterns of clay fraction samples………………….………46 4.3.2. Groupings of clay mineral assemblages based on XRD patterns…………..……50 4.3.3. Sybilla Modelling of the experimental XRD spectra……………………...……53 4.3.4. Illustration of the clay mineral abundances……………………….…………..…54 4.1.1. BSE images of select samples……………………………………………...……57 4.5.1. XRF major element plots………………………………………..………….……63 4.5.2. XRF trace element plots…………………………………………...….…….……64 4.6.1. LA-ICP-MS trace element plots…………………………………….….…..……69 4.6.2 Diffractograms of dated paleosols……………………………………………….71 5.0. Pedogenic framework………………………………………….………...………76 5.1. BSE image illustrating clay mineral paragenesis………………..…….…...…….82 5.2. Major oxide ratios of select clay minerals………………………….……………83

5.3. K-SiO2/Al2O3 plot illustrating smectitization trends……………………..………84 4.5.1. Rare earth and trace elements concentrations compared to standards………...…95

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CHAPTER 1

INTRODUCTION

A stacked series of Quaternary loess- paleosol couplets, known as the Blackwater

Draw Formation (BDF), forms a relatively thin mantle atop the expansive Southern High

Plains (SHP) in west Texas and eastern New Mexico (Fig 1.1; Reeves, 1976; Holliday,

1989, 1990; Gustavson and Holliday, 1990). The SHP is a 120,000 km2, slightly

southeastern sloping, relatively unmarked plateau that is only interrupted by southernly

dune fields, up to 19,000 tiny playa wetlands with associated lunettes and southeasterly

trending draws (Holliday, 1990). According to Holliday (1990), the origin of the plateau

is associated with the dissolution of Permian evaporites, which ultimately led to the

amalgamation of smaller river drainages to form the larger Pecos, Red and Canadian

Rivers. The development of these major river systems resulted in significant fluvial

incision that isolated the SHP from surrounding landforms at around 1.6 Ma (Fig. 1.1).

The current SHP represents the isolated remnants of the Rocky Mountain piedmont

alluvial plain and the subsequent aggradation of Pleistocene eolian sediments, which

overlie Mesozoic and Paleozoic bedrock (Reeves 1972).

Loess deposits and sediments derived from loess deposits, are the most

widespread terrestrial sediments on earth (Bronger and Heinkele, 1990). Loess is

generated through a variety of mechanisms and can be subject to modification and

reworking by pedological processes (Goudie and Middelton, 2006). Generally, these

sediments are non-stratified, unconsolidated, and comprised of predominately silt, fine sand, clay, and carbonate (Smalley and Vita-Finzi, 1968). During periods of non- deposition or landscape stability, soils have developed at the surface. With regards to the

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BDF, deposition is interpreted to be cyclic in nature, such that during periods of landscape stability, soils formed at the surface of the most recent deposit, which was then covered by the ensuing depositional event (Holliday 1990). Over time, this sequence led to a stacked series of couplets defined by repeating loess deposits, individually capped by a soil. These couplets can form long sequences, which provide sources of paleoenvironmental information, as during pedogenesis, climate conditions will dictate the rate and type of weathering (Rutter et al. 2003).

Figure 1.1. Extent of the Southern High Plains (Llano Estacado) physiographic province in relation to major rivers, U.S. boundaries and adjacent physiographic features. The SHP shapefile was modified from US Geological Survey, 1946.

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Clay minerals are a group of minerals in the phyllosilicate family, specifically hydrous aluminosilicates, which are typically smaller than 2 µm in size (Millot, 1971).

The formation of clay minerals at Earth surface operates through the contact of pre- existing minerals with the atmosphere, hydrosphere, and biosphere. Consequently, clay minerals are ubiquitous in surface or near-surface sediments. Clay minerals form predominately through the hydrolysis of rock-forming aluminosilicates (Chamley, 1989).

As weathering progresses, more stable hydrous phases will eventually form. This process is regulated by the precursor material in conjunction with the modes and intensity of weathering. As such, the formation of authigenic clay minerals in soils is controlled by the initial mineralogical composition and the type and intensity of weathering (Paquet and Clauer, 1997). Furthermore, weathering types and rates are dictated by the prevailing climate conditions.

The subject research area is located adjacent to the Bushland Playa, located roughly 24 km west of Amarillo, Texas (Fig 1.2). This location was chosen because the locality represents the northernmost extent of the BDF, defined by the caprock escarpment in this location (Fig 1.2). While previous research has explored the paleoclimate implications of sediments in the BDF at the Lubbock type section, few similar investigations have been conducted on the BDF at its northernmost extent (Stine et al., 2020). Furthermore, recent studies have discussed the interplay of sediments sourced from the Pecos River Valley and from northern mid-continental loesses within the BDF, and while the BDF at the Lubbock type section is composed of two members or sections, little is known about the interaction of these members in more northern sources

(Stine et al., 2020; Tewell, 2020). It is expected that at northern localities of the BDF, the

Texas Tech University, Colton Mallett, May 2021 upper member should be relatively thickest, therefore, if there were a provenance shift from a southern source to a northern source, this locality would record it clearly (Stine et al., 2020).

The Bushland Playa Core is 13.9 m in length and likely represents the entirety of the BDF, as the base of the core is interpreted to be the top of the Caprock Caliche, which directly underlies the BDF. The “Caprock” is a calcrete that forms the distinct, well indurated, uppermost strata of the Ogallala Formation (Reeves, 1970). Aside from the surface soil horizon, the formation at this locality is comprised of accumulation horizons containing varying amounts of pedogenic carbonate and loess. The soils from this locality are generally a reddish brown, the grain size varies from fine sand to silt with varying abundances of clay and contains varying amounts of pedogenic features such as rhizoliths as well as carbonate and ferromanganese nodules.

The aim of this study is to utilize clay mineral speciation and its mineralogical and geochemical variability to build a pedogenic framework and extrapolate prevailing paleoclimatic conditions during the Quaternary period. Additionally, K-Ar dating of authigenic and detrital illite in select samples will be utilized to ascertain if any change in provenance occurred during the deposition of the BDF in this locality. This study furthers the understanding of changing climate conditions across the Mid-Pleistocene Transition and works to establish a correlation between provenance change and changing climate conditions.

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Figure 1.2. The location of the Bushland Playa Core sample location in relation to the cities of Bushland and Amarillo, the Caprock Escarpment, and the Canadian River. Esri. “ESRI Satellite” [basemap]. Scale Not Given. “World Imagery”. November 16, 2016. https://server.arcgisonline.com/ArcGIS/rest/services/World_Imagery (August 22, 2020). Made with QGIS.

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CHAPTER 2

BACKGROUND 2.1 Current Landscape and Climate of the Southern High Plains

2.1.1 Modern Geomorphology

The Southern High Plains (SHP), or Llano Estacado (Staked Plains), are defined as the approximately 120,000 km2 elevated plateau, east of the Pecos River, south of the

Canadian River, west of the upper Red, Brazos and Colorado rivers and north of the

Trans Pecos and Edwards Plateau (Fig 2.1; Reeves, 1976).

A well-indurated calcrete in the upper Ogallala Formation, referred to as the

“Caprock”, delineates the western, eastern, and northern extent of the SHP. Along these

extents, the calcrete forms the Caprock Escarpment, providing a clear boundary to the

relatively elevated SHP. The Caprock Escarpment, in the east, forms a sharp cliff face,

separating the SHP from the Permian and Triassic “red beds” of the West Texas Rolling

Plains. Similarly, in the north and west, the Caprock Escarpment separates the SHP from

the Canadian River Basin and the Pecos River Valley, respectively. In contrast to the

well-defined Caprock Escarpment, the southern extent of the SHP is notably less distinct

(Fig 2.2). Here, the SHP and the defining “Caprock” caliche slowly grades into the top of

the Edwards Limestone of the Edwards Plateau (Reeves, 1976).

Locally, the SHP appears to be flat, but in fact has a very shallow southeastward

slope of about 1.9 to 3.75 meters per kilometer (Fig 2.1.2). The landscape is marked by

up to 19,000 shallow ephemeral lakes, or playas, that often have associated lunettes.

Approximately 30 larger depressions form saline lakes, or salinas, which are dispersed

across the SHP. The surface of the SHP is additionally cleft by several NW-SE,

Texas Tech University, Colton Mallett, May 2021 subparallel draws, that form the tributaries of the Colorado, Brazos, and Red River. These same tributaries form vast re-entrant canyons on the eastern edge of the SHP, such as Mt.

Blanco and Tule Canyon (Fig 2.1.1; Fig 2.1.2). The western and southwestern extent of the SHP is dominated by extensive Holocene dune fields (Reeves 1972, Holliday, 1990).

Figure 2.1.1. Geographic extent of SHP in relation to major cities, landforms and localities referenced in the text.

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A C

Figure 2.1.2. (A) Elevation profile of the SHP from the Colorado River, due south, to the Pecos River, passing through the study area. (B) Elevation profile of the SHP from the Pecos River, due east, to the West Texas Rolling Plains, passing through Lubbock, Texas. The red vertical lines illustrate the extent of the SHP and landmarks discussed in the text are labeled. (C) Map illustrating the profile lines A-A` and B-B.

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2.1.2 Modern Climate

The modern climate on the SHP is generally considered semi-arid, and there exists a slight trend of increasing precipitation from the west to the east (Carr, 1967). According to the National Centers for Environmental Information, National Oceanic and

Atmospheric Administration (NOAA) Climate Divisional database (nCLIMDIV), the mean annual precipitation ranges from approximately 380 to 635 mm along this gradient, although there is considerable variation in precipitation from month to month. Overall, the average precipitation for the whole of the SHP is 496.57 mm per year. Additionally, the mean temperature for the region varies between 5 to 15.6 ºC (41-60 ºF) at a minimum, to 21.7 to 26.7 ºC (71-80 ºF) at a maximum, and overall averages 14.61 ºC (58.3 ºF)

(NOAA).

The Köppen-Geiger climate classification is a system widely used to classify regional climate. This classification system divides climates into five main categories based upon mean annual precipitation volume (MAP), seasonal trends, and mean annual temperature (MAT). Furthermore, this classification system is best used to represent long term mean climate (Köppen, 1936).

An updated version of the Köppen-Geiger climate classification by Peel et al. (2007) classifies the climate of the SHP as predominately cold arid-steppe (BSk), with minor hot

arid-steppe (BSh) and cold arid-desert in the south (BWk) (Fig 2.1.3). An arid-steppe

climate is contemporarily referred to as a semi-arid climate and an arid-desert is referred

to as a desert climate.

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Figure 2.1.3. Köppen-Geiger climate classification of the U.S illustrating that the climate on the Southern High Plains is predominately cold arid-steppe (BSk), with minor hot arid-steppe (BSh) and cold arid-desert (BWk). Modified from Peel et al. (2007).

2.2 Geomorphic Evolution and Pertinent Strata of the Southern High Plains

2.2.1 Development of the Southern High Plains (SHP)

Following the progressive isolation of the SHP, which concluded at approximately

1.6 Ma, sedimentation in this region transitioned from largely fluvial, to predominately

eolian (Gustavson and Holliday, 1999). This shift in deposition is attributed to the

dissolution of underlying Permian evaporites. As the dissolution of Permian evaporites

proceeded, the region underwent a period of significant subsidence, inducing the

coalescences of minor drainages into the Pecos and Canadian rivers. The Pecos and

Canadian rivers incised into the High Plains, and the SHP became progressively isolated

from the surrounding landforms, resulting in the reduction of fluvial deposition and the

relative increase in eolian deposition. The Miocene- Pliocene Ogallala Formation records the fundamental change in depositional mechanisms, and is concluded by the overlying, purely eolian, Quaternary deposits (Fig. 2.2.1; Reeves 1972; Holliday, 1990). These

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Quaternary eolian sediments comprise the Blackwater Draw Formation (BDF) and

blanket the extent of the isolated SHP (Fig 2.2.1). While the strata underlying the BDF formation vary across the SHP, the Ogallala Formation composes a vast majority of it

(Fig 2.2.1). Locally, the lacustrine Tule and Blanco formations form isolated, unconformable deposits between the Ogallala Formation and the BDF (Holliday, 1988a).

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Figure 2.2.1. Simplified stratigraphy of the Southern High Plains. Modified from Holliday, 1990.

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2.2.2 Ogallala Formation

The Ogallala Formation is a vast alluvium and eolian deposit, which extends over

1285 km, or 450,000 km2 from southeastern New Mexico to southern South Dakota and underlies the entirety of the SHP (Darton, 1905; Reeves, 1972). The original eastern extent of the Ogallala Formation is ambiguous, but gravels from this formation have been discovered in Dallas County, Texas, indicating that these deposits once extended much further than the current extent of the SHP (Menzer and Slaughter, 1970; Reeves, 1972).

These Miocene- Pliocene deposits are predominately composed of alluvium derived from the westward lying Rocky Mountains and related eolian sediments (Winkler, 1985;

Gustavson and Winkler, 1988).

The Ogallala Formation was deposited in incised bedrock valleys following the post-Cretaceous erosional event, such that the Ogallala Formation lies unconformably over Permian through Cretaceous strata (Frye et al., 1956; Reeves, 1972). The relief of the underlying surface was established during the post-Cretaceous erosional event and was likely further modified by local dissolution of the underlying Permian salts (Reeves,

1972; Gustavson and Winkler, 1988). The thickness of the Ogallala Formation is controlled by this underlying paleotopography and varies from less than 1 m to 150 m, but the average thickness on the SHP is 30 m (Gustavson and Holliday, 1999). The

Ogallala Formation thins and become discontinuous as the SHP begins to merge with the

Edwards Plateau at the southern margin of the SHP (Frye and Leonard, 1957).

The Ogallala Formation can be disaggregated into a lower section and an upper section. The lower section corresponds to the sediments deposited prior to the incision of the Pecos River, and are predominately fluvial with minor eolian sediments (Gustavson

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and Holliday, 1999). As minor drainages coalesced into larger rivers, namely the Pecos

and Canadian rivers, fluvial sedimentation was reduced, allowing for eolian deposition to

become dominant, resulting in the deposition of the upper section (Reeves, 1972).

The depositional history of the Ogallala Formation, specifically prior to the

existence of major river drainages in the region, is particularly complex. As rivers cut across the landscape, repeatedly incising and filling drainage valleys, sedimentation was predominantly fluvial, although eolian deposition occurred in uplands and drainages.

Vertically stacked, sandy and gravelly fluvial facies occur within extensive, southeastward trending paleovalleys that incise into upper Paleozoic and Mesozoic strata

(Reeves, 1972; Hawley, 1984; Winkler, 1985; Gustavson and Winkler, 1988). The fluvial sediments in the lower section are often interbedded with eolian sediments. These eolian sediments were deposited atop the paleouplands and paleodivides between paleovalleys

(Winkler, 1985; Gustavson and Winkler, 1988).

The upper section of the Ogallala Formation is composed of eolian silt and loamy sands. These sediments were likely derived from the floodplains of rivers whose sediments compose the lower fluvial section (Winkler, 1985; Gustavson and Winkler,

1988; Gustavson and Holliday, 1999). The upper section of the Ogallala Formation contains a well-preserved series of stacked paleosols-eolian couplets, similar in nature to the BDF. These paleosols-eolian couplets, like the BDF, record alternating and repeating periods of sediment accumulation and then landscape stability and pedogenesis (Reeves,

1984; Gustavson and Winkler, 1988; Gustavson and Holliday, 1999). The cessation of deposition of the Ogallala Formation is defined by the distinct “Caprock Caliche”, a 3 to

4 m thick, highly indurated calcrete (Reeves, 1970). This style of cyclic deposition and

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subsequent pedogenesis of eolian sediments has occurred on the SHP since the Pliocene,

initially with the upper Ogallala Formation, and now with the BDF (Holliday, 1990;

Gustavson and Holliday, 1999).

2.2.3 Lacustrine Strata

The Blanco Formation underlies the BDF along the eastern margin of the SHP

(Frye and Leonard, 1957). The Pliocene Blanco Formation is composed of lacustrine and alluvial sediments, which were deposited in individual paleobasins inset in the underlying

Ogallala Formation (Dolliver and Holliday, 1988; Holliday, 1988b). The paleobasins are likely remnants of stream channels, or basins formed by solution- induced subsidence

(Gustavson and Finley, 1985; Gustavson, 1986). The Blanco Formation was deposited prior to the BDF, such that the Blanco Formation is situated unconformably between the underlying Ogallala Formation and overlying BDF. The Blanco Formation is generally

14 m thick, but can reach thicknesses of 27 m (Pierce, 1973). Mt. Blanco, the type section for this formation is in Blanco Canyon, a reentrant to the SHP formed by the White River

(Fig. 2.1.1; Dolliver and Holliday, 1988). Dating of glass shards from the Blanco Ash, which is near the center of the Blanco Formation at Mt. Blanco, gives an age of 2.8 ± 0.3

Ma (Boellstorff, 1976). The Mt. Blanco section also contains the type fauna for the

Blancan Land Mammal Age of North America (Schultz, 1977). According to Holliday

(1988), the minimum age for the end of deposition of the Blanco Formation is 1.6 Ma, based on the date of the Guaje Ash and inferences regarding the time required for the formation of conformably underlying calcisols. At the Mt. Blanco locality, there exists a

Stage IV calcic horizon within a calcrete at the top of the Blanco Formation, and is likely related to the formation of the overlying BDF.

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The lower Pleistocene Tule Formation is a lacustrine formation similar in character to the Blanco Formation. The sediments of the Tule Formation fill abandoned valleys incised in the BDF and the Miocene-Pliocene Ogallala Formation and the Triassic

Dockum Group. Characteristic sediments of the Tule Formation include thin limestone and dolostone beds and laminated, horizontally bedded mudstones. The Lava Creek B

Ash near the top of the Tule Formation, the Cerro Toledo-X ash near the base of the Tule

Formation and vertebrate remains suggest that the Tule Formation is early or middle

Pleistocene in age (Izett, 1977; Izett and Wilcox, 1982). The type section for this formation is in Tule Canyon (Fig. 2.1).

2.3 Blackwater Draw Formation

2.3.1 Physical Characteristics of the Blackwater Draw Formation

The BDF is characterized by a series of stacked strata, interpreted as loess-paleosol couplets, that form a thin mantle over the extent of the ~120,0002 km SHP (Holliday,

1989; Gustavson and Holliday, 1999). These couplets occur as distinct packages of sediment, each capped by a paleosol. The number of individual paleosols within the BDF varies between 4 and 14 at distinct localities. The thickness of the BDF also varies significantly, averaging between less than 1 m, and up to 9 m, but can locally be up to 27 m thick (Reeves 1976; Hovorka, 1997). The range in thickness seemingly correlates with the number of paleosols present, with sections of greater thickness having a greater number of preserved paleosols (Hovorka, 1997). The variation in the number of individual paleosols at different locations is likely due to wind deflation and soil welding processes (Holliday, 1989). Furthermore, different paleosols display varying degrees of pedogenesis.

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The character of the BDF varies across the SHP. The sediments of the BDF are generally red to reddish brown sands and silts. These sediments exhibit a fining sequence, such that sediments in the southwest are fine to very fine sand and the sediments in the northeast are very fine sand to silt and clay (Fig 2.3.1; Seitlheko, 1975;

Holliday, 1989). The thickness also trends such that deposits are thinner in the west and southwest of the SHP, but are thicker in the northeast (Reeves, 1976; Hovorka, 1997).

The number of individual layers and paleosols is fewest in the west and southwest and greatest in the northeast. The sediments in the BDF are generally 80-90% quartz, with smaller percentages of feldspar, mica, and clay minerals, and while these averages do not generally do not vary across the SHP, they do vary withing the vertical profile (Stine et al., 2020).

Project

Figure 2.3.1. Sediment classification within the SHP. Adapted from Seitlheko,1975.

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2.3.2 Depositional Models

The preliminary depositional model suggested that the Pleistocene sediments of the

SHP, currently recognized as the BDF, were split into a distinct northeastern formation

and a distinct southwestern formation. The northeastern sediments are predominantly

eolian loesses, in contrast to the fine sands in the southwestern extent of the SHP. The

fine sands were interpreted to be sourced from the alluvium in the nearby Pecos River

Valley (Frye and Leonard, 1965; Reeves, 1971, 1976). These sediments were initially

called “cover sands” by Frye and Leonard (1957) but were later renamed to the current

BDF by Reeves (1976). These models did not account for the pervasive soils and did not

extensively explore the interplay between the two apparent sediment provenances.

A more evolved depositional model accounts for the pervasive, well developed soils,

defined by a repeating series of argillic horizons topped by calcic horizons, and proposed

that the deposition of the BDF was cyclic in nature (Holliday, 1989; Gustavson and

Holliday, 1989). Each cycle includes a period of eolian deposition followed by a period

of landscape stability and pedogenesis. In this model, sediments were sourced from the

Pecos River Valley and carried onto the SHP during depositional periods, creating a

continuous sheet of sediments across the SHP. As deposition slowed or ceased,

pedogenesis predominated. The evolving soils were likely affected by wind deflation,

which further augmented the variability in the number of soils and thickness of soils at

different locations across the SHP. This cycle of deposition and pedogenesis resulted in

the series of soils, with younger soils disconformably overlying previous soils. The process began at more than 1.6 Ma ago and continues across the SHP today (Holliday,

1989).

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The most recent depositional model (Stine et al., 2020) builds upon the two previous

models but additionally proposes that a change in the frequency of astronomical climatic forcing during the mid- Pleistocene transition (MPT) was the driving factor in the evident depositional cyclicity and variation in weathering, texture and possible sediment sources across the SHP (Stine et al., 2020). This model proposes that the BDF can be partitioned into sediments deposited pre-, syn- and post- the MPT. Prior to the MPT, sediments were sourced from the Pecos River Valley and deposited as sand sheets onto the SHP during periods of deposition, after which periods of landscape stability allowed for widespread pedogenesis. While this was occurring, the climate was oscillating through numerous, short lived, 41 ka climate cycles that resulted in a relatively arid climate. As the MPT progressed and eventually terminated, 100 ka strongly asymmetric climate cycles were characterized by long periods of cooling and glaciation, followed by a shorter interglacial period. This transition is interpreted to be the cause of more humid conditions during deposition and to be responsible for the introduction of silts and finer sediments into the upper section of the BDF. The longer and more intense glacial periods allowed for the production and influx of finer sediments onto the SHP. This resulted in syn- and post-

MPT BDF sediments to be a combination of northern derived loesses and southern derived alluvium and reworked sands sourced from the Pecos River Valley (Stine et al.,

2020).

2.3.3 Age Estimates

Currently, the ambiguous age constraints and age models, combined with the extensive variability in the thickness and number of soils in the BDF, has led to numerous uncertainties in regard to the minimum and maximum age of deposition of the BDF, as

Texas Tech University, Colton Mallett, May 2021 well as uncertainties related to the depositional ages of each individual paleosol. The age of the BDF is constrained by fission track dates on glass shards from the few in-situ ashes and from thermoluminescence and paleomagnetic dates and the inference of soil formation from well-developed calcisols (Izett et al., 1972; Lindsay et al., 1975; Patterson et al., 1988; Holliday, 1988a; Holliday, 1989; Patterson and Larson, 1989; Gustavson and

Holliday, 1999; Stine et al., 2020).

Two ash beds, which occur locally in the BDF, supply relative age constraints. The maximum age of the BDF is constrained by the 1.4 Ma Guaje ash, which occurs locally near the base of the BDF sequence (Izett et al., 1972). At the Mt. Blanco section, there exists a well-developed calcrete below the Guaje ash, which would have required several hundred thousand years to form, giving a minimum age of 1.6 Ma of the BDF at this location (Holliday, 1988a). Additional age constrain is provided by the Lava Creek B volcanic ash which was dated to be 0.6 Ma. The Lava Creek ash occurs within the BDF, approximately 90 cm above the basal contact of the sequence (Izett and Wilcox, 1982).

This 90 cm of BDF contains a well-developed soil with a Bt horizon (Izett and Wilcox,

1982).

In addition to fission track dates of glass shards, age control is supplied by thermoluminescence and paleomagnetic reversal data from BDF sediments. Thus, thermoluminescence ages were delivered by samples of the upper two Bt horizons at the

BDF type section. The uppermost soils were dated at 118 ± 14 ka and the second buried soil was dated at greater than 270 ka (Holliday, 1989). These data are consistent with magnetostratigraphic studies conducted by Patterson and Larson (1989) which found that the most recent and uppermost buried soils are normally magnetized, suggesting that

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these soils were formed within the last 730 Ka (Patterson et al., 1988; Patterson and

Larson, 1989). The third soil is magnetically reversed, which suggests the soil was

deposited in a period of reversed polarity, likely 0.9 Ma to 0.73 Ma (Singer, 2014). The

soils lower in the section exhibit both normal and reversed components, which only

allows the determination that they are older than 0.9 Ma (Patterson et al., 1988; Patterson

and Larson, 1989).

Finally, the underlying Ogallala Formation and locally underlying Blanco Formation serve as the ultimate maximum age control for the BDF. The Blanco Formation contains the Blanco Ash which has returned glass-shard fission track ages of 2.32 ± 0.15 Ma (Izett et at., 1972; Izett, 1981). Additional paleomagnetic studies indicate the Blanco Formation was deposited between 2.4 Ma and 1.4 Ma (Lindsay et al., 1975). Age models for the upper Ogallala Formation posit that deposition of this formation did not occur after approximately 2 Ma. These data indicate that the BDF is not likely older than 2 Ma

(Gustavson and Holliday, 1999).

2.3.4 Recent Observations

Recent work of Stine et al. (2020) and Tewell (2020) have shed new light on the character of the BDF at the Lubbock type section and at the Bushland Core, respectively.

Stine et al. (2020) used additional textural, geochemical, and rock-magnetic data to revise the pedogenic framework of the Blackwater Draw Formation type section initially proposed by Patterson and Larson (1989) and Holliday (1989). This new pedogenic framework describes a surface soil and 5 underlying paleosols. Paleosols 4 and 5 are interpreted to have been deposited prior to the mid-Pleistocene transition (MPT), while the surface soil and paleosols 1 through 3 were deposited during and after the MPT.

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Furthermore, this study posits that paleosols 4 and 5 are composed of predominantly sands transported from the Pecos River Valley, while the younger paleosols and surface soil are a mixed-source provenance from both the Pecos River Valley and midcontinent loess. Pre-MPT sediments were deposited during shorter duration, less humid climate cycles, while syn- and post- MPT sediment were deposited during longer duration, more humid climate cycles.

The Bushland Core was previously studied by Tewell (2020). This study concluded that the Bushland Core records a surface soil and four paleosols. The core can be divided into an upper member and a lower member based on grain size, geochemistry, and weathering indices. The lower member contains paleosol 4, which has a geochemical signature similar to sediments from the Pecos River Valley, and likely was deposited under relatively more arid, hence less weathering conditions. The uppermost 3 paleosols and the surface soil have a geochemical signature similar to Nebraskan loesses.

Furthermore, the upper member is thicker in this locality, compared to the Lubbock type section, indicating less sediment input from the Pecos River Valley and more input from northern sources. These findings are consistent with the work of Stine et al. (2020).

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Figure 2.3.2. Stratigraphic columns of the Lubbock type section described by Stine et al. 2020 and the Bushland Playa Core defined by Tewell (2020). The sections are proportionately scaled and the Caprock Caliche was used as a datum for correlation.

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CHAPTER 3

MATERIALS AND METHODS 3.1 Core Procurement

Three 5.08 cm (2 inch) diameter cores were acquired using a hydraulic push corer

and plastic soil sampling sleeves approximately 0.7 m adjacent from each other, near a

playa in Bushland, Texas (Fig. 3.1). Each core records approximately 14 m of unconsolidated sediment. Coring was ceased at approximately 14 m below the grounds surface, due to the inability of the coring rig to penetrate through a dense caliche horizon.

This impenetrable calcite-rich horizon was interpreted to be the “caprock caliche” of the

Miocene to Pliocene Ogallala Formation. It is therefore assumed that the entirety of the

Blackwater Draw Formation in this locality is represented within the core. One core is

housed at Texas Tech University, while the other two cores are housed at the University

of Texas at Dallas.

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Figure 3.1. The location of the sample location in relation to Bushland, Texas and the Playa in Bushland.

3.2 Core Preparation

The core was preserved continuously, as it was being collected, in 12 individual,

1.2 m-long plastic sleeves labeled A-L. Following the completion of coring, the preserved core was transported to TTU. There, the sealed plastic sleeves were placed four at a time, in sequence, on a table, then cut open with a utility knife so that the core could be laid out on the table aside the scale. The core was then bisected with a utility knife and splayed so that the internal structure and lithology could be described. The core was

visually inspected and described in terms of grainsize, pedogenic features and color. The

Munsell Soil Color Chart was used to assess the color of the sediment. The appearance,

or lack thereof, of pedogenic features were used to differentiate between individual

accumulation, or “B”, soil horizons. These observations were recorded with high

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resolution photos and a core log. Following the completion of the core log and collection

of photos, 15-30 cm increments of the core were placed in sample bags labeled with their

corresponding plastic sleeve name and the depth in centimeters below the grounds

surface. The core sample bags were placed in storage. The core log is presented in

Appendix F.

3.3 Sampling

Using the preliminary visual descriptions and interpretation of soils and soil

horizons, a total of 18 individual sample locations were chosen explicitly, so that the

global trend of the core was represented. Following the completion of high-resolution

geochemistry analysis and the reclassification of individual soils by Tewell (2020) an

additional 4 samples, BL-19 through BL-22 were prepared. The samples were extracted from the previously prepared sample bags and labeled BL-1 through BL-22 (Fig. 3.2).

For the purposes of this research, and because the soil horizons are seemingly repetitive, it was not necessary to gather high resolution data, but rather have a data points from each soil and from each soil horizon type.

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BL-1 BL-2

BL-3 BL-4

BL-5 BL-6

BL-7

BL-8

BL-8b

BL-9

BL-9b

BL-10 BL-10b

BL-11

BL-12

BL-13

BL-14

BL-15 BL-15b

BL-16

BL-17

BL-18

Figure 3.2. Location in the core for each of the 22 samples chosen for analysis.

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3.4 Optical Microscopy

Sample locations BL-1 through BL-18 were chosen for thin section analysis.

Pieces of sediment with an average weight of 15 g were sent to Petrolab (Sardegna, Italy) to be prepared as thin sections. Thin sections were observed using a Leica DM750 P under plane polarized light (PPL) and cross polarized light (XPL) at magnifications between 4x and 20x. Samples were assessed for types of pedogenic features, for the intensity and type of pedogenesis and mineralogy. Thin section petrography results are given in appendix B.

3.5 Loss on Ignition

Loss on ignition was conducted in the Texas Tech GeoAnalytical Laboratory for a

series of samples from BL-1 through BL-18. Porcelain crucibles were heated in a muffle furnace at 1000 °C for 30 minutes. After heating, the crucibles were placed in a desiccator to cool for 15 minutes. Once cool, the crucibles were weighed, then approximately 3-4 g of previously milled sample were measured out into the crucible, then the crucible was placed into the muffle furnace for an additional 30 minutes at 1000

°C. Following heating of the samples, the crucibles were placed in the desiccator and allowed to cool for 20 minutes. Once cooled, the weight of the crucibles and samples were re-weighed to calculate the LOI using the following equation.

( ) ( ) 100 LOI= Crucible+ Wet Sample − Crucible+ Dry Sample Crucible + Wet Sample ∗ The results of the LOI are recorded in appendix E.

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3.6 Glass Disc Fusions

Glass disks were prepared by borate fusion with 1:5 sample-to-flux ratio per

Texas Tech GeoAnalytical Lab standard operating procedures. Borate flux was represented by a mixture of Li-tetraborate (Li2B4O7), Li-metaborate (LiBO4) and

anhydrous Na tetraborate (Na2B4O7). Each sample was prepared by mixing of 2 g of flux

and 2 g of sample in the porcelain crucible. Then an additional 2 g of flux was added to

the sample. The sample and flux were homogenized and then poured on the flux in the

platinum crucible. Two samples at a time were attached to the M4 Claisse Fluxer and

fused into glass disks, placed in plastic storage containers, labeled, and sent to the TTU

GeoAnalytical lab for XRF analysis.

3.7 X-Ray Fluorescence

Global fraction samples, prepared as glass disc fusions, of BL-1 through BL-18

were analyzed for SiO2, TiO2, Al2O3, Fe2O3, MnO, MgO, CaO, Na2O, K2O, P2O5, Ba,

Co, Cr, Cu, Ni, Sc, Sr, V, Zn and Zr using the Thermo ScientificTM ARLTM PERFORM’X

Sequential X-Ray Fluorescence Spectrometer housed at the Texas Tech GeoAnalytical

Laboratory. The data was collected using 60 kV, under vacuum, and processed using the

UniQuant software. Major element concentrations were measured as oxides in weight

percent, while the trace element concentrations were measured in parts per million. The

results of these analysis are provided in appendix E.

The Thermo ScientificTM ARLTM PERFORM’X Sequential X-Ray Fluorescence

Spectrometer was calibrated using the in-house standards AGV-120, GSP-1 and BHVO-1

(Appendix D). The three standards were analyzed twice, once before BL-1 through BL-

18 were analyzed, and once after. The three standards were analyzed under the same

Texas Tech University, Colton Mallett, May 2021 conditions and for the same elements as the BL samples. The measurements of the in- house standards were used as a standard of comparison, to calculate and confirm the accuracy and precision of the machine.

3.8 Laser Ablation Inductively Coupled Plasma Mass Spectrometry

Following XRF analysis, the glass disc fusions BL-1 through BL-18 were analyzed for trace elements. The previously prepared glass discs were trimmed on a

Beuhler trim saw outfitted with a diamond blade wheel and prepared as thin sections by the Texas Tech University Thin Section Lab. The thin sections were polished using

Beuhler Beta Polisher with an automated vector head using 9 and 3 µm diamond polishing pastes per lab standard operating procedures.

The polished thin sections of BL-1 through BL-18 were analyzed for Mg, Sc, Ti,

V, Ga, Rb, Sr, Y, Zr, Nb, Ba, La, Ce, Pr, Nd, Sm, Eu, Gd, Tb, Dy, Er, Yb, Lu, Hf, Ta, Pb,

Th and U at the Texas Tech GeoAnalytical Lab using an Agilent 7500cs inductively coupled plasma quadrupole mass spectrometer and a New Wave UP123 Nd: YAG laser ablation system. Each sample was analyzed three times.

The mass spectrometer was calibrated using the glass value of Calcium 43 from the in-house standard BHVO. The in-house standards BHVO and GSD were used as a standard of comparison, to calculate and confirm the accuracy and precision of the machine. Prior to, and following the analysis of the selected samples, BHVO and GSD were analyzed. GSD was further analyzed every two samples (6 laser spots), and BHVO was analyzed every 6 samples (18 laser spots).

The raw data was reduced using the GSD and BHVO in-house standard through excel per lab standard operating procedures. Following the reduction of the trace element

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concentration data, the three analysis for each sample were averaged to give a single

concentration value for each analyte (appendix F).

3.9 X-Ray Powder Diffraction

X-ray powder diffraction was performed on the global sample and clay fraction

for samples BL-1 through BL-18. Sample preparation initially included material

powdering in an agate mortar prior to whole rock measurements. Clay fraction was

separated according to Zanoni et al. (2016), which includes sonication and centrifugation.

The oriented mounts were prepared by pipetting the suspension on glass slides and left to

dry at room temperature. The thickness of the prepared mounts exceeded 50 μm, which is

required for semi-quantitative determination of the clay mineral content (e.g. ‘infinite

thickness’ of Moore and Reynolds, 1997). The measurements were undertaken in air-

dried (AD) conditions and after ethylene-glycol (EG) saturation as well as heated to 450

°C at the Geosciences Clay Laboratory of Texas Tech University using a Bruker D8

Advance diffractometer. The instrument is featured by the horizontal goniometer axis and synchronized rotation of both the X-ray source and detector arms. Measurement parameters consisted of a step scan in the Bragg-Brentano geometry with CuKα radiation

(45 kV and 40 mA). Sample mounts were scanned at a counting time of 1.8 s per

0.02°2Θ from 3 to 70 and from 3 to 30°2Θ for the whole rock and the clay fraction,

respectively.

3.10 Interpretation and modeling of XRD patterns

The interpretation of XRD diffractograms was done using Bruker EVA software

and compared against the PDF4+ database issued by the International Centre for

Diffraction Data. X-ray diffraction patterns of clay minerals were examined using

Texas Tech University, Colton Mallett, May 2021 methods described by Moore and Reynolds (1997) and Środoń (2006). For chlorite- mineral interpretation, the recommendations of Lagaly et al. (2006) were followed. The values of Kübler index (KI: Kubler , 1964; Ferreiro Mähmann et al., 2012 ) and defined by the width of the illite 001 peak measured at half of the peak height above the background (Eberl and Velde, 1989), were measured using the same software package.

Characterizing mixed layering only by analyzing the 001/002 and 002/003 I-S peak positions was practically impossible given the fact that the samples analyzed consisted of multiple I-S mixed-layer minerals both rich and poor in illite (Moore and Reynolds,

1997). To characterize the mixed-layer clay minerals, type of chlorite, and 10Å species

XRD clay patterns were modeled using Sybilla™ (Drits and Sakharov, 1976) software package. The modeling included a trial-and-error procedure that provided optimal clay- mineral structural and probability parameters to obtain the best fit between experimental and calculated patterns and of the intensities of the 00l reflections for each of the clay phases present. For mixed-layer minerals, the number, nature, and stacking sequence of different compositional layers were taken as modifiable values (e.g. Uzarowicz et al.,

2011; Šegvić et al., 2020). The whole rock and clay fraction spectra, the modeled clay fraction spectra and the quantification of whole rock and clay fraction spectra are included in appendix C.

3.11 Scanning Electron Microscopy – Energy Dispersive X-Ray Spectroscopy

The collection of back scatter electron (BSE) images and energy dispersive X-ray

(EDS) spectra was acquired using a Zeiss Crossbeam 540. The samples BL-7, BL-11,

BL-13 and BL-14 were chosen because their clay fraction diffraction spectra represented the 3 groups established from the modeling of XRD patterns. Samples BL-13 and BL-14

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were analyzed at 4 sites using the EDS, while BL-7 was analyzed at 3 sites and BL-11 was analyzed at two sites. A total of 260 EDS spectra were collected from the four samples. A BSE image, with the location of the gathered spectra denoted, were collected for each site.

The BSE images were collected with an EHT of 15.00 kV at a working distance between 4.4 mm and 8.3 mm. EDS analyses were performed on spots with a diameter of

1 µm, with acquisition live time of 20 s and a general working distance around 8 mm.

Chemical data were used as atomic percentages and were normalized to 100% (appendix

D).

3.12 K-Ar Dating of Illite

Samples BL-2, BL-10 and BL-17 were radiometrically dated using potassium and

argon at the Laboratorio Nacional de Geoquímica y Mineralogía; Instituto de Geología,

Universidad Nacional Autónoma de México. The select samples were centrifuged and

accumulated into <1 µm and 1-2 µm size fractions using the methods described in Zanoni

et al. (2016). The 1-2 µm fraction is expected to contain predominately detrital illite,

while the <1 µm fraction is expected to contain predominately authigenic illite.

The K-Ar method used in this study is described in detail by Solé and Enrique

(2001) and Solé (2009). The potassium was measured by X-ray fluorescence using a

specific regression for measuring K (Solé and Enrique, 2001). Argon was extracted by

complete sample fusion using a 50 W CO2 laser defocused to 1-3 mm diameter. The

measurements were performed in a static vacuum with a MM-1200-B noble gas mass

spectrometer, using an electromagnet with peak switching controlled by a Hall probe.

The total uncertainties were better than 0.2%, and 0.5% for 40Ar-38Ar and 36Ar

Texas Tech University, Colton Mallett, May 2021 respectively. The K–Ar ages were calculated using the decay constants recommended by

Steiger and Jäger (1977). Age errors were obtained at 1σ level.

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CHAPTER 4

RESULTS

A summary regarding the location of samples within the core, and the analyses conducted on said samples, are provided in Table 4.0.1. Initially, an original 18 samples were collected, labeled BL-1 through BL-18, then Tewell (2020) provided supplemental evidence of distinct soils within the core. To properly investigate modal and clay fraction mineralogy variation between these soil profiles, and additional 4 samples were collected, labeled BL-8b, BL-9b, BL-10b and BL-15b. The BL-15b XRD CF sample file was corrupted and was unfortunately not included in the XRD CF sample suite.

The results of the analyses performed in this study will be displayed in this chapter, while interpretation of soil profiles will be explored in the discussion.

Table 4.0.1. Table illustrating the location of each sample in the core and what analyses were conducted for each sample. Optical Depth XRD LA-ICP- K-Ar Sample Petrograph XRD WR SEM XRF (m) CF MS Dating y BL-1 0.3 X X X X X BL-2 0.7 X X X X X X BL-3 1.45 X X X X X BL-4 1.7 X X X X X BL-5 2.4 X X X X X BL-6 2.6 X X X X X BL-7 3.5 X X X X X X BL-8 4.2 X X X X X BL-8b 5.1 X X BL-9 5.9 X X X X X BL-9b 6.5 X X BL-10 7.3 X X X X X X BL-10b 7.75 X X BL-11 8.35 X X X X X X BL-12 8.9 X X X X X BL-13 9.55 X X X X X X BL-14 10.2 X X X X X X BL-15 11.1 X X X X X BL-15b 11.6 X BL-16 12.4 X X X X X BL-17 13.1 X X X X X X BL-18 13.8 X X X X X Abbreviations: CF: Clay fraction <2µm, WR: Whole rock fraction. The samples chosen for analysis using X-ray diffraction were each analyzed three times; once prepared as an air-dried sample, once after glycolation and once after heating.

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4.1 Optical Petrography

The thin section samples were prepared primarily for analysis via SEM-EDS, which required the disaggregation of the sample sediments prior to thin section preparation, therefore the captured photomicrographs do not represent the original in-situ

textural characteristics of the samples. Additionally, due to the lab’s preparatory

techniques, the thickness of the grains within the sample is inconsistent, resulting in a

high degree of variation of pleochroism between similar minerals across the thin section.

Accordingly, any attempt at point count quantification of modal mineralogy would be

less effective than XRD quantification. Optical petrography was therefore used to

identify prevalent minerals, their morphology, cements, and some accessory phases.

The primary phases present in the samples are quartz, plagioclase, orthoclase, and

microcline (Fig. 4.1.1). Other identified phases include biotite, muscovite, kaolinite, and

various oxides (Fig. 4.1.1). The greatest variant observable through petrographic

techniques is the quantity and texture of clay minerals and pedogenic calcite. In addition

to the observed variation of clay and pedogenic carbonate quantity, the grainsize of the

major phases in BL-18 appear to be slightly larger than samples shallower in the

Bushland Playa Core (Fig. 4.1.1).

Quartz grains occur as singular, subequant, anhedral to subhedral crystals that

generally display slight undulose extinction, although several instances of larger, sutured

quartz grains were observed. Orthoclase appears as the most abundant feldspar, with

microcline and plagioclase appearing in lesser abundances. The feldspars within the

selected samples show significant chemical weathering and sericitization. The general

character of the sediment is illustrated in Figure 4.1.1. The oxides exist as explicit grains,

Texas Tech University, Colton Mallett, May 2021 as a cement and within cutans. Clay minerals appear as distinct agglomerations encompassing feldspar. The character of the clay and oxides is illustrated in Figure 4.1.1.

The carbonate present occurs as both spar and micrite cement on and between grains (Fig

4.1.1).

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Figure 4.1.1. Selected photomicrographs illustrating variations in mineralogy, grain size and authigenic calcite and clays. (A) Sample BL-1 under 10x magnification and plane polarized light (PPL). (B) Sample BL-1 under 10x magnification and cross polarized light (XPL). (C) Sample BL-11 under 20x magnification and PPL. (D) Sample BL-11 under 20x magnification and XPL. (E) Sample BL-9 under 20x magnification and PPL. (F) Sample BL-9 under 20x magnification and XPL. (G) Sample BL-18 under 10x magnification and PPL. (H) Sample BL- 18 under 10x magnification and XPL.

4.2 X-Ray Diffraction of Whole-rock Fraction

Whole-rock modal mineralogy data, acquired through the quantification of X-ray diffraction spectra (Fig 4.2.1), was integrated with optical petrographical observations to assess the variation in global sample and clay-size-fraction mineralogy. The Bushland

Core’s mineralogy is relatively homogenous, which is evident when the diffractograms of the select samples are compared (Fig. 4.2.1). The greatest variability is attributed to differences in the relative abundance of shared mineral assemblages (Fig. 4.2.2). Phases within the Bushland Core are predominately quartz, albite, and K-feldspar, accompanied by lesser quantities of mica/illite, mixed layer minerals (MLMs), and variable amounts of pedogenic calcite. Detectable accessory phases, oxides and heavy minerals include hematite, pyroxenes, and amphiboles.

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The mineralogical data for the Bushland Core, normalized without calcite, are

supplied in Table 4.2.1, while multiple mineralogical trends are illustrated in Figure

4.2.2. Appendix E contains the non-normalized mineralogical data.

Calcite has the highest degree of variability relative to other phases present in the samples (Fig. 4.2.2; Appendix E). The uppermost sample, BL-1 has virtually no calcite present, while BL-18, at the base of the Bushland Core has approximately 30 wt% calcite, furthermore, the deepest four samples contain the highest proportion of calcite

(Fig. 4.2.1). The variability of calcite affects the relative weight percent of other phases

(Fig. 4.2.2), such that, when calcite is removed from the modal mineralogy results and the remaining phases are normalized to 100%, a more illustrative trend is revealed (Fig

4.2.2, B). There are several justifications for why the modal mineralogy should be described without calcite, including the following: 1. The observed heterogeneity of calcite accumulations in these calcisols precludes meaningful interpretations of laterally adjacent sediments based on the calcite present in the core; 2. The degree of variability of calcite could have been exacerbated by the sampling process, as addressed in the

Methods chapter; 3. Most importantly, the calcite in these soils is secondary, and do not reflect the original mineralogy of the sediments at the time of deposition. Due to these reasons, it is prudent to assess mineralogical variations without the impact of calcite. The remaining mineralogical data in this section will be reviewed using the data following normalization without calcite (Table 4.2.1). The role of calcite will be reviewed in the discussion.

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Figure 4.2.1. Diffractograms of all the samples analyzed using XRD, which illustrate all identified phases between 3-70 °2θ. Major mineral peaks are labeled. Mineral abbreviations: cal – calcite; fsp – feldspars; IS – illite-smectite; qtz – quartz. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

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The most prevalent mineral is quartz, which averages 63.5 wt% in the samples,

furthermore, the relative abundance of quartz varies between 81.3 wt%, in sample BL-17, at a maximum, and 51.2 wt%, in sample BL-8b, at a minimum. Aside from calcite, the quantity of quartz has the highest degree of variability within the Bushland Core (Table

4.2.1). The proportion of quartz generally diminishes down the core profile, from sample

BL-1 to sample BL-8b, at which point, the trend reverses, and generally increases the remainder of the way to the deepest sample, BL-18 (Fig. 4.2.2, B). Although this major trend is observed, the quantity of quartz is slightly undulose, with divergences from the general trend appearing as increased “peaks” in samples BL-6, BL-BL-9, BL-11, and BL-

15 (Fig 4.4.2, B). The deepest 3 samples, BL16 through BL-18 have relative percentages of quartz distinctly greater than the remaining samples.

Feldspar abundances are significantly lower than quartz within the extent of the core, and combined, average around 20 wt%, with K-feldspar approximately 23% more abundant than albite, although samples BL-1, BL-9, and BL-9b have albite abundances greater than K-feldspar abundances (Table 4.2.1). Proportion of feldspars tend to be relatively consistent down the core profile, although the lowest 3 samples are notably deprived in feldspar content. Due to the similar crystallography of K-feldspars, and to adequately quantify the modal percentages of feldspar, albite has been taken as a measure of all plagioclases while K-feldspar stands for all potassium bearing feldspars (Table

4.2.1).

Mica, in this instance, is the sum of muscovite, biotite, plus discrete illite, because these phyllosilicates are crystallographically similar which leads to the overlap of their diffraction maxima (Fig. 4.2.1). Mica/illite comprises approximately 5.1 wt% of the

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sample and composes up to 10 wt% in select samples. Kaolinite does not generally occur

in abundances greater than 1 wt%, although several samples contained up to, or slightly

more than 2 wt%. Hematite occurs throughout the core, averaging less than a percent of

the total weight. Enstatite, riebeckite and magnesiohornblende all occur in abundances

less than 5 wt% and were quantifiable in only a few samples.

While the preparation of non-oriented, global fraction samples does not allow for

a precise characterization of expandable clay phases, due to the occurrence of several

generations of low crystallinity (observed as a broad peak) mixed layer minerals and

smectites, this preparation mode does allow for a relatively accurate mineral

quantification. MLMs and smectite were quantified from a montmorillonite phase and an

50-50 illite-smectite phase selected from the PDF4+ database issued by the International

Centre for Diffraction Data using RIR.

The average content of phyllosilicates in the samples is approximately 15.4 wt%.

Smectite within the Bushland Core averages less than 3.7 wt%. Illite-smectite is slightly more abundant than smectite (montmorillonite) and averages 5.8 wt% within samples.

Based on petrography and SEM observations (see following subchapter), it may be that, due to the poorly crystalline nature of clay intermediates and smectite minerals, the quantity of clay phases was somewhat underrepresented. The trends of the clay minerals in the core are illustrated in Figure 4.2.2.

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D Figure 4.2.2. Stacked bar charts illustrating the quantified whole rock mineralogy of select samples identified using XRD (A) Complete representation of minerology in select samples. (B) Representation of mineralogy within select samples, normalized to 100%, following the removal of calcite values. (C) Representation of quartz, clay minerals and mica, normalized to 100%. (D) Representation of only clay minerals and mica, normalized to 100%.

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Table 4.2.1. Modal mineralogy of the global fraction based on quantification of XRD diffractograms, normalized to 100% following the removal of quantified calcite. A similar table, which includes calcite, is supplied in the appendices. Depth Mica + ∑ ∑ Swell. Sample Qz Ab Kfs Sme IS Kln Hem Pyx Amp (m) Ilt Phyllo Clays BL-1 0.3 70.01 11.84 10.33 1.91 1.91 2.71 1.00 7.52 4.61 0.30 0.00 0.00 BL-2 0.7 70.82 7.50 12.46 2.94 2.23 3.14 0.71 9.02 5.37 0.20 0.00 0.00 BL-3 1.45 61.76 7.72 9.14 6.89 5.82 7.13 0.71 20.55 12.95 0.24 0.00 0.59 BL-4 1.7 60.50 12.66 14.97 5.43 1.31 4.52 0.40 11.66 5.83 0.20 0.00 0.00 BL-5 2.4 56.47 9.13 10.93 5.12 4.91 8.22 0.70 18.96 13.14 0.10 3.91 0.50 BL-6 2.6 62.84 10.11 15.16 3.05 2.11 5.79 0.63 11.58 7.89 0.32 0.00 0.00 BL-7 3.5 52.92 11.58 15.55 7.94 4.74 6.17 0.88 19.74 10.92 0.22 0.00 0.00 BL-8 4.2 56.29 12.35 15.85 4.55 4.55 5.24 1.05 15.38 9.79 0.12 0.00 0.00 BL-8b 5.1 51.17 11.17 13.30 8.12 4.67 9.14 2.23 24.16 13.81 0.20 0.00 0.00 BL-9 5.9 63.55 12.78 11.01 3.63 3.63 3.85 1.32 12.44 7.49 0.22 0.00 0.00 BL-9b 6.5 52.06 15.26 8.76 10.00 5.26 6.70 1.34 23.30 11.96 0.21 0.00 0.41 BL-10 7.3 57.86 8.53 11.00 4.11 8.43 4.52 0.92 17.99 12.95 0.10 4.52 0.00 BL-10b 7.75 60.88 9.66 14.67 7.22 2.33 3.84 1.05 14.44 6.17 0.35 0.00 0.00 BL-11 8.35 69.11 10.26 12.37 2.52 2.31 2.41 0.91 8.15 4.73 0.10 0.00 0.00 BL-12 8.9 63.26 8.09 12.47 3.26 3.26 8.99 0.45 15.96 12.25 0.22 0.00 0.00 BL-13 9.55 62.32 10.39 10.69 4.68 6.52 4.58 0.61 16.40 11.10 0.20 0.00 0.00 BL-14 10.2 60.74 6.46 13.44 6.67 5.50 6.46 0.53 19.15 11.96 0.21 0.00 0.00 BL-15 11.1 70.61 5.92 10.51 4.39 3.06 4.49 0.61 12.55 7.55 0.41 0.00 0.00 BL-15b 11.6 62.33 8.89 9.15 9.68 1.72 6.50 1.59 19.50 8.22 0.13 0.00 0.00 BL-16 12.4 76.23 4.56 7.08 3.12 3.36 5.04 0.48 12.00 8.40 0.12 0.00 0.00 BL-17 13.1 81.31 3.66 5.30 3.03 1.89 3.79 0.76 9.47 5.68 0.25 0.00 0.00 BL-18 13.8 81.75 2.19 5.84 3.07 1.17 5.55 0.29 10.07 6.72 0.15 0.00 0.00 Average 63.52 9.09 11.30 5.09 3.70 5.77 0.87 15.44 9.47 0.21 0.38 0.07 Abbreviations: Qz: quartz, Ab: albite, Kfs: potassium feldspar, Mica + Ilt: sum of muscovite and illite, Sme: smectite, IS: illite/smectite, Kln: kaolinite, ∑ Phyllo: sum of phyllosilicate phases, Swell. Clays: sum of swelling clays, Hem: hematite, Pyx: pyroxene, Amp: amphibole. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

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4.3 X-Ray Diffraction of Clay Fraction

The clay-mineral assemblage is characterized by the predominant occurrences of illite (Ill), low crystallinity illite (Ill low), multiple generations of interstratified illite-

smectite (IS), discrete smectite (SSS), minor amounts of kaolinite (Kln) and occasionally

chlorite (Chl). While the assemblages do not necessarily vary in terms of present phases,

there are distinct trends in the modal percentages of the existing phases. Most notable is

the dramatic increase in illite and decrease in expandable phases in BL-17 and BL-18. IS

occurs as multiple generations, not only within individual samples, but additionally from

sample to sample.

The phases in the clay fraction (<2 µm) were first investigated using X-ray

diffraction techniques on each sample, 1 oriented (AD), 1 treated with ethylene glycol

(EG) and 1 heated (HT) to 500°C for 1 h to observe peak migrations. Following initial

observation of AD, EG and HT diffractogram patterns, the clay mineral assemblages

 were further deconvoluted and quantified using the Sybilla modeling software (Table

4.3.1).

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Figure 4.3.1. Representative XRD patterns of clay fraction analyzed samples. EG stands for sample treatment via saturation with ethylene-glycol, HT stands for sample treatment via heating and AD stands for air-dried samples. Abbreviations: ISS: illite- smectite rich in smectite components, SSS: discrete smectite, I/S: ordered (R1) illite- smectite, Kln: kaolinite, Chl: chlorite, Qtz: quartz, Cal: calcite. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

4.3.1 Qualitative Identification of Phases in the Clay Fraction

Illite: Discrete illite (Ill) occurs in all analyzed samples and is recognized in XRD

diffractograms by a basal 001 diffraction peak at 8.82 °2θ (10 Å), that does not shift

when heated or saturated in ethylene-glycol (Fig. 4.3.1) (Moore and Reynolds, 1997).

The rational illite 00l reflection series, compared to the non-periodical diffraction

maxima of mixed layer illite-smectite, as well as the potential for IS to swell when

glycolated, can be used to differentiate between discrete illite and interstratified illite-

smectite (Fig. 4.3.1; Środoń and Eberl, 1984). Diffractogram patterns indicate a likely

Texas Tech University, Colton Mallett, May 2021 occurrence of a secondary, low crystallinity illite phase at 8.5 °2θ occurring as a shoulder to the left of the main basal 001 diffraction peak at 8.82 °2θ (10.6 Å).

Illite-Smectite. Expandable interstratified illite-smectite was initially identified by preliminary inspection of XRD spectra with a broad 5.1 °2θ (17 Å) peak, which emerged following EG solvation (Reynold and Hower, 1970).

Smectite. Discrete smectite was initially identified in AD sample diffractograms by a relatively broad peak at 6 °2θ (15 Å) and a shift to 5.1 °2θ (17 Å) following EG treatment.

Chlorite and Kaolinite. Chlorite (Chl) and Kaolinite (Kln) were observed to occur discretely or in combination within the selected samples (Fig. 4.3.1). Chlorite has a sequence of basal diffraction peaks, which can overlap, or nearly overlap the kaolinite peaks, but depending on the content, chemistry, and thickness of the phases, in combination with the attenuation of the instrument, the overlap can vary (Moore and

Reynolds, 1997). In sample BL-17, the distinction between kaolinite and chlorite is made clear by a division of the kaolinite 001 diffraction peak at 12.33 °2θ (7.2 Å) and the chlorite 002 diffraction peak at 12.51 °2θ (7.1 Å) (Fig. 4.3.1). Even clearer evidence of the existence of both phases in Sample BL-17 is illustrated by the complete separation of the kaolinite 002 peak at 24.9 °2θ (3.58 Å) and the chlorite 004 peak at 25.1 °2θ (3.5 Å)

(Fig. 4.3.1.) Given a weak reflection for both phases at 12.5 and 25 °2θ , the interpretation of which phase is present may not be straightforward. At present, only kaolinite, only chlorite or both phases may be present. Simply by heating, the chlorite

002 reflection is weakened but not eliminated, allowing for the determination of which

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phases are present (Moore and Reynold, 1997; Figure 4.3.1). No change is seen at 12.5

°2θ after heating for any of the remaining samples, indicating that chlorite only occurs in

Sample BL-17; all other samples do not have chlorite, but do contain varying amounts of

kaolinite.

4.3.2 Grouping Clay Mineral Assemblages

Although the clay minerals assemblages are largely correspondent in all the

analyzed samples their crystallinity and interstratification particularities allowed for

identification of 3 distinct categories, which are further divided into 6 individual groups

(Fig. 4.3.2). Category 1 represents the most common assemblages and is composed of

groups 1 through 4. Category 2 and 3 are less prevalent and are composed of groups 5

and 6, respectively.

The category 1 is defined by an assemblage of well-defined smectite, several generations of illite-smectite, low crystallinity illite, illite, and kaolinite (Fig.4.3.2). The 4 groups in category 1 are distinguished from each other based on the location and shape of the expandable mineral peaks. Samples in Group 1 display a relatively narrower first basal peak of IS, centered at 14.29 Å, and have a slight “shoulder” to the left of the main peak. Samples in Group 2 exhibit a broad, symmetrical smectite peak at 15.04 Å.

Samples in Group 3 are defined by a very large, symmetrical 001 of smectite centered at

15.58 Å. The final group in category 1, Group 4 is defined by a broad 001 smectite peak, centered at 15.04 Å, with a “shoulder” to the left of the main peak representing an additional, >15.04 Å smectite phase (Fig 4.3.2).

Group 5 is defined by a broad IS peak similar to Group 4 but has a very distinct peak of 10.6 Å low illite (Fig. 4.3.2). Group 5 is also the only group to have a sample that 48

Texas Tech University, Colton Mallett, May 2021 contains chlorite (Fig. 4.3.2). The final group 6 varies greatly from the other groups; namely the first basal peaks of several IS species form an intensity plateau between 6.5 and 8 °2θ, whilst the 001 reflex of smectite is rather broad (Fig 4.3.2).

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Figure 4.3.2. Groupings of clay mineral assemblages based on air dried X-ray diffractogram patterns. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

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4.3.3 Modeling and Qualitative Description of Clay Minerals

Following an initial investigation of clay minerals based on clay mineral

treatments and repetitive measurements under different conditions (Zanoni et al., 2016),

Sybilla software package was used to model (Fig. 4.3.3) and quantify the clay minerals

present in the clay fraction (Fig. 4.3.4; Table 4.3.2).

As outlined in the previous section (c.f. Methods) the modelling relies on a trial-

and-error procedure that provides optimal clay mineral structural and probability

parameters to get the best fit between experimental and calculated patterns. Modelling is

especially helpful with mixed-layer clay minerals whose number, nature and stacking

sequence are considered as modifiable values (e.g. Uzarowicz et al., 2012). In this

research, the experimental patterns were fitted with: (a) discrete smectite (~15.1 Å), (b) an R1 IS (~11.6 Å-13.6 Å), (c) additional R0-R1 ISS (~12.8 Å -14.5 Å), (d) a low crystallinity illite (~10.2-10.6 Å), (e) a discrete illite and (f) kaolinite (Table 4.3.1).

Sample BL-17 was the only sample to contain measurable amounts of chlorite

(Table 4.3.2). Taking into account that this research looks into discrete smectite and smectitic layers in illite-smectite as indicators of successive hydrolysis which leads to the soil formation at middle latitude regions (Millot, 1971), it was important to accurately assess their amounts in modelled diffractograms. Therefore, the trends adumbrated by discrete smectite, the sum of the all illite-smectite intermediates and the sum of illitic minerals were provided separately (Fig. 4.3.4C).

Discrete smectite and ∑ IS phase quantities tend to trend inversely of each other, with smectite occurring in lesser quantities, with a few exceptions (Fig 4.3.4). This is

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however an expected trend for warm-temperature Quaternary soils where in a given profile smectite evolves from illite-smectite through the process of layer-by-layer transformation (i.e. pseudo-bisiallitization, Paquet and Clauer, 1997). The quantity of smectite was observed to reach a maximum and appeared in greater amounts of IS in samples BL-7, BL-10, and BL-13. This inverse trend continues down the Bushland Playa

Core profile until sample BL-16, where the occurrence of these phase mirror each other and both decrease drastically. The abundances of illite, except at the bottom of the profile, remain relatively consistent. Kaolinite is relatively consistent down the core profile (Fig 4.3.4).

Table 4.3.1. Summarized Sybilla modelling parameters and characterization of identified clay mineral phases. Sample SSS ISS R0 to R1 IS R1 Ill Low Ill (~15.1Å) (~12.8-14.5Å) (~11.6-13.6Å) (~10.2- 10.6Å) (~10.02Å) 2θ d- Space 2θ d- Space 2θ BL-1 6.12 14.44 6.7 13.19 7.56 11.69 8.63 10.23 8.8 10.04 BL-2 5.78 15.29 6.46 13.63 7.6 11.63 8.57 10.32 8.82 10.02 BL-3 5.91 14.97 6.61 13.38 7.63 11.58 8.64 10.23 8.8 10.04 BL-4 5.98 14.78 6.91 12.78 7.69 11.49 8.55 10.34 8.8 10.04 BL-5 5.8 15.24 - - 7.63 11.58 8.61 10.27 8.8 10.04 BL-6 5.96 14.83 6.9 12.82 7.33 12.06 8.55 10.34 8.82 10.04 BL-7 5.87 15.06 6.9 12.82 7.45 11.86 8.52 10.38 8.82 10.02 BL-8 5.92 14.92 - - 7.36 12 8.52 10.38 8.82 10.02 BL-8b 5.69 15.53 6.45 13.71 7.49 11.8 8.59 10.3 8.84 10 BL-9 5.82 15.2 6.29 14.12 7.22 12.24 8.59 10.29 8.84 10 BL-9b 8.94 14.87 7.06 12.52 7.24 12.21 8.53 10.36 8.86 9.98 BL-10 5.67 15.58 - - 7.42 11.92 8.46 10.45 8.82 10.02 Bl-10b 5.94 14.87 - - 7.04 12.56 8.35 10.58 8.82 10.02 BL-11 5.92 14.92 6.54 13.52 7.49 11.8 8.62 10.25 8.79 10.06 BL-12 5.83 15.15 - - 6.9 12.81 8.41 10.52 8.82 10.02 BL-13 5.71 15.58 - - 7.42 11.92 8.46 10.45 8.8 10.02 BL-14 5.78 15.29 6.03 14.65 7.47 11.83 8.59 10.3 8.82 10.02 BL-15 5.78 15.29 5.96 14.83 7.45 11.86 8.59 10.3 8.82 10.02 BL-16 5.96 14.83 - - 7.38 11.97 8.55 10.34 8.82 10.02 BL-17 5.82 15.2 6.1 14.48 7.06 12.52 8.37 10.56 8.82 10.02 BL-18 5.82 15.2 - - 6.5 13.6 8.32 10.63 8.84 10

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Figure 4.3.3. Sybilla modelling of the experimental XRD spectra as the example of a representative sample of each of the identified groups. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

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A C

Figure 4.3.4. Illustrations of the clay mineral abundances within the Bushland Playa Core. (A) Stacked bar graph illustrating the respective quantities of all identified clay minerals. (B) A simplified bar graph illustrating the relationship between discrete smectite, the sum of interstratified illite-smectites and the sum of illite phases. (C) Line graph illustrating the simplified trends of expandable clay minerals and illites in relation to depth in the Bushland Playa Core. Abbreviations: SSS: smectite, ISS: illite- smectite (rich in smectite) R0 to R1, IS: illite-smectite R1, Ill Low: low crystallinity illite, Ill: illite, Kln: kaolinite, Chl: chlorite. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

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Table 4.3.2. Modal mineralogy of the clay fraction based on quantification of XRD spectra and a breakdown of mixed layer illite and mixed layer smectite within all mixed layer phases. ∑ ∑ISS R0- Depth ∑ Swell. ISS R0- Kln/ Sample Phyllo SSS IS R1 Ill Low Ill Kln Chl R1 and IS (m) Clay R1 SSS (wt%) R1 BL-1 0.3 7.52 4.61 9.6 19.9 29.7 27.9 8.9 4.1 0.0 49.6 0.43 BL-2 0.7 9.02 5.37 7.1 28.9 27.2 23.5 9.1 4.1 0.0 56.1 0.58 BL-3 1.45 20.55 12.95 20.6 16.5 32.3 16.7 8.3 5.6 0.0 48.8 0.27 BL-4 1.7 11.66 5.83 18.5 6.8 31 22.8 16.2 4.7 0.0 37.8 0.25 BL-5 2.4 18.96 13.14 20.9 0.0 32.3 24.6 17.2 5 0.0 32.3 0.24 BL-6 2.6 11.58 7.89 19.2 2.8 30.8 29 13.6 4.4 0.0 33.6 0.23 BL-7 3.5 19.74 10.92 30 0.8 23.4 28.4 10.7 6.7 0.0 24.2 0.22 BL-8 4.2 15.38 9.79 19.2 0.0 33.2 28.9 11.8 6.8 0.0 33.2 0.35 BL-8b 5.1 24.16 13.81 9.9 2.6 40.5 29.9 13.4 3.7 0.0 43.1 0.37 BL-9 5.9 12.44 7.49 20 1.4 40 28.2 7 3.4 0.0 41.4 0.17 BL-9b 6.5 23.30 11.96 12.8 4.9 30.9 30.4 15 6 0.0 35.8 0.47 BL-10 7.3 17.99 12.95 35.1 0.0 23.4 25.9 12.1 3.4 0.0 23.4 0.10 BL-10b 7.75 14.44 6.17 12.5 0.0 27 37.6 18.8 4.1 0.0 27.0 0.33 BL-11 8.35 8.15 4.73 15.1 11.1 29.4 33.8 5.8 4.7 0.0 40.5 0.31 BL-12 8.9 15.96 12.25 16.6 0.0 36.8 31.2 12.5 3 0.0 36.8 0.18 BL-13 9.55 16.40 11.10 29.8 0.0 22.4 28.8 14.8 4.3 0.0 22.4 0.14 BL-14 10.2 19.15 11.96 11.5 16 28.9 28.2 11.1 4.3 0.0 44.9 0.37 BL-15 11.1 12.55 7.55 10.2 12.7 33.7 29.1 10.1 4.1 0.0 46.4 0.40 BL-16 12.4 18.17 8.22 23.4 0 34.3 31.3 7.4 3.6 0.0 34.3 0.15 BL-17 13.1 12.00 8.40 19.3 9.2 12.5 48.4 7.1 2.4 1.00 21.7 0.12 BL-18 13.8 9.47 5.68 5.6 0.0 18.3 73.1 2.2 0.8 0.0 18.3 0.14 Avg. 15.24 9.53 17.5 6.4 29.43 31.32 11.1 4.25 1 35.8 Abbreviations: Ill: Discrete illite, SSS: Discrete smectite, Kln: kaolinite, Chl: chlorite, IS: illite-smectite, ∑ I/S: sum of all IS phases/ generations, ∑ Ill in ISS: sum of all illite in IS phases, ∑ Sm in IS: sum of all smectite in IS phases. ∑ IS includes IS R1, all generations of IS R0 and all generations of ISS R0. Total Clays represents the sum of clay minerals from the quantified whole rock XRD results. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

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4.4 Scanning Electron Microscopy – Energy Dispersive Spectroscopy

Scanning electron microscopy – energy dispersive spectroscopy techniques and backscatter electron (BSE) imaging were utilized to further address the apparent variations in clay mineral chemistry and speciation observed through X-ray diffraction analysis (Fig. 4.2.2; Fig. 4.3.1). Modeling of clay fraction X-ray diffractograms revealed that multiple species of mixed layer illite-smectites and smectites exist within individual samples (Figure 4.3.3). These variations in chemical composition of clay minerals, which exists as matrix in the Bushland Playa Core (Fig. 4.4.1; Table 4.4.3), were also observed in EDS spectra. The clay minerals in the Bushland Playa Core (Table 4.4.3) exhibit an array of chemical compositions which tend to stray from typical compositions (Table

4.4.1). Furthermore, clay mineral precursors, particularly muscovite (Table 4.4.2), exhibit chemical compositions that do not necessarily match typical compositions (Fig. 4.4.1;

Table 4.4.1). Several other minerals were also analyzed (Table. 4.4.2)

Table 4.4.1. Idealized chemical composition (wt%) of pertinent minerals. Mineral Mnt IS Ill Ms Kln Ab An Or SiO2 66.66 63.26 54 45.87 52.95 68.71 45.88 65.39 Al2O3 25.76 23.84 17.2 38.69 45.72 19.63 34.31 18.45 FeO 1.08 3.25 1.85 0.66 0.83 MgO 4.12 3.75 3.11 0.10 0.16 K2O 0.05 1.49 7.26 10.08 0.03 0.03 0.11 14.76 CaO 25.11 0.15 0.47 0.22 18.28 Na2O 0.14 0.06 0.64 11.72 0.56 1.08 Mnt- montmorillonite; ISS- smectite rich illite-smectite; IS- ordered illite-smectite; Ill- illite; Ms-muscovite; Kln-kaolinite; Ab- albite, An- anorthite, Or- orthoclase. Values adapted from Rock-Forming minerals v.4 framework silicates; v3 sheet silicates after. Środoń et al. (1986). Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

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The analyzed non-clay minerals include the dominant phases such as feldspar, quartz, calcite, and accessory phases, zircon, ferrohornblende, magnetite, and rutile.

Analyzed plagioclase has chemistry typical of albite endmember, although several grains were slightly enriched in CaO. Most of alkali feldspars were found to be compositionally orthoclase. BSE imaging revealed that feldspars display varying degrees of alteration, although due to the paucity of analyzed grains, little can be said about trends from sample to sample (Fig. 4.4.1). Calcite appears in the groundmass as well as in the form of large

“spar” nodules. Other identified minerals including ferrohornblende, zircon, magnetite and rutile have expected chemical compositions (Table 4.4.2).

Figure 4.4.1. BSE images of sample BL-7, BL-14 and BL-13, illustrating character of major minerals and the clay composed matrix. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

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Table 4.4.2. Representative EDS analyses (wt%) of select minerals. Sample BL-7 BL-13 BL-11 BL-7 BL-13 BL-14 BL-11 BL-7 BL-7 Site 2 2 1 3 1 4 1 1 1 Analysis 15 96 209 46 54 207 208 1 6 Mineral Ab Ab Ab Or Or Or Or Qz Cal SiO2 69.5 66.9 70.8 66.4 64.9 66.3 65.1 100 3.7 Al2O3 19.5 20.2 19.3 18.7 18.9 19.1 18.3 0.0 1.2 FeO 0.0 1.5 0.0 0.0 0.0 0.0 0.0 0.0 0.0 MgO 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 1.5 K2O 0.0 1.2 0.0 12.7 16.3 11.2 15.6 0.0 0.9 CaO 0.0 2 0.6 0.0 0.0 0.0 0.0 0.0 92.7 Na2O 11 8.3 9.3 2.2 0.0 3.4 0.3 0.0 0.0 TiO2 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 Total 100 100.1 100 100 100.1 100 99.3 100 100

Sample BL-13 BL-11 BL-11 BL-13 BL-13 BL-11 BL-11 BL-11 BL-14 Site 2 2 2 2 2 1 1 2 1 Analysis 89 250 260 74 83 225 242 252 153 Mineral Ms Ms Ms Fe Hbl Rt Fe Ox Zrn Mag Cal No SiO2 58.3 53.9 51.5 43.3 0.7 0.0 32.4 1.1 29.1 Al2O3 27.3 28.8 34.8 12.4 0.0 7 0.0 0.0 5.7 FeO 4.6 4 3 21.4 0.0 84.9 0.0 97.8 0.0 MgO 2.4 2.9 1.3 8.6 0.0 2.5 0.0 0.0 1.8 K2O 5.6 10.5 9.4 0.6 0.0 0.0 0.0 0.0 0.0 CaO 0.9 0.0 0.0 11.8 0.0 0.0 0.0 0.0 63.4 Na2O 0.5 0.0 0.0 1.1 0.0 0.0 0.0 0.0 0.0 TiO2 0.5 0.0 0.0 0.9 99.3 4.5 0.0 1.2 0.0 Zr 0.0 0.0 0.0 0.0 0.0 0.0 67.6 0.0 0.0 Total 100.1 100.1 100 100.1 100 98.9 100 100.1 100 Ab- albite; Or- orthoclase; Qz- quartz; Cal- Calcite; Cal No- calcite nodules, Ms- Muscovite; Fe Hbl- ferrohornblende; Rt- rutile; Fe Ox- iron oxide; Zrn- zircon; Mag- magnetite. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

The paleosol substrate made of the clay minerals and discrete muscovite account for most of recorded EDS spectra, of which the representative results are illustrated in

Table 4.4.3. Interestingly, all 10Å phases along with IS share similar trends in terms of the enrichment or impoverishment of its cations from various structural positions (e.g. Fe,

Al, and K). The chemical composition of analyzed muscovite is significantly reduced both in tetrahedral Al and interlayer K, such that the composition of the muscovite in the selected samples are intermediate between an unaltered muscovite and a typical illite.

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This is a common feature of the ongoing hydrolysis whereby illite stems through the

physical and chemical erosion and weathering of crystalline rocks and can be related

either to the weathering of feldspars or degradation of true mica (e.g. Meunier and Velde,

2004; Liu et al., 2007). A typical muscovite is composed of approximately 40 wt% Al2O3

and 12 wt% K2O (Rieder et al., 1988), but the muscovite from the Bushland Playa Core is

generally composed of less than 30 wt% Al2O3 and less than 10 wt% K2O. Several

muscovite grains contained less than 20 wt% Al2O3 and less than 5 wt% K2O, and thus represent an illite-smectite pseudomorph (e.g. Šegvić et al., 2016). In addition to the apparent reduction in Al2O3 and K2O, these former muscovites are enriched in FeO, with

specific grains containing as much as 15 wt%. Considering the distinct enrichment in

FeO and other chemical signatures of illitization mentioned heretofore, once can argue that most of the muscovite grains analyzed in this study should be termed illitized muscovite or simply illite.

The progressive loss of K is clear in EDS spectra of illite and IS species. Illite

minerals contain as much as 8 wt% FeO and as little as 3 wt% K2O. Typically illite

should contain over 9% K2O and approximately 2 wt% FeO (Środoń et al., 2006). Illite-

smectite and ISS are also reduced in K2O and enriched in FeO. EDS chemistry did not

permit a differentiation between discrete smectite and smectite-rich illite-smectite because both are marked with similar identification signatures (contents of CaO (0.7-1 wt%), octahedral FeO/MgO (2-3), tetrahedral Al2O3/SiO2 (0.20-0.25); Welton, 1984).

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Table 4.4.3. Representative EDS analyses (wt%) of clay mineral phases from selected samples. Sample BL-7 BL-7 BL-13 BL-13 BL-11 BL-11 BL-7 BL-14 BL-14 BL-14 BL-11 BL-7 BL-7 BL-7 Site 2 3 2 2 1 2 3 1 1 2 1 1 1 3 Analysis 20 39 75 95 235 249 41 155 I69 174 211 5 11 32 Mineral I/SS I/SS I/SS I/SS I/SS I/SS I/S I/S I/S I/S I/S Ill Ill Ill SiO2 66.8 65 69.4 73.5 67 72.1 65.6 61.2 62.3 60.9 58.9 60 56 50.9 Al2O3 18.3 18.7 18.6 14.2 19.4 16.4 19.8 20.9 23.9 24.6 20.4 21.9 21.3 22.2 FeO 8.6 9.3 5.5 6.3 6.2 5.1 7.1 6.5 4.9 8.2 9.2 6.3 7.7 13.9 MgO 2.6 3.2 2.6 2.9 3.1 2.8 2.6 3.3 3.7 3.4 4.6 3.2 3.2 4.4 K2O 2.4 2.5 2.3 2.2 2.8 2.6 3 3.1 3.9 3 4.6 4.5 3.3 5.2 CaO 0.7 1 1 0.7 0.9 0.7 0.9 4.6 1 0 0.9 3 7.3 1 Na2O 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.3 0.4 0 0.0 0.2 0.3 0 TiO2 0.7 0.0 0.4 0.0 0.5 0.0 0.0 0.0 0 0 1.3 0.7 0.0 2.2 Total 100.1 99.7 99.8 99.8 99.9 99.7 99 99.9 100.1 100.1 99.9 99.8 100 99.8 MgO/K2O 1.1 1.3 1.1 1.3 1.1 1.1 0.9 1.1 0.9 0.9 1,0 0.7 1.0 0.8

Sample BL-7 BL-13 BL-14 BL-14 BL-7 BL-13 BL-14 BL-11 BL-14 BL-14 BL-11 Avg. Avg. Avg. Site 3 3 2 2 2 1 1 1 2 2 2 Analysis 45 126 177 179 25 65 165 233 183 186 248 Mineral Ill Ill Ill Ill Ms I M I M I M Kln Kln Kln I/S Ill Kln SiO2 58.6 60 62.3 62.9 52.9 47 48.4 51.7 62.1 55.8 50.2 64.6 59.7 55.5 Al2O3 26.5 24.2 22.7 23.2 28.8 22.9 19.5 35.5 32 44.2 36.7 20.7 23.4 39.9 FeO 6.7 7.2 6.8 5.4 4 13.3 15.1 1.5 2.6 0 1.4 6.6 7.9 2.5 MgO 2.1 3.4 3.8 2.9 3.8 11 6.9 1.4 1.5 0 1 3.1 3.6 0.7 K2O 3.3 3.4 3.2 5.3 10.3 3.8 4.6 8.6 1.4 0 8.1 3.0 3.3 0.9 CaO 2 0.6 2 0.7 0.0 0.7 2.1 0.0 0.4 0 0 1.1 1.5 0.4 Na2O 0.8 0.0 0 0.5 0.2 0.0 0.4 0.6 0 0 0 0.2 0.03 0.08 TiO2 0 0.8 0 0 0.0 1.2 2.8 0.7 0 0 0 0.3 0.5 0 Total 100 99.8 99.8 100.1 100 99.9 99.8 100 99.9 100 100 - - - MgO/K2O 0.6 1.2 1.2 0.5 0.4 2.9 1.5 0.2 1.1 - 0.1 - - - Ill- illite; I/S- illite - smectite; I/SS- illite – smectite rich in smectite; Kln kaolinite; I M- illitized muscovite. I/S (S)- includes the average of both illite-smectite and illite- smectite rich in smectite. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

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4.5 X-Ray Fluorescence

The major element geochemistry of analyzed samples and standards are included

in Appendix B. The major and trace element geochemistry is summarized in Tables 4.5.1 and 4.5.2. The sum of the major oxides and LOI for the analytes were between 102 and

114 wt%, therefore the values of major oxides were normalized to 100. The trace element values, reported in ppm, were not adjusted. The trends of major oxides in relation to stratigraphy are illustrated in Figure 4.5.1

The Bushland Core is generally geochemically homogenous, with SiO2 as the

most abundant major oxide. Samples contain SiO2 in abundances between 58 wt% and 76

wt %, and values average 69 wt%. There is no overall trend, although there are several

instances of dramatic decreases in the wt% of SiO2. The first deflection from the

surrounding samples occurs at 1.45 meters below ground surface (mbgs) (BL-3), where

the concentration of SiO2 drops from approximately 70 wt% down to 58 wt%. The two

other instances of decreases SiO2, in relation to surrounding samples, occurs at 3.5 mbgs

(BL-7) and 8.9 mbgs (BL-12). These three deflections are coincidentally the three lowest

concentrations of SiO2. In the instances of BL-3 and BL-12, the samples directly below have maximum SiO2 abundance. The next most abundant oxides, Fe2O3 and Al2O3 follow a trend established by SiO2, with Al2O3 abundances always slightly greater than

Fe2O3. Fe2O3 content averages 3 wt% and Al2O3 averages 9 wt%. These two elements

show a slight decreasing trend in the bottom 5 m of the core. The remaining major

oxides, TiO2, MnO, MgO, Na2O, K2O and P2O5¸ occur in abundance less than 3 wt% and

follow a trend of Fe2O3 and Al2O3. Calcium shows the greatest degree of variation, with

abundances between <1 wt% and 15 wt%, and samples averaging 5 wt %. The abundance

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of CaO appears to control the trends of the other oxides, which are inverse to the

concentration of CaO, such that increased abundance of calcite relatively reduces the

abundance of the other oxides. An exception to this is at 7.3 mbgs (BL-10), where SiO2

shows a relative increase while all other oxides show a decrease, including CaO. Note

that at the very deepest sample, BL-18, there is an increase in CaO and MgO, indicating

the presence of a Mg-rich calcite.

Ba and Zr are the most abundant trace elements, averaging 576 and 351 ppm,

respectively (Fig. 4.5.2). These two trace elements follow similar trend to the major

oxides, although Ba shows a large increase to 1828 ppm at 1.45 mbgs (BL-3). This

increase corresponds to a peak in CaO, though the remaining samples fluctuate very little

down the remainder of the core. The remaining trace elements average below 130 ppm

(Figure 4.5.2). Cobalt, Cu, Ni, Sc and Zn follow trends similar to that of the major oxides

while Cr and Sr follow an inverse trend similar to the trend of CaO.

Note that the relative increase in CaO does not necessarily match the characteristics of the described core profile (Fig. 4.5.1). While this issue was previously addressed in methods chapter, this discrepancy is worth mentioning again. The sampling method, in combination with the high degree of heterogeneity in accumulated pedogenic calcite in any given sample zone, has resulted in sampled sediment that may not contain a representative amount of calcite, and thus a non-representative quantity of CaO.

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Figure 4.5.1. Major element as oxides, plotted as weight percent against meters below grounds surface (mbgs).

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Figure 4.5.2. Trace elements reported in ppm versus meters below grounds surface. The trace element abundances are illustrated on a log scale.

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Table 4.5.1. Normalized XRF chemistry of paleosols from the Bushland Core (Amarillo, TX). Values are in wt% (oxides) and ppm (elements).

Sample Depth SiO2 Al2O3 CaO Fe2O3 MgO K2O Na2O TiO2 P2O5 MnO LOI Sum BL-1 0.3 68.81 12.32 0.82 4.39 1.18 2.26 0.72 0.70 0.08 0.07 8.7 100 BL-2 0.7 70.80 10.37 2.66 3.44 1.10 2.02 0.67 0.60 0.06 0.04 8.2 100 BL-3 1.45 58.04 7.74 14.96 1.90 1.11 1.61 0.59 0.39 0.09 0.03 13.6 100 BL-4 1.7 73.09 10.24 1.67 3.34 1.15 2.24 0.87 0.61 0.09 0.07 6.6 100 BL-5 2.4 72.62 10.51 1.83 3.47 1.17 2.26 0.87 0.63 0.08 0.07 6.5 100 BL-6 2.7 71.97 10.41 2.14 3.42 1.20 2.18 0.84 0.61 0.09 0.06 7.1 100 BL-7 3.5 61.84 9.15 10.67 2.41 1.24 1.99 0.77 0.44 0.14 0.03 11.3 100 BL-8 4.2 64.83 10.34 7.20 2.84 1.28 2.28 0.91 0.52 0.12 0.04 9.6 100 BL-9 5.9 65.69 12.78 3.36 3.76 1.79 2.52 0.85 0.53 0.09 0.04 8.6 100 BL-10 7.3 75.75 9.42 1.69 2.86 1.25 2.04 0.69 0.40 0.07 0.05 5.8 100 BL-11 8.35 69.94 12.36 0.84 4.20 1.72 2.58 0.77 0.57 0.08 0.05 6.9 100 BL-12 8.9 64.03 9.36 8.71 2.42 1.21 1.99 0.60 0.41 0.09 0.03 11.2 100 BL-13 9.55 74.51 10.41 1.37 2.98 1.24 2.11 0.76 0.50 0.05 0.05 6.0 100 BL-14 10.2 74.35 9.52 2.47 2.79 1.22 2.11 0.68 0.47 0.10 0.04 6.2 100 BL-15 11.1 69.33 9.49 5.55 2.73 1.24 2.06 0.66 0.45 0.16 0.04 8.3 100 BL-16 12.4 68.00 5.77 11.01 1.44 0.99 1.33 0.17 0.25 0.16 0.01 10.9 100 BL-17 13.1 71.87 5.26 9.40 1.32 0.99 1.25 0.18 0.25 0.17 0.01 9.3 100 BL-18 13.8 72.65 3.78 10.45 0.77 1.16 0.95 0.09 0.16 0.03 0.01 10.0 100

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Table 4.5.2. Normalized XRF chemistry of paleosols from the Bushland Core (Amarillo, TX). Values are in wt% (oxides) and ppm (elements). Sample Depth Ba Co Cr Cu Ni Sc Sr Zn Zr BL-1 0.3 615 6 51 23 28 6 112 74 439 BL-2 0.7 564 3 41 20 22 5 98 51 491 BL-3 1.45 1827 2 17 6.7 12 16 172 29 319 BL-4 1.7 557 3 39 15 19 5 114 52 458 BL-5 2.4 545 4 40 14 19 4 113 58 472 BL-6 2.7 559 3 3 10 19 5 112 53 473 BL-7 3.5 579 1 27 7 14 13 172 37 320 BL-8 4.2 585 2 31 10 17 10 183 47 328 BL-9 5.9 516 4 42 11 24 6 157 64 336 BL-10 7.3 482 2 29 13 18 3 99 49 301 BL-11 8.35 504 6 44 20 25 3 116 70 348 BL-12 8.9 509 2 28 9 14 10 138 41 321 BL-13 9.55 471 3 32 10 21 3 111 48 405 BL-14 10.2 476 2 30 12 17 4 99 50 390 BL-15 11.1 507 2 29 10 17 6 121 45 338 BL-16 12.4 395 2 17 4 8 11 120 25 183 BL-17 13.1 377 2 17 3 7 10 115 20 223 BL-18 13.8 299 3 12 5 4 8 116 11 176

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4.6 Laser Ablation Inductively Coupled Plasma Mass Spectrometry

The concentration of trace elements varies as a function of their chemical

properties, the source of the sediment and weathering, diagenesis, and sediment sorting

processes (McLennan, 1989). Though trace elements can be subject to change by these

factors, the rare earth elements (REE) are relatively poorly soluble and immobile

(Michard, 1989). Due to this immobility, REE are preferentially retained in

phyllosilicate-rich sediment, which allow them to be used as an insight into the source

rocks composition (Andersson et al., 2004).

The trace element geochemistry of analyzed samples and standards are included

in Appendix C. The trends of the trace element geochemistry in relation to stratigraphy

are illustrated in 4.6.1. Similar to the trends seen in the major element geochemistry, the

trace element geochemistry is relatively consistent, with major changes in concentration seen as deflections from the overall trend in several samples (Fig. 4.6.1). The overall trend is seen as a relative decrease in concentration down the core profile, with the greatest decrease in concentration seen in the last 3 samples, BL-16 through BL-17 (Fig.

4.6.2).

Ti is the most abundant trace element and occurs in concentrations between 4740 ppm and 1350 ppm, and gradually decreases in concentration down the core profile, with the most noticeable decrease in concentration in samples BL-16 to BL-18 (Fig. 4.6.1).

Significant deflections from the overall trend appear as a decrease in sample BL-3, a decrease in BL-10 followed by an increase in BL-11.

V, Rb, and Ce are the second most abundant trace elements and occur in concentrations between ~110 ppm and ~10 ppm. These trace element concentrations

Texas Tech University, Colton Mallett, May 2021 follow the same trend as Ti (Fig. 4.6.1). La, Y, Nd, and Nb are the third most abundant trace elements and occur in concentrations between ~45 ppm and ~1 ppm (4.6.1, C).

These trace elements concentrations also follow a similar trend to Ti, V, Rb, Ce, La, Y,

Nd and Nb (4.6.1). The remaining trace elements, shown in Fig. 4.6.1, D follow a similar trend to the aforementioned trace elements, but occur in concentrations less than ~15 ppm.

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A B C D

Figure 4.6.1. Trace elements reported in ppm versus meters below grounds surface. The trace element abundances are illustrated on a log scale. (A) Concentration of Ti. (B) Concentration of V, Rb and Ce. (C) Concentration of Y, Nb, La and Nd. (D) Concentration of Pr, Sm, Eu, Gd, Tb, Dy, Er, Yb, Dy, Er, Yb, Lu, Hf, Ta, Pb, Th, and U.

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4.7 K-Ar Dating of Illite

K-Ar dating operates under the assumption that there is an increase in authigenic

illite particles as the size fractions are progressively finer (Środoń et al., 2002; Clauer et

al., 2017). Here, we assume that the 1-2 µm fraction represents predominately detrital

illite, while the <1 µm fraction represents a mixture of detrital and an increased

concentration of authigenic illite particles. A total of 6 illite fractions (both <1 and 1 to 2

µm) were obtained from three individual paleosols, which were sampled at the top,

middle, and bottom of the core profile. XRD measurements of selected samples (Fig.

4.6.2.) showed a significant decrease in the intensity of quartz, which is a clear marker of

clay fraction uniformity and hence its authigenic nature (Zanoni et al., 2019).

These fractions returned K-Ar ages between 93.7± 1.4 Ma and 168.9± 3.4 Ma

(Table 4.7.1). Sample BL-2 returned dates of 168.9± 3.4 Ma for the detrital (1-2µm) illite and 137.2± 1.9 Ma for the mixed (<1 µm) illite fraction. Sample BL-10 returned dates of

144.5± 2.4 Ma for the detrital illite fraction and 93.7± 1.4 Ma for the mixed illite fraction.

Sample BL-17 returned dates of 123.5± 2.8 Ma for the detrital illite fraction and 120.1±

1.7 Ma for the mixed illite fraction.

Samples BL-2 and BL-10 ages indicate sizeable differences between the detrital illite fraction and the mixed illite fraction. The differences in ages between the 1-2 µm fraction and the <1 µm are 31.7 Ma and 50.8 Ma in samples BL-2 and Bl-10, respectively. Sample Bl-17 is markedly distinct compared to the other two samples. Not only is the detrital fraction in BL-17 much younger compared to the detrital fraction of the other samples, but the mixed fraction is also nearly equal to detrital fraction (Table

4.7.1).

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Table 4.7.1. Potassium- Argon age data of the size fractions from the selected samples Sample (depth) BL-2 (0.7 m) BL-10 (7.3 m) BL-17 (13.1 m) Fraction (µm) 1-2 <1 1-2 <1 1-2 <1 K 1.96 1.85 2.18 2.09 1.6 2.05 40 Ar* (mol/g) 6.019E-10 4.573E-10 5.688E-10 3.485E-10 3.556E-10 4.414E-10 %40 Ar* 80.9 75.9 77.9 68.8 72.7 69.8 Age (Ma) 168.9 137.2 144.5 93.7 123.8 120.1 ±1s 3.4 1.9 2.4 1.4 2.8 1.7

Figure 4.6.2. X-ray diffractograms of dated paleosol samples showing lower concentrations of quartz in the finer fraction. For details see the text. Blue (glycolated), Red/Black (air-dried). Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

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CHAPTER 5

DISCUSSION

5.1. Clay-Mineralogy-Based Pedogenetic Scheme

Paleosols represent former soils, which have been incorporated into the geological

record, and as such stand for ancient interaction zones between the bedrock, and the

lithosphere, hydrosphere, biosphere, and atmosphere (Chamley, 1989; Tabor and Myers,

2015; Pevehouse et al., 2020). Taking this into account, it is clear why the paleosols have

been looked at in numerous studies that tried to unveil the physical, biological, and

chemical information about the conditions near Earth’s surface in the geological past

(Kraus, 1999; Mark and James, 1994; Cotton and Sheldon, 2012). In the last few decades,

potentially the most important driving force behind paleosol studies was to examine past

rates of weathering and pedogenesis in order to infer on paleoenvironmental and

paleoclimatic conditions at the time that the paleosols had formed (Sheldon and Tabor,

2009 and references therein).

The topics of paleosol research usually include soil morphology (Marković et al.,

2004), classification (Mack et al., 1993), geochemistry (Rye and Holland, 2000), rock

magnatism (Stine et al., 2020) and clay mineralogy (Presley et al., 2010; Daviel et al.,

2011). When it comes to the clay mineralogy of paleosols, there are generally four major components that should be considered: (a) detrital phyllosilicates from crystalline rocks,

(b) detrital phyllosilicates from a prior pedogenetic weathering, (c) pedogenic clay minerals formed through the authigenic in-situ processes, and (d) diagenetic clay minerals which came into existence as a result of sediment burial (Tabor and Myers,

2015). Detrital clays are dominant in paleosols formed under arid and cold climates,

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which simply do not permit an advanced hydrolysis (Southard and Miller, 1996;

Chamley, 1989). Detrital phyllosilicates are accounted for by muscovite, chlorite and

some minor vermiculite and kaolinite; however, it may be difficult to assess their

pedogenetic stages, due to weak or incipient chemical weathering under such climatic

conditions (Babechuk and Kamber, 2013). Detrital clays normally include resistant

phases such as quartz and less commonly feldspars, while the previously mentioned

chlorite readily converts into its hydroxyl-interlayered variety, (hydroxyl-interlayered minerals, Meunier, 2007) documented principally in alpine soils developed on a glacial till (e.g. Egli et al., 2001; Šegvić et al., 2018). The presence of chlorite is thus typically related to weak chemical weathering, unfavorable for advanced pedogenesis. Pedogenetic clay minerals however are typical in middle latitude regions where significant chemical weathering takes place, leading to the formation of a variety of brown soils (Gradusov,

1974). These soils generally constitute a rather thick blanket above the parent rocks.

Their pedogenetic clay minerals are consisted of “open” (i.e. exfoliated) primary phyllosilicates. Successive steps of clay mineral evolution proceed if hydrolysis progresses or becomes more active which leads toward the formation of irregular vermiculite mixed-layers, smectitic mixed-layers and finally a discrete (usually degraded) smectite (Chamley, 1989). Such a mineralogical series tends to develop from the base to the top of a given soil profile and from cool- to warm-temperature regions (Sheldon and

Tabor, 2009). Bisiallitization processes, which governs the transformation of primary minerals into pedogenetic clays, typically do not allow for the entire budget of Fe and Al to be incorporated in the octahedral positions of secondary phases which may result in the

Texas Tech University, Colton Mallett, May 2021 occurrence of sizeable amounts of soil Fe and Al oxyhydroxides (e.g., gibbsite and goethite; Egli et al., 2008).

Identification of paleosols can pose a significant challenge but is commonly based on the organization of morphological features within a soil layer (Sheldon and Tabor,

2009). In that sense, it is important to keep in mind that paleosols are solitary when formed during a period of landscape stability, but more commonly the paleosols are vertically stacked because of being formed in sedimentary systems undergoing net aggradation (e.g. Pace et al., 2009). In the case of minimal erosion, coupled with rapid and unsteady sedimentation, compound paleosols will generally form. Conversely, when the rate of pedogenesis exceed the rate of deposition, composite paleosols will ensue

(Dubiel and Hasiotis, 2011).

The Blackwater Draw Formation (BDF) of the Southern High Plains (SHP) consists of Pleistocene aeolian mantle, that contains a series of stacked paleosols and a surface soil that covers nearly the entire surface of the SHP (Holiday, 1988a; Holiday,

1989). These paleosols are thought to have developed under prevailing arid conditions, as suggested by the abundance of pedogenetic calcite and a general mineralogy devoid of clay mineral accumulation (e.g. Hovorka, 1997). The coarse-grained magnetite/maghemite populations were detected in the pre–mid-Pleistocene horizons while its fining-upward grain size was reported for the syn– and post–mid-Pleistocene paleosols which was attributed to increased rates of weathering (Stine et al., 2020).

Recent studies based on geochemistry and physical properties of BDF at the 29 m thick type section near Lubbock, TX (Fig. 2.1.1) have identified five buried paleosols and a surface soil (Baird et al., 2015), whereas the 13.9 m push core from Bushland Playa west

Texas Tech University, Colton Mallett, May 2021 of Amarillo, TX (Fig. 2.1.1; Fig. 2.3.2) revealed four buried soils and a surface soil

(Tewell, 2020).

In this study the paleosol architecture of the BDF was assessed using the distribution of smectite layers from pedogenetic smectite and interlayered illite-smectite

(Fig 4.3.4; Table 4.3.2), where a maximum of these phases is expected at the top of a given profile (e.g., Righi et al., 1999; Ryan and Huertas, 2009). This pedogenetic scheme consists of five paleosols and one surface soil in the core acquired from Bushland Playa

(Fig. 5.0). The pedogenetic scheme proposed herein is based on the total clay content of analyzed samples (Fig. 4.2.2.; Table 4.2.1), which was further refined by utilizing the clay speciation within the clay fraction (Fig. 4.3.4.; Table 4.3.2.).

While it is useful to use the distribution of clay minerals and clay mineral speciation to determine the intensity of chemical weathering, without higher resolution data, including grain-size distributions and geochemistry, a more complete pedogenic scheme cannot be constructed. These data truly only describe the intensity of chemical weathering using clay mineral authigenesis as a proxy. There are other processes which alter the concentration of authigenic clay minerals within the profile aside from those previously mentioned, most notably clay illuviation and mechanical clay infiltration.

These processes allow for the translocation of clay phases deeper into the profile. The effects of these processes are further convoluted by subsequent pedogenic and erosional events, which act to alter the original characteristics of a previously developed paleosol.

Therefore, the boundaries we identified are not necessarily boundaries between paleosols, rather they are boundaries between a section of least chemical weathering, and the section of highest chemical weathering located directly below (Fig. 5.0).

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Figure 5.0. Inferred pedogenic framework of the Bushland Playa Core. Zones displaying the highest intensity of weathering are at the top of each paleosol, and are denoted with a red arrow. Quartz content was assessed from XRD of the global fraction, clay mineral speciation was assessed using XRD of the clay fraction and elemental data were quantified using LA-ICP-MS.

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5.1.1. Surface Soil

Surface soil accounts for the first two samples, BL-1 and BL-2 (Fig. 5.0.), which corresponds to the interval between the surface and 0.7 meters below ground surface

(mbgs). Both samples have about 70 wt% of quartz and approximately 10 wt% of K- feldspar and albite; no calcite has been observed whatsoever (Fig. 4.2.2.). This is to be expected as calcium carbonate precipitates in a deeper soil horizon (Bretz and Horberg,

1949; James, 1972). The quantity of phyllosilicates in this top layer is around 10 wt%

(Table 4.2.1.), which is consistent with the composition of organic topsoil, since these soils are typically devoid of significant clay mineral accumulation (e.g. Gupta et al.,

2008). Clay fraction of this top layer is poor in smectite (˂10 wt%) and dominated in illite-smectite (~50 wt%), while the content of illite revolves around 30% (Fig. 4.3.4.).

Kaolinite content is around 5 wt% within the clay fraction. Both samples belong to the

category 1/group 1 clay mineral assemblage characterized with low-crystallinity smectite and several generations of poorly crystallized illite-smectite (Fig. 4.3.3.). An ongoing bisiallitization via several I-S species is documented herein with poorly crystallized discrete smectite rendering an end product of intensive chemical weathering (Pal et al.,

1989).

5.1.2. Paleosol I

The first paleosol layer is defined by its top portion at 1.45 mbgs (sample BL-3)

reaching 2.6 mbgs (sample BL-6) (Fig. 5.0.). An average quartz content here is somewhat

lower than in the case of surface soil being around 60 wt% (Table 4.2.1.). The content of

feldspar is equal to that of the surface soil (~10 wt%), while mica can attain up to 5 wt%

(Fig. 4.2.2.). Diagenetic calcite has been found only in the top layer of this paleosol

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horizon (Fig. 4.2.2.). The total clay content in this paleosol ranges from 10 to 20 wt%, the

highest being in the top sample (BL-3), which is to be expected (Table 4.2.1.). Clay fraction analyses further showed that smectite and smectite rich illite-smectite are dominant at the top of this paleosol (~35 wt%, Fig. 4.3.4.; Table 4.3.2.) thus testifying to the high rates of weathering in the exposed part of the profile (e.g., Šucha et al., 2001). In the deeper parts of the horizon, the content of smectite decreases while the fractions of illite and I-S both increase. The positive coorelation between the abundances of smectite and kaolinite in this horizon testifies to the high stages of hydrolysis (Karathanasis and

Hajek, 1983).

5.1.3. Paleosol II

The second paleosol is defined by the top at 3.5 mbgs (sample BL-7) and bottom at 6.5 mbgs (sample BL-9b) (Fig. 5.0.). In this paleosol the content of quartz is even more reduced (~55 wt%, Table 4.2.1.). On the other hand, the abundances of both plagioclase and K-feldspar are on par with those of paleosol I (Fig. 4.2.2.). The contents of 10Å phyllosilicates however, and that of discrete smectite and illite-smectite, are on average

30% higher than in the first paleosol (Table 4.2.1.), which likely posits that feldspars are

a main source of swelling clays and not detrital phyllosilicates. Indeed, feldspar readily

dissolves in eogenetic environment where a flow of meteoric water may lead to the

removal of excess cations such as alkalis and silica and then the formation of 2:1 clay

minerals (Glasmann, 1992; Bjørlykke, 1998, Worden and Burely, 2003). In the clay

fraction the content of discrete smectite and smectite-rich illite-smectite progressively

decreases down from the top of the paleosol giving lead to illite-smectite, which was found to be progressively enriched (Fig. 4.3.4.). Similar to paleosol I, a positive

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correlation between smectite and kaolinite abundances was also reported in this paleosol

profile (Table 4.3.2.).

5.1.4. Paleosol III

A third paleosol commences at 7.3 mbgs (sample BL-10) and ends at 8.9 mbgs

(sample BL-12) (Fig. 5.0.). This layer of ancient soil is analogues to the previous one in

terms of modal mineralogy, except for quartz, whose content is somewhat higher (~60

wt%, Table 4.2.1.), while the content of phyllosilicates decreases from 10 to 30 wt% (Fig.

4.2.2.). Clay fraction mineralogy reveals that the top layer of this paleosol is rich in

discrete smectite (~35 wt%, Table 4.3.2.), but steadily decreases toward the bottom of the

layer, adjusting for the increase of illite-smectite (Fig. 4.3.4.). Phases of the clay fraction in this paleosol belong to the groups 1 and 4 which are both characterized by discrete smectite of varying crystallinity with less amount of illite-smectite and detrital illite (Fig.

4.3.2.).

5.1.5. Paleosol IV

A fourth paleosol is defined by sample BL-13 at 9.55 mbgs to sample BL-15b at

11.6 mbgs (Fig. 5.0.). In comparison to the previous paleosol, quartz is somewhat enriched (~65 wt%, Table 4.2.1.), with lower albite (~8 wt%, Table 4.2.1.) and similar K- feldspar content (~10 wt%, Table 4.2.1.). Total phyllosilicates are higher compared to paleosol III and revolve around 15 wt% (Table 4.2.1.). Analysis of the clay fraction makes it apparent that the top of this paleosol is, as expected, rich in discrete smectite

(~25 wt%, Table 4.2.1.). Lower in the sequence, illite-smectite, as in all other present paleosol horizons, becomes a dominant clay mineral reflecting a reduced rate of pedogenesis.

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5.1.6. Paleosol V

The fifth paleosol commences at 12.4 mbgs (sample BL-16) and ends at 13.8

mbgs (sample BL-18). This basal paleosol has a strikingly different mineralogy compared

to the rest of analyzed sequence. First, the content of quartz is somewhat higher (~65-70

wt%, Table 4.2.1.), while the amount of diagenetic calcite may attain 30 wt% (Fig.

4.2.2.). The abundances of feldspars, both albite and K-feldspar, are significantly lower, centered around ~4 and ~7 wt%, respectively. Total phyllosilicate share is around 5% lower compared to other paleosol and averages around 10 wt%. Within the clay fraction one may observe a standard trend established earlier with smectite enrichment in the top layer, which in deeper portions of the paleosol is progressively diminishing because of illite-smectite increase. However, both smectite and illite-smectite in this layer appear as minor phases, with the major phase, detrital illite, attaining about 80% of clay fraction in the sample BL-18 (Fig. 4.3.4.). Further proof of the dominance of detrital phases in this paleosol is the presence of chlorite in samples BL-17 and BL-18, while kaolinite is barely half of the amount present in the rest of the section (Fig. 4.3.4.; Table 4.3.2.). Chlorite does not normally form as a pedogenetic phase, and as a detrital component, it is very susceptible to chemical weather (Ji et al., 2006); therefore, its presence in this paleosol is a clear sign of a reduced rate of weathering and colder climatic conditions during pedogenesis (Herbillon and Makumbi, 1975; Alekseev et al., 2003).

5.1.7. General Remarks

Observing the paleosol scheme presented herein, several inferences can be made.

First, discrete smectite and smectite-rich illite-smectite are found to be concentrated in the top sample of each paleosol thus representing the periods of extended pedogenesis.

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Such cyclicity is typical for the aeolian Quaternary sequences whose composition is

susceptible to climate oscillations and rates of sedimentations (e.g. Bucher et al., 2019;

Bruno et al., 2020). The abundances of kaolinite are also found to follow the dynamics of

smectite getting their maxima in the top portion of each paleosol (Fig. 4.3.4.), which

further corroborates intensive rates of weathering these top sequences of investigated

paleosols were exposed to. Kaolinite is indeed a pedogenetic phase that forms out of

dioctahedral smectite once the water activity in the soil increase (Liivamägi et al., 2018).

The BSE figure below (Fig. 5.1.) neatly depicts a well crystallized kaolinite in finer

sediment containing illite-smectite and discrete smectite, which suggests their genetic relation (Shen et al., 2012).

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Qtz Qz Or I-S

Kln

Figure 5.1. BSE image of sample BL-13 showing a typical clay paragenesis of the top portion of Paleosol IV. Abbreviations Or- orthoclase; Qz- quartz; Kln- kaolinite; I-S- illite-smectite. Mineral abbreviations after Kretz (1983) and Whitney and Evans (2010).

The clay mineral quantification also showed a lack of direct relation between

detrital 10Å phyllosilicate and diagenetic clays (Fig. 4.3.4.), which likely defines

feldspars as the main source of newly formed 2:1 clays through a series of illite-smectite intermediates. Saprolitic paleosol (paleosol V) is on the other hand dominated by presence of detrital phases such as illite and chlorite (Fig. 4.3.3., sample BL-17) which once the climate changes served as a feedstock of pedogenetic phyllosilicates (Chamley,

1989).

The transitional nature of the bulk of swelling clays in analyzed paleosols (i.e. illite-smectite) as suggested by X-ray diffractometry is further corroborated by their phase chemistry projected in multiple diagrams using different ratios of octahedral and

Texas Tech University, Colton Mallett, May 2021 interlayer cations (Mg, Fe, Al, and K; Fig. 5.2.). Moreover, the typical deficiency in K showed by illite-smectite as well as enrichment in silica is outlined in Figure 5.3.

Figure 5.2. Major oxide data plotted in comparison to the standard clay compositions to delineate clay compositions of illites (blue), illite-smectites (orange) and kaolinites (green). Modified from Pal et al. (2015) and Weaver and Pollard (1973).

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Figure 5.3. K-SiO2/Al2O3 plot depicting the relationship of interlayer K net loss with Al-Si replacement during smectitization. Illites (blue), illite-smectites (orange).

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5.2. Paleosols Geochemistry

Soil and paleosol profiles contain a variety of mineral components, the most

chemically and physically reactive being clay minerals. The clay-sized fraction in terrigenous deposits often includes, but not always, both authigenic and detrital minerals.

Clay-sized material is useful for the investigation of paleosols as it is the fraction that typically retains minor elements (i.e. REE, high-field-strength elements) after erosion and weathering of the source rock. These elements are either allocated in the interlayer positions of these clays or adsorbed on their surface, preserving the initial abundances inherited from the source material (McLennan, 1989). Thus, the investigation of the clay fraction is especially helpful to trace the origin of clay material in sediments and sedimentary rocks since REE abundances reflect the chemistry of their source (Cullers et al., 1997). On the other hand, major elements in soils are mainly controlled by the global mineralogy, and are especially useful for the interpretation of cemented horizons (e.g. calcite) or to assess the degree of weathering.

The geochemistry of the Black Water Draw Formation (BWDF) in the Southern

High Plains (SHP) was investigated through a 13.9 m core from Bushland Playa west of

Amarillo, TX. Eighteen samples analyzed by XRF and LA-ICP-MS show some degree of major and minor elements variability with a clear general trend throughout the core.

Using geochemistry alone, differentiating between different paleosol horizons was found to be difficult. While the use of geochemistry would not allow for a definitive assessment of paleosol architecture, it appears to have recorded the pre–mid- Pleistocene transition, well characterized on the North American Mid-Continent, towards the lower portion of the core (Mudelsee and Schulz, 1997; Head and Gibbard, 2005; Clark et al., 2006).

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5.2.1. Major elements signature

Calcite-rich horizons accounts for the most important variation in major elements.

Samples BL-3, 7, 8, 12, 16, 17 and 18 show high Ca values due to the presence of pedogenetic calcite (ranging from 10 to 30%, Table 4.2.1), resulting in a relative lower content in other major elements (Fig. 4.5.1.). A second noteworthy observation is the general decrease of all major elements with depth except for Si and Ca, especially in the bottom three samples, BL-16, 17 and 18. Additionally, both CIA and Al/Si values show a clear decrease with depth (Tewell, 2020). These observations coincide with the evolution of the mineralogy throughout the core; namely, in the lower portion of the profile, from

BL-15b to BL-18 (ca. 11.5 m to the end of the core) a relative enrichment in calcite and quartz was documented accompanied by the depletion of feldspars and clay minerals.

Major elements appear to be richer in the interval from sample BL-8 to BL-12, corresponding to the paleosols II and III. These two paleosols are featured by the highest clay minerals content, as well as high Al/Si values.

5.2.2. Minor elements signature

Both observations made using major elements distribution seem to be observed, and are even more pronounced, using Rare Earth Elements (REE) and other trace elements. Most of them are shown to be relatively depleted in calcite-rich horizons (Fig.

4.5.2; Fig. 4.6.1), and up to 12 times less concentrated in the lower portion of the core

(samples BL-16, 17 and 18, Paleosol V).

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5.2.3. General Remarks

Geochemical data from this study indicates that Paleosol V is distinct from the other soils in the Bushland Core. Based on the results, we are not able to precisely characterize individual, clay mineralogy based paleosols, mainly due to the paucity of

analyzed samples. However, the presence of a recorded pre–mid- Pleistocene transition is suggested, dividing the core into an upper and lower member of the BDF. This division is emphasized by both major and minor elements negative excursions with depth, indicating a weakly developed soil in the lower part of the core (Paleosol V) formed in arid environment. This nicely aligns with the mineralogical results as evidenced by intensive and well-expressed pedogenic carbonate content in Paleosol V as well as a lack of clay

authigenesis. In this regard the present study corroborates the findings of Tewell, (2020)

and Stine et al. (2020).

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5.3. Paleosols and Inferences on Quaternary Climate Across the Southern High Plains

Paleosol deposits which may occur worldwide as vast belts across the mid-

latitudes are critical terrestrial archives for reconstructing glacial to interglacial climate

changes during the Quaternary (Muhs et al., 2003; Li et al., 2018). Numerous studies

have shown that there is a relationship between clay mineral assemblages on one hand

and climatic conditions and weathering intensity in the source region on the other hand

(Biscaye, 1965; Thiry, 2000). Bearing in mind the clay minerals may have various origins

in natural environment (Weaver, 1956; Park and Khim, 1992) a caution should be

exercised when it comes to their usage as proxies for paleoclimate and weathering

intensity as only the authigenic clays found in soil serve that purpose (Liu et al., 2018).

The formation of clay minerals in soil profile is likely controlled by the climate;

namely the climate changing from dry and cold to hot follows a general trend of

bisiallitization which ultimately leads to Al-oxide formation

(smectite + illite + kaolinite → smectite + kaolinite → kaolinite + gibbsite → gibbsite, (Du

et al., 2012). Smectite commonly forms abundantly in low reliefs where poor drainage

prevents the removal of silica and alkalis. This is related to a warm climate with

alternating humid and dry seasons (Chamley, 1997). On the other hand, kaolinite is

normally a product of highly hydrolytic alteration environment and therefore forms in

perennially warm and humid climates with a least temperature of ~15 °C (Righi et al.,

1999). Relative proportions of those two pedogenetic minerals can be used as a proxy for

climate with an elevated kaolinite/smectite ratio (Table 4.3.2.) being indicative of humid

and warm to more dry and seasonal climate variations (Zhang et al., 2016). Earlier work has shown that the geochemistry of clay fraction shows more susceptibility to climate

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than that of the global sample (Zhang et al., 2015), which necessitates a careful

examination of diffraction patterns of the clay fraction to inspect on the clay mineral

compositional variations (Moore and Reynolds, 1997).

The eolian Blackwater Draw Formation (BDF) of the Southern High Plains of

New Mexico and Texas investigated in this study was deposited between 1.6 Ma and the

Holocene which corresponds to the Early to Middle Pleistocene (Gustavson and Holliday,

1999). Within the formation, the intervals of sediment aggradation were separated by

times of landscape stability which further permitted regionally extensive pedogenesis

(Holliday, 1989, 1990). Considering the lifespan of BDF it is reasonable to believe it had

recorded the major mid-Pleistocene transition characterized by a global increase in ice

volume and a strong influence by the change from a 41 k.y. obliquity cycle to a 100 k.y.

eccentricity cycle that happened between ~1.2 and 0.7 Ma (Head and Gibbard, 2005;

Clark et al., 2006). The onset of the mid-Pleistocene transition at the northern hemisphere

is followed by the cooler and drier climate that was conditioned by the high atmospheric

pressures (Head and Gibbard, 2005). Investigating the BDF at the type section near

Lubbock, TX (Stine et al., 2020) demonstrated using the sedimentological, geochemical and rock-magnetic data that the time following the mid-Pleistocene transition saw different sediment provenance during glacial intervals and a higher degree of weathering during interglacial intervals.

Clay mineralogy of analyzed core has indicated an abrupt change in the rate of weathering and thus pedogenesis at the transition of paleosols IV and V (Fig. 5.0.). While the major phyllosilicates in the latter are detrital illite (and chlorite) the former is marked by diagenetic swelling phases such as discrete smectite and smectite rich illite-smectite

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(Fig. 4.3.4.). Although there is a lack of age control in the studied core this major change

in weathering can be assigned to the previously mentioned mid-Pleistocene transition

(Clark et al., 2006). Abundance of diagenetic calcite in that zone further testifies on

weakly developed soils in an arid environment which is consistent with the finding of

Stine et al., (2020). On the other hand, the syn- and post-mid Pleistocene transition intervals are marked by the formation of regular soil cycles characterized with the increased hydrolysis under warmer and more humid climate which is also in line with the interpretation of the paleosols sequence at the type section near Lubbock, TX (Stine et al., 2020). The climate must have been warm during interglacials with efficient runoffs as suggested by kaolinite enrichment at the top sample of each paleosol (Dura et al., 2009).

The same may be inferred from EDS chemistry of pedogenetic clays where an increase in

Mg/K suggests a higher share of smectitic component (Wilson, 2013; Table 4.4.3.). A good example would be BL-7 which originates from the top of paleosol II has Mg/K 1.4 times more compared to BL-14 which define the mid-section of paleosol IV.

This research has demonstrated that clay mineral association can be used as reliable paleoclimatic proxies once possible diagenetic and authigenic bases have been identified and excluded (Bertier et al., 2008; Pace et al., 2009). Combined with other climate-susceptible proxies (geochemistry, grain size, stable isotopes) the use of clay mineral assemblages, concentrations and degree of crystallization has been shown effective for reconstructing regional paleoclimatic changes (Singer, 1980; Crawford

Elliot et al., 1997; Wang et al., 2020).

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5.4. Implications of K-Ar Ages of Detrital Illite and Whole-Rock Trace Element Chemistry

According to the results presented herein, the Bushland core of the Blackwater

Draw Formation can be divided in two main sections at about 12 mbgs, as suggested by

fundamental differences in both the global and clay mineralogy, as well as geochemical

characteristics (Fig. 5.0). Previous studies have shown a similar division, clearly

identifying an upper and lower part of the BDF (Baird, 2015, Tewell, 2020, Stine et al.,

2020) emphasized by the differences in grain-size distribution, magnetic minerals as well

as geochemical trends. Whilst it was previously concluded that BDF sediment was

derived solely from the Pecos River drainage area, the clear geochemical and lithological

changes recently reported in the BDF have been attributed to a different sediment source

provenance between the lower and upper section of the studied type sections. Using

sedimentological, geochemical, and rock-magnetic data, Stine et al. (2020) found the

lower section of the BDF to be consistent with a sediment source derived from the

neighboring Pecos River. However, those authors proposed a mixture of southern and

northern sources during the deposition of the upper section of the core (i.e. after the mid-

Pleistocene transition) consisting of an eolian-suspension of silt and clay material. The similar study of Tewell (2020) corroborates the finding of Stine et al. (2020), through the comparison of relatively immobile major and trace elements from the upper and lower sections of the Bushland Core to potential source areas described in the literature, in addition to new data from the Pecos River. The study of clay mineralogy outlined in this thesis concur with the aforementioned studies, showing a dominance of detrital and sandier material in the lower section of the core overlain by a lithology richer in silt- and clay-sized sediments.

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While the difference in provenance between the lower and upper sections of the

BDF has been unequivocally confirmed, the exact provenance, especially of the upper section, still lacks scientific consensus. In order to identify the exact sediment source terrains of the Bushland core, a thorough comparison with the geochemical fingerprint of potential provenance regions is crucial. However, due to the nature of this study and the sampling strategy, there was a limit to the number of samples that could be analyzed.

Hence, the geochemical data collected from the core is insufficient to undertake the task of assessing different source provenance via a detailed fingerprinting work. Yet, with the available minor element data, normalization diagrams were constructed to infer on the differences between the two sections of the core.

As previously mentioned, REE abundances in sediment reflect chemistry of their source (McLennan, 1989) and are therefore useful for provenance studies (Nesbitt et al.,

1990; Cullers et al., 1997). Figure 4.5.1 displays REE concentrations obtained through the laser ablation mass spectrometry normalized to chondrite (Boynton, 1984) as well as

a spider diagram with trace elements normalized to the primitive mantle (Hofmann,

1988). All normalized REE curves show smooth parallel trends with concentration levels

in the range of 4 to about 100x chondrite. Light REEs (LREEs) are enriched relative to

heavy REEs (HREEs), of about 3x. Both LREEs and HREEs show a systematic

magmatic fractionation pattern, whereas consistently higher LREE concentrations might

reflect source enrichment. Additionally, a clear negative Eu anomaly, characteristic of

intermediate to felsic magmas, serves as additional evidence of magmatic fractionation

(Rollinson, 1993). However, although most paleosols show similar concentrations,

Paleosol V curves are of much lower REE abundances. This has already been discussed

Texas Tech University, Colton Mallett, May 2021 and is explained by the difference in clay content in the deepest paleosol profile. Trace elements trends from spider diagrams (Fig. 5.4.1) tend to be mutually similar, except for

Ba and Sr in sample BL-18. The latter is known to be very rich in calcite, which is likely the cause of Ba and Sr enrichment in the sample.

A comparison of different paleosols using normalization diagrams to discriminate between different sources is proven to be inconclusive but show a clear intermediate to felsic signature for the material in both the upper and lower sections of the Bushland core.

Although REE and trace elements data collected are insufficient to assess the different provenances suggested from our mineralogical results and previous studies, K-

Ar ages acquired from three different samples (i.e. BL-2, Paleosol I; BL-10, Paleosol III;

BL-17, Paleosol V) provide further insight in distinguishing the upper and lower sections

(Table 4.7.1). Two size fractions, 2-1 and <1 µm have been separated from each of the three selected samples and are believed to represent mainly detrital and authigenic material in the coarser and finer fraction, respectively. Sample BL-17 from Paleosol V shows a detrital age of about 124 Ma, while the fraction believed to be mainly authigenic

(<1 µm) is dated to about 120 Ma. Both ages are surprisingly close to each other, proving almost no clay neoformation during hydrolysis from the time of deposition to present, characteristic of the pre-mid-Pleistocene transition section of the core, as reported by previous authors (Holliday, 1989; Holliday, 1990; Baird, 2015; Stine et al., 2020). The material from Paleosol V appears to originate from a terrane of Cretaceous age. Samples

BL-10 and BL-2, showing detrital ages of about 145 and 169 Ma respectively, are part of the upper section of the Bushland core and are believed to have a mixed provenance

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(Stine et al., 2020). Both samples show ages approximately 30 to 50 Ma younger in their authigenic fraction. Since both dated fractions are unlikely to represent 100% of detrital or authigenic material, it may be hypothesized that detrital ages are likely older than reported while authigenic ages tend to be younger than reported (Schomberg et al., 2019).

Accordingly, detrital ages could potentially be as old as 190-200 Ma if one accounts for a

100% detrital illite fraction. The IAA would typically require a precise quantification of illite polytypes, especially detrital illite (2M1), however, this study did not quantify illite polytypes, therefore the detrital and authigenic illite proportions were estimated following the findings of Zanoni (2020, PhD thesis to be submitted). The upper section of the core appears to be characterized by Jurassic ages, drastically different from the material constituting the lower portion of the core. As previously outlined, Stine et al.

(2020) found that post–mid-Pleistocene transition paleosols (i.e., Paleosols I, II, III and

IV in this study), most closely, but imperfectly align with a Nebraska loess provenance.

They concluded that the upper section of the BDF is consistent with a mixed provenance of Nebraska loess and sand derived from the Pecos River drainage. Our ages show that the upper section of the core is different from the lower section thus confirming a shift in provenance through the Mid-Pleistocene transition.

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Chondrite- normalized REE A 1000 BL1 BL2 BL3 BL4 BL5 BL6 100 BL7 BL8 BL9 BL10

Sample/ Chondrite Sample/ BL11 10 BL12 BL13 BL14 BL15 BL16 1 BL17 La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu BL18 Primitive Mantle- normalized trace elements B 1000 BL1 BL2 BL3 BL4 BL5 BL6 100 BL7 BL8 BL9 BL10 BL11 10 BL12

Sample/ primitive mantle primitive Sample/ BL13 BL14 BL15 BL16 BL17 1 Rb Ba Th U Nb Ta K La Ce Pr Sr Nd Zr Hf Sm Eu Ti Gd Dy Y Er Yb Lu BL18 Figure 5.4.1 (A) REE concentrations obtained through the laser ablation mass spectrometry normalized to chondrite After Boynton 1984. (B) Spider diagram with trace elements normalized to the primitive mantle Hofmann (1988)

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5.5. Final Remarks

The results of this research, which builds a pedogenic framework based on the

concentration of authigenic clay minerals within the core profile, should be used in

conjunction with geochemistry and grain-size distribution data to build a more robust

description of the paleosols present at this location. Tewell (2020) uses geochemistry and grain-size distribution of the Bushland core to describe a series of four paleosols, 2 of which were recognized as compound soils. The integration of these data is the next step in resolving a more precise pedogenic framework, including the true number of paleosols and weathering intensity during pedogenesis.

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CHAPTER 6

CONCLUSIONS

The Blackwater Draw Formation is composed of a series of stacked paleosols, deposited since the early to mid-Pleistocene, which preserve a record of prevailing paleoclimate conditions since the time of deposition. The geochemistry, whole rock mineralogy and clay mineral assemblages of the Blackwater Draw Formation at Bushland are indicative of varying climate regimes and multiple sediment sources thus corroborating the findings of previous studies.

• Due to the complexities of illuviation and translocation of clay minerals during

pedogenesis within a profile, the unique complications at this locality, and the

lack of high resolution geochemical and grain-size distribution data, the

“Paleosols” we identify are realistically simply zones denoted by the intensity of

weathering.

• The geochemistry, mineralogy, and clay mineral speciation of the analyzed

samples indicate that the Blackwater Draw Formation at Bushland consists of five

buried paleosols and a surface soil, as opposed to the four paleosols, two of which

were recognized as compound soils, and the surface soil described by Tewell

(2020). These data also suggest the existence of an upper member, composed of

the surface soil down to Paleosol IV, and a lower member, composed of Paleosol

• Each paleosol has been distinguished using the distribution of smectite and

interlayered illite-smectite (Figure 5.0) with their concentration maxima standing

as a top section of given profile.

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• In Paleosol V, the general deprivation of minor and some major elements with

depth, extensive pedogenic calcite accumulation and lack of clay mineral

authigenesis indicate that the lower member of the Blackwater Draw Formation at

Bushland was subject to more arid condition, compared to the upper member

which on the contrary does contain extensive authigenic clay minerals, indicating

a warmer and wetter climate.

• K-Ar dating of detrital and authigenic illite clearly differentiate the upper and

lower section of the Bushland Core. Not only is the age of detrital illite in the

upper section significantly older than the detrital illite in the lower section, but the

age difference between the authigenic and detrital illite fraction is much greater in

the upper section than in the lower section. The difference in age between the

authigenic and detrital illite fraction demonstrates that the upper section was

subject to climate conditions favorable to clay neoformation, while the lower

section was not. The change in age of the detrital fraction between the upper and

lower section is indicative of a provenance shift.

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APPENDICES

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APPENDIX A. X-RAY FLUORESCENCE

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XRF Data for SiO2, Al2O3, CaO, Fe2O3, MgO, K2O, Na2O, TiO2, P2O5 reported in weight %

Sample SiO2 Al2O3 CaO Fe2O3 MgO K2O Na2O TiO2 P2O5 MnO LOI Sum BL-1 72.866 13.050 0.866 4.650 1.251 2.394 0.761 0.713 0.088 0.071 9.186 105.895 BL-2 74.642 10.934 2.806 3.629 1.164 2.126 0.711 0.624 0.065 0.047 8.673 105.421 BL-3 66.133 8.813 17.048 2.118 1.270 1.833 0.673 0.444 0.097 0.033 15.473 113.935 BL-4 75.618 10.590 1.732 3.451 1.184 2.321 0.898 0.635 0.092 0.069 6.868 103.458 BL-5 74.932 10.841 1.892 3.579 1.202 2.332 0.898 0.653 0.082 0.069 6.703 103.183 BL-6 74.564 10.790 2.216 3.544 1.241 2.261 0.869 0.637 0.093 0.059 7.336 103.612 BL-7 68.160 10.085 11.762 2.656 1.370 2.196 0.844 0.484 0.152 0.038 12.472 110.217 BL-8 69.438 11.079 7.716 3.039 1.376 2.438 0.973 0.555 0.129 0.046 10.326 107.115 BL-9 69.377 13.500 3.546 3.971 1.886 2.666 0.896 0.563 0.095 0.046 9.071 105.618 BL-10 77.728 9.661 1.734 2.938 1.283 2.088 0.703 0.405 0.068 0.048 5.951 102.607 BL-11 72.445 12.777 0.873 4.349 1.783 2.672 0.795 0.591 0.082 0.053 7.161 103.581 BL-12 70.358 10.284 9.574 2.663 1.332 2.182 0.659 0.447 0.093 0.030 12.267 109.889 BL-13 76.340 10.664 1.407 3.056 1.270 2.164 0.780 0.512 0.051 0.051 6.153 102.450 BL-14 76.590 9.810 2.540 2.871 1.256 2.172 0.701 0.486 0.101 0.038 6.444 103.008 BL-15 73.239 10.019 5.861 2.881 1.307 2.172 0.693 0.475 0.164 0.038 8.787 105.636 BL-16 74.783 6.350 12.105 1.586 1.088 1.457 0.190 0.273 0.174 0.014 11.953 109.971 BL-17 77.629 5.680 10.131 1.420 1.064 1.354 0.197 0.268 0.183 0.016 10.067 108.008 BL-18 78.753 4.097 11.327 0.831 1.256 1.025 0.095 0.172 0.032 0.009 10.802 108.397 AGV-120 58.2356 17.2387 4.7806 6.9393 1.5548 2.881 4.0347 1.0653 0.4932 0.0958 97.319 AGV-120 58.2711 17.2196 4.7861 6.9379 1.5596 2.8865 4.033 1.0636 0.4927 0.0958 97.3459 GSP-1 66.0114 15.06 1.8334 4.2595 1.0492 5.4568 2.6165 0.6422 0.2773 0.0376 97.2439 GSP-1 66.0694 15.0504 1.8332 4.2476 1.0456 5.4587 2.6185 0.6457 0.2783 0.0371 97.2845 BHVO-1 48.9961 13.3921 11.3501 11.6368 6.8017 0.5017 1.9525 2.733 0.2739 0.1659 97.8038 BHVO-1 49.0058 13.397 11.3493 11.6232 6.8106 0.5006 1.9564 2.7341 0.2713 0.1643 97.8126

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Normalized XRF Data for SiO2, Al2O3, CaO, Fe2O3, MgO, K2O, Na2O, TiO2, P2O5 reported in weight %

Sample Depth SiO2 Al2O3 CaO Fe2O3 MgO K2O Na2O TiO2 P2O5 MnO LOI Sum BL-1 0.3 68.810 12.323 0.818 4.391 1.181 2.260 0.719 0.673 0.083 0.067 8.675 100.0 BL-2 0.7 70.804 10.372 2.661 3.442 1.105 2.017 0.674 0.592 0.062 0.044 8.227 100.0 BL-3 1.45 58.044 7.735 14.962 1.859 1.115 1.609 0.590 0.390 0.085 0.029 13.581 100.0 BL-4 1.7 73.091 10.236 1.674 3.336 1.145 2.243 0.868 0.613 0.089 0.067 6.638 100.0 BL-5 2.4 72.621 10.507 1.833 3.468 1.165 2.260 0.871 0.633 0.079 0.067 6.496 100.0 BL-6 2.7 71.965 10.414 2.139 3.420 1.198 2.182 0.839 0.614 0.090 0.057 7.080 100.0 BL-7 3.5 61.841 9.150 10.672 2.410 1.243 1.992 0.766 0.439 0.138 0.034 11.315 100.0 BL-8 4.2 64.826 10.343 7.203 2.837 1.284 2.276 0.908 0.519 0.120 0.043 9.640 100.0 BL-9 5.9 65.687 12.782 3.357 3.760 1.786 2.524 0.849 0.533 0.090 0.044 8.589 100.0 BL-10 7.3 75.753 9.416 1.690 2.863 1.250 2.035 0.685 0.395 0.066 0.047 5.800 100.0 BL-11 8.35 69.940 12.335 0.843 4.199 1.722 2.580 0.767 0.571 0.079 0.051 6.913 100.0 BL-12 8.9 64.027 9.358 8.712 2.423 1.212 1.986 0.600 0.407 0.085 0.027 11.164 100.0 BL-13 9.55 74.514 10.409 1.374 2.983 1.240 2.112 0.761 0.500 0.050 0.050 6.006 100.0 BL-14 10.2 74.353 9.524 2.466 2.787 1.219 2.109 0.680 0.472 0.098 0.037 6.256 100.0 BL-15 11.1 69.332 9.485 5.548 2.728 1.238 2.056 0.656 0.450 0.155 0.036 8.318 100.0 BL-16 12.4 68.002 5.774 11.007 1.442 0.989 1.325 0.173 0.248 0.158 0.013 10.869 100.0 BL-17 13.1 71.873 5.258 9.379 1.315 0.985 1.254 0.182 0.248 0.170 0.015 9.321 100.0 BL-18 13.8 72.653 3.780 10.449 0.766 1.159 0.946 0.087 0.158 0.029 0.008 9.965 100.0 AGV-120 58.235 17.2387 4.7806 6.9393 1.5548 2.881 4.0347 1.0653 0.4932 0.0958 97.32 AGV-120 58.271 17.2196 4.7861 6.9379 1.5596 2.8865 4.033 1.0636 0.4927 0.0958 97.35 GSP-1 66.011 15.06 1.8334 4.2595 1.0492 5.4568 2.6165 0.6422 0.2773 0.0376 97.24 GSP-1 66.069 15.0504 1.8332 4.2476 1.0456 5.4587 2.6185 0.6457 0.2783 0.0371 97.28 BHVO-1 48.9961 13.3921 11.3501 11.6368 6.8017 0.5017 1.9525 2.733 0.2739 0.1659 97.80 BHVO-1 49.0058 13.397 11.3493 11.6232 6.8106 0.5006 1.9564 2.7341 0.2713 0.1643 97.81

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XRF Data for Ba, Co, Cr, Cu, Ni, Sr, Zn and Zr reported in PPM Sample Ba Co Cr Cu Ni Sc Sr Zn Zr BL-1 615.500 6.800 51.200 23.000 28.300 5.500 112.100 73.700 438.500 BL-2 564.400 3.600 40.900 20.300 22.700 5.100 98.300 51.200 491.100 BL-3 1827.7 1.700 17.400 6.600 12.000 16.400 172.100 29.200 319.000 BL-4 557.500 3.300 38.700 14.800 19.100 4.900 113.700 52.200 457.500 BL-5 545.600 3.500 40.400 13.500 19.200 4.900 113.200 58.100 472.200 BL-6 559.100 3.300 38.100 9.700 18.800 5.300 112.000 53.300 472.700 BL-7 579.000 1.400 26.700 7.400 13.700 13.200 172.400 37.200 319.500 BL-8 585.600 1.800 31.400 9.900 16.800 10.900 183.200 46.500 327.800 BL-9 516.100 3.900 42.000 11.100 23.700 6.700 156.800 64.300 335.500 BL-10 482.700 2.400 29.200 12.800 17.500 3.500 98.500 49.200 300.800 BL-11 504.900 5.700 44.000 19.800 25.300 3.400 116.300 70.300 348.200 BL-12 509.900 1.500 27.600 9.000 13.700 10.400 137.500 40.900 320.800 BL-13 471.200 2.600 32.300 10.400 20.900 3.100 111.200 48.300 404.900 BL-14 476.900 2.300 30.400 11.900 17.400 4.700 98.700 49.800 390.200 BL-15 507.900 1.800 29.200 10.000 16.500 6.900 120.500 44.900 337.900 BL-16 395.100 1.800 16.800 3.900 8.100 11.500 119.900 24.900 183.400 BL-17 377.600 1.800 17.000 2.600 6.800 10.100 115.200 20.100 223.000 BL-18 299.700 2.900 11.900 5.400 3.900 8.100 116.300 11.100 176.400 AGV-120 1154.2 14.3 11.8 41.4 11.1 11.7 650.9 79.8 249.6 AGV-120 1152 14.5 11.8 41.8 10.8 12 650.8 80.6 250.5 GSP-1 1276.2 4.1 8.2 29.7 13.2 3.7 216.2 93.9 515.8 GSP-1 1288.7 4.1 8.6 28.2 12.5 4 217.3 94.9 515.7 BHVO-1 146.3 51 283.9 114.3 100.2 36.6 369.6 87.7 159.6 BHVO-1 156.5 50.1 280.6 115.3 100 33.5 369.4 87.3 159

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APPENDIX B. LASER ABLATION-INDUCTIVELY COUPLED-PLASMA MASS SPECTROMETRY

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Reduced LA-ICP-MS trace element abundances for Sc, Ti, V, Ga, Rb, Sr, Y, Zr, Nb, Ba and La reported in PPM Sample Sc Ti V Ga Ga Rb Sr Y Zr Nb Ba La BL-1 13.07 4735.23 108.21 42.61 15.47 96.53 116.20 39.62 586.81 18.15 545.03 42.66 BL-2 11.36 4159.93 91.85 37.41 13.10 81.47 114.34 37.17 697.47 15.84 486.54 37.47 BL-3 9.29 3433.91 82.56 95.91 10.33 63.00 256.76 32.21 576.37 13.07 1551.30 34.42 BL-4 11.28 4200.34 93.23 36.17 12.42 83.78 129.52 35.77 637.92 16.46 476.82 36.43 BL-5 11.43 4254.55 90.72 33.96 12.60 83.39 126.49 34.49 658.96 17.09 452.83 36.11 BL-6 11.20 4040.33 90.88 34.93 12.34 79.66 122.57 36.83 670.34 16.00 463.05 36.45 BL-7 9.96 3416.06 86.50 34.47 11.69 70.95 229.68 32.00 521.28 14.98 477.25 34.75 BL-8 10.38 3660.95 92.10 36.17 12.77 81.58 225.20 34.80 497.57 18.20 486.98 38.01 BL-9 11.18 3500.53 94.44 36.65 16.65 99.56 167.71 32.18 462.72 26.53 437.80 36.26 BL-10 8.90 2540.23 78.60 28.91 11.65 78.20 108.62 23.83 402.32 15.02 384.63 27.02 BL-11 11.67 3610.07 107.61 35.03 14.93 97.01 118.47 27.93 460.68 20.14 449.06 32.78 BL-12 9.20 3035.35 81.72 31.56 12.24 75.18 177.15 31.80 488.44 19.85 418.61 32.26 BL-13 9.12 3089.73 74.69 29.04 12.76 77.05 115.26 28.99 529.84 23.39 375.72 29.61 BL-14 8.75 2997.33 74.45 28.49 11.51 76.93 111.09 29.80 521.64 18.83 387.61 29.38 BL-15 9.11 2996.22 80.41 30.69 11.79 78.46 144.96 29.13 484.68 18.16 421.11 31.64 BL-16 6.36 1876.90 57.82 21.27 7.39 50.88 163.87 16.33 290.34 8.46 298.64 19.64 BL-17 6.02 1832.52 53.37 19.93 6.32 45.25 153.65 15.95 340.12 7.59 291.27 18.24 BL-18 4.70 1346.30 39.25 16.30 4.94 35.83 160.53 9.33 264.05 4.73 240.15 12.12 BHVO-1 31.86 15600.67 385.38 26.15 20.78 9.13 452.02 24.93 179.67 17.59 132.43 15.37 BHVO-1 25.95 12844.18 316.02 22.11 18.87 7.61 389.26 20.56 149.13 14.96 107.58 12.51 BHVO-1 26.76 13290.99 323.10 20.86 18.31 7.42 379.89 20.71 144.23 14.24 108.02 12.63 BHVO-1 27.23 12740.45 314.21 20.55 17.28 7.22 372.74 20.98 149.65 14.15 102.40 12.77

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Reduced LA-ICP-MS trace element abundances for Ce, Pr, Nd, Sm, Eu, Gd, Tb, Dy, Tb, Dy, Er, Yb, Lu, Hf, Ta, Pb, Th and U (PPM) Sample Ce Pr Nd Sm Eu Gd Tb Dy Er Yb Lu Hf Ta Pb Th U BL-1 81.74 9.52 36.80 7.00 1.29 6.57 0.99 6.34 3.89 3.90 0.63 13.35 1.27 16.17 14.33 2.85 BL-2 73.70 8.56 33.22 6.37 1.21 5.90 0.93 5.70 3.65 3.80 0.60 15.54 1.11 14.10 12.35 2.76 BL-3 66.02 7.72 28.94 5.57 1.07 5.25 0.78 5.10 3.14 3.32 0.52 12.93 0.91 9.36 10.35 2.65 BL-4 73.60 8.42 32.64 6.21 1.18 6.14 0.88 5.79 3.58 3.76 0.59 14.30 1.15 12.66 12.56 2.65 BL-5 74.48 8.24 31.92 6.21 1.13 5.68 0.87 5.54 3.47 3.59 0.57 15.13 1.18 11.90 12.77 2.73 BL-6 72.07 8.38 32.46 6.36 1.32 6.18 0.89 5.98 3.68 3.88 0.60 15.32 1.13 10.59 12.15 2.76 BL-7 64.63 7.85 29.48 5.86 1.08 5.12 0.77 4.95 3.20 3.32 0.53 11.98 1.02 10.21 11.10 2.62 BL-8 72.20 8.61 31.90 6.45 1.19 5.67 0.85 5.52 3.41 3.67 0.57 11.42 1.22 11.51 12.12 2.77 BL-9 78.44 8.25 30.68 5.97 1.01 5.26 0.84 5.34 3.23 3.30 0.52 10.85 1.77 13.37 14.14 2.66 BL-10 59.56 6.03 22.66 4.45 0.79 4.27 0.62 3.94 2.47 2.54 0.41 9.46 1.13 14.62 10.40 2.12 BL-11 70.39 7.71 29.04 5.55 0.97 5.08 0.77 4.77 2.87 3.12 0.48 10.90 1.45 16.67 13.37 2.75 BL-12 63.18 7.42 28.18 5.55 0.96 5.17 0.78 5.09 3.16 3.51 0.51 11.52 1.36 10.79 11.84 2.54 BL-13 62.40 7.11 26.62 5.48 0.86 4.96 0.75 4.95 3.01 3.34 0.50 12.43 1.63 11.10 11.99 2.38 BL-14 58.11 6.70 25.55 5.02 0.85 4.79 0.71 4.64 2.97 3.21 0.51 12.05 1.27 12.02 10.38 2.41 BL-15 61.62 7.24 27.36 5.23 0.93 4.70 0.72 4.58 2.85 2.99 0.48 11.27 1.23 10.75 10.93 3.13 BL-16 36.61 4.49 16.45 3.24 0.56 2.71 0.43 2.61 1.68 1.75 0.27 6.66 0.57 6.16 6.27 1.80 BL-17 32.40 4.13 15.42 2.95 0.56 2.58 0.40 2.49 1.56 1.72 0.27 7.90 0.53 5.37 5.51 1.95 BL-18 22.00 2.69 9.96 1.75 0.36 1.53 0.24 1.49 0.97 1.16 0.18 6.14 0.33 5.37 3.95 1.18 BHVO-1 37.55 5.27 24.34 6.01 2.05 5.83 0.88 4.99 2.35 1.96 0.28 4.19 1.08 2.02 1.27 0.47 BHVO-2 31.31 4.25 19.26 4.52 1.67 4.95 0.72 4.18 2.00 1.55 0.22 3.27 0.81 bdl 1.00 0.58 BHVO-3 31.14 4.43 19.86 4.77 1.59 4.98 0.71 3.92 1.83 1.70 0.18 3.28 0.83 bdl 0.84 bdl BHVO-4 30.92 4.55 21.25 4.76 1.57 4.56 0.75 4.14 1.99 1.75 0.23 3.61 0.92 bdl 1.03 0.33

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APPENDIX C. ACCURACY AND PRECISION OF XRF AND LA-ICP-MS ANALYSES

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Known composition of in-house standards in weight %

Sample SiO2 Al2O3 CaO Fe2O3 MgO K2O Na2O TiO2 P2O5 MnO

AGV-120 58.24 16.72 16.72 7.09 1.77 2.96 4.87 1.1 0.493 0.112 GSP-1 66.01 14.8 2.28 4.52 1.16 5.68 3.2 0.698 .278 0.045 BHVO-1 49 12.4 11.95 12.11 7.63 0.5 2.43 2.72 0.251 0.182

Percent deviation of experimental analytes from in-house standards

Sample SiO2 Al2O3 CaO Fe2O3 MgO K2O Na2O TiO2 P2O5 MnO

AGV-120 0.02 3.05 -13.03 -2.14 -12.02 -2.58 -17.17 -3.23 -.001 -14.46 GSP-1 0.05 1.72 -19.59 -5.89 -9.71 -3.91 -18.20 -7.74 -0.07 -17.00 BHVO-1 00 8.02 -5.02 -3.96 -10.80 0.23 -19.57 0.50 8.61 -9.29 Average 0.02 4.26 -12.55 -4.00 -10.84 -2.09 -18.31 -3.49 2.84 -13.58

Relative standard deviation of analyes from standards

Sample SiO2 Al2O3 CaO Fe2O3 MgO K2O Na2O TiO2 P2O5 MnO AGV- 0.0304 0.0798 0.0554 0.0100 0 0.1541 0.0574 0.0210 0.0953 0.0507 120 GSP-1 0.0439 0.2717 0.0318 0.1398 0.6693 0.1718 0.0054 0.0382 0.0174 0.1799 BHVO-1 0.0099 0.0201 0.0182 0.0584 0.4845 0.0653 0.0035 0.0997 0.1097 0.4768 Average 0.0280 0.1239 0.0352 0.0694 0.3846 0.1304 0.0221 0.0530 0.0741 0.2358

Relative standard deviation of analyes from standards

Sample Ba Co Cr Cu Ni Sc Sr V Zn Zr AGV- 0.0954 0.69444 0 0.48077 1.36986 1.26582 0.00768 -14.2 0.49875 0.17996 120 GSP-1 0.4873 0 2.3809 2.59067 2.72374 3.8961 0.25375 0 0.52966 0.00969 BHVO- 3.368 0.8902 0.5845 0.4355 0.0999 4.4222 0.0270 0 0.2285 0.1883 1 Average 1.317 0.5282 0.9885 1.1689 1.3978 3.1947 0.0961 -4.76 0.419 0.1259

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APPENDIX D. X-RAY DIFFRACTION

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Modal mineralogy of the global fraction based on quantification of XRD spectra K- Sample Qtz Alb Cal Mica+Ill Sm IS Kln Hem Pyx Amph Spar BL-1 69.8 11.8 10.3 0.5 1.9 1.9 2 1 0.3 0 0 BL-2 69.9 7.4 12.3 1.4 2.9 2.2 3.4 0.7 0.2 0 0 BL-3 52 6.5 7.7 15.7 5.8 4.9 3.4 0.6 0.2 0 0.5 BL-4 60.2 12.6 14.9 0.6 5.4 1.3 2.4 0.4 0.2 0 0 BL-5 56.3 9.1 10.9 0.5 5.1 4.9 5.9 0.7 0.1 3.9 0.4 BL-6 59.7 9.6 14.4 5.1 2.9 2 4.8 0.6 0.3 0 0 BL-7 48 10.5 14.1 9.2 7.2 4.3 3.4 0.8 0.2 0 0 BL-8 48.3 10.6 13.6 14.1 3.9 3.9 2.1 0.9 0.1 0 0 BL-8b 50.4 11 13.1 1.5 8 4.6 4.9 2.2 0.2 0 0 BL-9 57.7 11.6 10 9.2 3.3 3.3 2.5 1.2 0.2 0 0 BL-9b 50.5 14.8 8.5 2.9 9.7 5.1 4.9 1.3 0.2 0 0.4 BL-10 56.3 8.3 10.7 2.6 4 8.2 7 0.9 0.1 4.4 0 BL-10b 52.3 8.3 12.6 14.1 6.2 2 2.7 0.9 0.3 0 0 BL-11 68.7 10.2 12.3 0.7 2.5 2.3 1.4 0.9 0.1 0 0 BL-12 56.3 7.2 11.1 11 2.9 2.9 5.7 0.4 0.2 0 0 BL-13 61.2 10.2 10.5 1.8 4.6 6.4 2.5 0.6 0.2 0 0 BL-14 57.4 6.1 12.7 5.4 6.3 5.2 4 0.5 0.2 0 0 BL-15 69.2 5.8 10.3 2.1 4.3 3 2.8 0.6 0.4 0 0 BL-15b 47 6.7 6.9 24.6 7.3 1.3 1.8 1.2 0.1 0 0 BL-16 63.5 3.8 5.9 16.5 2.6 3.1 1.8 0.4 0.1 0 0 BL-17 63.1 2.7 3.8 20.2 2.4 1.9 2.5 0.6 0.2 0 0 BL-18 53.8 1.2 3.5 30 2.7 0.8 2.5 0.2 0.1 0 0 Average 57.80 8.45 10.46 8.62 4.63 3.43 5.19 0.80 0.19 0.38 0.02 Std.Dev. 7.13 3.24 3.16 8.44 2.08 1.81 1.79 0.41 0.08 1.20 0.08 Abbreviations: Qtz: quartz, Alb: albite, K-spar: potassium feldspar, Cal: calcite, Mica + Ill: sum of muscovite and illite, Sm: smectite, IS: illite/smectite, Kln: kaolinite Hem: hematite, Pyx: pyroxene, Amph: amphibole.

Comparison of phase name in Sybilla program and the true/ interpreted phase name for models below Low Ordered, Disordered, Discrete Discrete Crystallinity interstratified smectite rich, Kaolinite Illite Smectite Illite illite-smectite illite-smectite Interpreted Ill Ill Low R1 IS R0-R1 ISS SSS Kln Phase Name Id’ed Illite IS R1 IS R0 ISS R0 SSS Kaolinite Sybilla Phase

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Summarized Sybilla modelling parameters Phase D 2θ σ* Tmean L1 L2 ± L3 Paa P2/(P2+P3) (%) (%) SSS BL-4 14.78 5.98 12 2.74 46 54 0.49 BL-5 15.24 5.8 12 2.45 51 49 0.42 BL-8b 15.53 5.69 12 2.45 58 42 0.24 Bl-9b 14.87 5.82 12 3.32 65 35 0.33 BL-10 15.58 5.67 12 2.45 63 37 0.32 BL-17 15.2 5.82 12 2.74 61 39 0.3 ISS R0 BL-4 12.78 6.91 12 8.25 48 52 0.5 BL-5 ------BL-8b 13.71 6.45 12 3.61 47 53 0.5 Bl-9b 12.52 7.06 12 4.19 60 40 0.62 BL-10 ------BL-17 14.48 6.1 12 2.74 49 51 0.5 IS R0 BL-4 11.49 7.69 12 7.38 72 28 0.5 BL-5 11.58 0.69 12 7.38 69 31 0.44 BL-8b 11.8 0.76 12 4.48 76 24 0.5 Bl-9b 12.21 0.65 12 5.93 65 35 0.5 BL-10 11.92 0.6 12 21.3 60 40 0.5 BL-17 12.52 0.7 12 4 70 30 0.5 Ill Low BL-4 10.34 8.55 12 5.64 99 1 0.5 - BL-5 10.27 8.61 12 6.8 96 4 0.18 - BL-8b 10.3 8.59 12 6.22 99 1 0.5 - Bl-9b 10.36 8.53 12 5.64 97 3 0.5 - BL-10 10.45 8.46 12 5.06 99 1 0.5 - BL-17 10.56 8.37 12 4.48 99 1 0.5 - Ill BL-4 10.04 8.8 12 15.04 - - - - BL-5 10.04 8.8 12 13.09 - - - - BL-8b 10 8.84 12 15.43 - - - - Bl-9b 9.98 8.86 12 14.65 - - - - BL-10 10.02 8.82 12 15.04 - - - - BL-17 10.02 8.82 12 17.38 - - - - D- D-spacing of phase; σ*- orientation of particles on the mounted X-ray slide; Tmean- crystallinity of phase; L1 (%)- percent of 1st layer in multi-layer phase; L2 ± L3 (%)- percent of 2nd or sum of 2nd and 3rd layers in multi-layer phase; Paa- probability of finding layer 1 after the preceding layer 1 in a 2-component system; P2/(P2+P3)- probability of

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APPENDIX E. SEM-EDS OF CLAY MINERALS

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Representative EDS analyses (wt%) of clay mineral phases from selected samples. Sample BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 Site 1 2 3 3 3 3 2 2 2 2 2 2 2 2 Analysis 5 20 31 39 41 43 75 79 80 82 87 90 95 105 Mineral IS ISS IS ISS IS IS ISS ISS IS IS IS IS ISS ISS SiO2 60 66.8 62.9 65 65.6 64.3 69.4 67.3 62.2 62.2 61.2 62.3 73.5 67.7 Al2O3 21.9 18.3 21.7 18.7 19.8 22.7 18.6 19.8 22.5 21.6 21.5 22.1 14.2 18.6 FeO 6.3 8.6 7.1 9.3 7.1 5.4 5.5 6 7 8.5 7.9 6.5 6.3 6.3 MgO 3.2 2.6 2.6 3.2 2.6 1.4 2.6 2.4 2.7 2.6 3 3.1 2.9 2.3 K2O 4.5 2.4 3.2 2.5 3 1.4 2.3 2.5 3.7 3.6 3.3 4.7 2.2 2.8 CaO 3 0.7 1.4 1 0.9 0.5 1 0.8 0.9 1.2 1.8 1.1 0.7 1.1 Na2O 0.2 0 0 4.2 0 0.9 0.2 0 0.2 0 0 0 TiO2 0.7 0.7 0.8 0 0 0.4 0.5 0.7 0 0.6 0 0 1.1 Total 99.8 100.1 99.7 99.7 99.7 99.9 99.8 100.2 99.9 99.7 99.5 99.8 99.8 99.9

Sample BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-14 BL-14 BL-14 BL-14 Site 2 2 2 2 2 2 2 3 4 4 1 1 1 1 Analysis 110 111 112 114 116 117 118 123 138 146 155 162 163 169 Mineral ISS ISS IS IS IS IS IS ISS IS IS IS IS IS IS SiO2 66.7 70.6 63.8 61.2 62.4 62.4 65.4 67 63.5 66.3 61.2 62.9 62 62.3 Al2O3 19.8 17.6 22.4 23.6 22.7 23.1 20.3 19.8 22.6 21.1 20.9 22.4 22.2 23.9 FeO 6.4 5.5 6.8 7 5.8 6.7 6.7 5.9 5.4 5.2 6.5 7.5 7.3 4.9 MgO 2.5 2.2 2.8 3 3.6 3.6 2.8 3.3 3.6 3 3.3 3.7 4.2 3.7 K2O 2.9 3 3.2 3.3 3 2.9 3.8 2.7 2.5 2.6 3.1 3.2 3.2 3.9 CaO 1.2 1.1 1 1.1 1.3 1.3 1 0.8 1.1 1.3 4.6 0 1.2 1 Na2O 0 0 0 0 0 0 0 0.3 0.4 0 0.3 0 0 0.4 TiO2 0.5 0 0 0.6 0 0 0 0 0 0.5 0 0 0 0 Total 100 100 100 99.8 98.8 100 100 99.8 99.1 100 99.9 99.7 100.1 100.1 Ill- illite; I/S- illite - smectite; I/SS- illite – smectite rich in smectite; Kln - kaolinite; I M- illitized muscovite

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Representative EDS analyses (wt%) of clay mineral phases from selected samples. Sample BL-14 BL-14 BL-14 BL-11 BL-11 BL-11 BL-11 BL-11 BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 Site 2 2 4 1 1 1 2 2 1 1 1 1 1 2 Analysis 172 174 206 211 232 235 247 249 2 3 9 10 11 14 Mineral IS IS IS IS ISS ISS IS ISS Ill Ill Ill Ill Ill Ill SiO2 66.5 60.9 63.3 58.9 67 67 64.9 72.1 56 58.4 57.5 48.1 56.9 60.8 Al2O3 20.2 24.6 21.7 20.4 19.5 19.4 20 16.4 20.8 22.4 22.8 19 21.3 23.6 FeO 5.6 8.2 7 9.2 6.1 6.2 7 5.1 7.9 8.4 6.9 7 7.7 6.8 MgO 3.2 3.4 3.4 4.6 3.5 3.1 3.7 2.8 3.2 2.4 4 3.4 3.2 3.9 K2O 2.8 3 2.9 4.6 2.4 2.8 2.9 2.6 3.3 5.3 2.5 2.3 3.3 3 CaO 1.5 0 1.1 0.9 0.9 0.9 0.9 0.7 6.7 2 5.9 19.7 7.3 1.2 Na2O 0 0 0 0 0 0 0 0 0 0.6 0 0 0.3 0 TiO2 0 0 0.6 1.3 0.6 0.5 0.4 0 0.9 0.4 0 0 0 0.5 Total 99.8 100.1 100 99.9 100 99.9 99.8 99.7 98.8 99.9 99.6 100 100 99.8

Sample BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 BL-7 Site 2 2 2 2 2 3 3 3 3 3 3 3 3 3 Analysis 19 21 24 28 29 30 32 33 35 36 37 38 40 42 Mineral Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill SiO2 60.7 56.5 61.6 59.2 60.4 59.7 50.9 57.4 55.5 55.5 58.5 59.2 61.3 59.9 Al2O3 23.9 24.3 22.4 23.7 24.3 23.7 22.2 25.2 23 23 25.2 22.5 23.5 23.5 FeO 7.2 8.3 7.8 9 7.4 8 13.9 6.3 12.6 12.6 9.2 8.2 7.3 8.1 MgO 3.8 5.5 3.4 3.8 3.6 3.1 4.4 2.6 3.2 3.2 3 2.8 2.9 3.4 K2O 3.2 3 3.2 3.3 3.4 3.3 5.2 3.4 3.3 3.3 3.1 3.5 3.2 3.3 CaO 0.6 0.9 0.8 0.5 0.8 1.1 1 4.5 1.3 1.3 1.1 1.3 1 0.9 Na2O 0 0 0 0 0 0 0 0 0.5 0.5 0 0 0 0 TiO2 0.5 1.5 0.8 0.5 0 0.9 2.2 0.4 0.5 0.5 0 2.5 0.6 0.5 Total 99.9 100 100 100 99.9 99.8 99.8 99.8 99.9 99.9 100.1 100 99.8 99.6 Ill- illite; I/S- illite - smectite; I/SS- illite – smectite rich in smectite; Kln - kaolinite; I M- illitized muscovite

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Texas Tech University, Colton Mallett, May 2021

Representative EDS analyses (wt%) of clay mineral phases from selected samples. Sample BL-7 BL-7 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 Site 3 3 1 1 1 1 1 1 1 1 1 2 2 2 Analysis 44 45 52 55 56 57 58 59 62 63 64 76 77 81 Mineral Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill SiO2 61.6 58.6 59 59.1 55.2 55.9 59.8 59.3 57.7 59.7 59.6 56.9 61.7 59.5 Al2O3 23.5 26.5 23.5 24.2 22 22.1 24.7 23.2 22.3 23.6 22 22.5 24.2 23.9 FeO 7.8 6.7 9.3 9.3 12 10.2 8.1 8.8 12.4 9 10.4 11.4 7 7.9 MgO 3.1 2.1 2.5 3.1 4.8 5.2 3.4 3.1 3.2 3 3.7 2.4 3.4 3.3 K2O 3 3.3 3.8 2.9 3.9 3.9 2.8 3.2 2.6 3.1 2.9 3.7 3 3.4 CaO 0.7 2 1.6 1.4 0.8 0.9 1.3 1.5 1.5 1.6 1.3 1.3 0.7 1.3 Na2O 0 0.8 0 0 0 0 0 0 0 0 0 0 0 0 TiO2 0 0 0 0 1.3 1.7 0 1 0 0 0 1.3 0 0.5 Total 99.7 100 99.7 100 100 99.9 100.1 100.1 99.7 100 99.9 99.5 100 99.8

Sample BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 Site 2 2 2 2 2 2 2 2 2 2 3 3 3 3 Analysis 84 85 86 92 94 101 103 107 108 115 121 124 125 126 Mineral Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill SiO2 58 59.5 60.9 60.8 60.1 60.4 59.4 59.2 60.2 60.1 62.6 60.7 60.2 60.2 Al2O3 23.3 25 22.6 23.9 24.1 24 22.8 23.6 23.3 23.1 23.8 24.1 24.3 24.2 FeO 7.3 7.4 8.5 7.6 7.8 7.5 8 8 7.9 8.1 5.8 7 7.6 7.2 MgO 3.6 3.5 2.8 3.3 2.7 3.3 3.7 3.3 3.3 3.5 2.9 3.6 3.6 3.4 K2O 2.8 3.3 3.1 2.7 3.2 3.1 4.1 2.7 3.2 3.2 2.5 3.1 3.3 3.4 CaO 1.4 1.3 1.6 1.1 1 1 1.1 1.3 1.2 1.3 1.2 1.2 0.5 0.6 Na2O 0 0 0 0 0 0 0 0 0 0 0 0 0 0 TiO2 0.5 0 0 0.5 0.6 0.6 1 1.7 0.9 0.5 0.9 0 0.5 0.8 Total 96.9 100 99.5 99.9 99.5 99.9 100.1 99.8 100 99.8 99.7 99.7 100 99.8 Ill- illite; I/S- illite - smectite; I/SS- illite – smectite rich in smectite; Kln - kaolinite; I M- illitized muscovite

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Texas Tech University, Colton Mallett, May 2021

Representative EDS analyses (wt%) of clay mineral phases from selected samples. Sample BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 BL-13 Site 3 3 3 3 3 3 4 4 4 4 4 4 4 4 Analysis 128 130 131 132 133 134 136 137 139 140 143 144 145 147 Mineral Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill SiO2 59.3 60.3 60.6 61.2 60.7 61.4 59.6 59.1 59.8 60.5 60.9 60.5 57.2 61 Al2O3 25.2 23.7 24.4 23.7 23.4 22.6 24.9 22.3 23.7 23.8 24.1 23.9 24.4 24.7 FeO 6 7.3 7.1 7 7.7 7.1 6.3 8.4 8.4 7.1 7.6 7.1 7.5 6.6 MgO 2.9 3.8 3.5 3.6 3.3 3 3.5 3.4 3.2 3.6 3.7 3.3 3.7 3.4 K2O 4.1 3.4 2.8 3.1 3.1 3.9 2.4 2.9 2.5 2.9 3.1 2.9 4.3 3 CaO 1 0.7 1.5 0.7 1.1 1.4 1.3 1.6 1.5 1.5 0.6 1.9 2.2 1.1 Na2O 0 0 0 0 0 0 0.3 0 0 0 0 0 0 0 TiO2 0.9 0.6 0 0.5 0.5 0.5 1 1.3 0.5 0.4 0 0 0.9 0 Total 99.4 99.8 99.9 99.8 99.8 99.9 99.3 99 99.6 99.8 100 99.6 100.2 99.8

Sample BL-13 BL-13 BL-13 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 Site 4 4 4 1 1 1 1 1 1 1 1 1 1 1 Analysis 148 149 150 151 155 156 157 158 159 160 161 164 166 168 Mineral Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill SiO2 60 61.3 60.5 60.4 61.2 62 60.8 60 60.2 61.7 61 61.3 54.5 61.1 Al2O3 23.9 24.3 23 23 20.9 23.5 23.5 23.3 23.9 23.4 23.1 22.7 22.4 24.2 FeO 7.6 6.1 7.1 7.7 6.5 6.7 8.4 8.5 7.9 7.6 7.7 7.7 12.8 6.8 MgO 3.7 3.4 4.2 3.7 3.3 3.6 3.7 3.4 3.7 3.7 3.9 4.4 4.9 3.8 K2O 2.9 3.9 4 3 3.1 3.5 3.6 3 3.2 3.6 3.3 3.5 3.6 3.1 CaO 1.6 1 0.8 1.1 4.6 0.5 0 0.7 1.1 0 1.1 0 1 0.8 Na2O 0 0 0 0 0.3 0 0 0 0 0 0 0 0 0 TiO2 0 0 0 0.9 0 0 0 0.8 0 0 0 0 0.8 0 Total 99.7 100 99.6 99.8 99.9 99.8 100 99.7 100 100 100.1 99.6 100 99.8 Ill- illite; I/S- illite - smectite; I/SS- illite – smectite rich in smectite; Kln - kaolinite; I M- illitized muscovite

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Texas Tech University, Colton Mallett, May 2021

Representative EDS analyses (wt%) of clay mineral phases from selected samples. Sample BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 Site 2 2 2 2 2 2 2 2 2 2 2 2 2 3 Analysis 170 171 173 175 176 177 178 179 180 181 182 184 185 190 Mineral Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill SiO2 60.8 60 60.7 60.6 59.9 62.3 53.5 62.1 61 60.5 59.9 59.5 60.5 60.4 Al2O3 23.2 23.2 25.2 23.1 24.5 22.7 22.8 23.2 23.8 23.6 23.3 21.7 24.4 23.6 FeO 7.3 7.6 7.5 7.1 6.7 6.8 13.4 5.4 7.2 8.4 7.8 7.1 6.8 6.8 MgO 3.3 3.5 3.4 3.7 3.5 3.8 6.1 2.9 3.6 3.4 3.8 3.9 3.2 3.6 K2O 2.9 3.3 3.2 3.4 4.5 3.2 3 5.3 3.2 3 2.8 2.7 2.7 3 CaO 1.7 1.8 0 1.3 1 1 0.5 0.7 0 0.5 1.4 4.7 2 1.8 Na2O 0 0 0 0 0 0 0.4 0.5 0 0 0 0 0 0 TiO2 0.5 0.4 0 0 0 0 0 0 0.9 0.6 0.7 0 0 0.7 Total 99.7 99.8 100 99.2 100.1 99.8 99.7 100.1 99.7 100 99.7 99.6 99.6 99.9

Sample BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 BL-14 Site 3 3 3 4 4 4 4 4 4 4 4 4 4 4 Analysis 191 192 193 194 195 196 197 198 199 201 202 203 204 205 Mineral Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill SiO2 60.2 60.6 59.7 59.6 61 60.3 60.8 60.6 60.6 61.3 61 61.5 59.7 61.2 Al2O3 23.7 24.1 23.7 23.7 23.5 23.6 23.6 23.3 23.6 22.4 23.2 23.6 23.9 23.6 FeO 6.9 7.9 7.7 8.1 5.7 6.7 7.1 7.9 6.9 6.2 6.3 6.2 8 5.9 MgO 3.8 3.6 3.6 3.7 3.2 3.8 3.5 3.6 3.8 3.7 3.9 3.7 3.8 4 K2O 3.3 3.3 3.3 3 5.3 3.4 3.2 3.3 3.1 3.9 3.3 3.5 3.5 3 CaO 1.5 0 1.3 1.2 1.3 1.4 1.1 0.8 1.2 1.4 1.2 1.4 0.8 1.2 Na2O 0 0 0 0 0 0 0 0 0 0 0.3 0 0 0 TiO2 0.4 0 0.6 0.5 0 0.7 0.6 0.5 0.6 1.2 0.5 0 0 0.9 Total 99.8 99.5 99.9 99.8 100 99.9 99.9 100 99.8 100.1 99.7 99.9 99.7 99.8 Ill- illite; I/S- illite - smectite; I/SS- illite – smectite rich in smectite; Kln - kaolinite; I M- illitized muscovite

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Texas Tech University, Colton Mallett, May 2021

Representative EDS analyses (wt%) of clay mineral phases from selected samples. Sample BL-11 BL-11 BL-11 BL-11 BL-11 BL-11 BL-11 BL-11 BL-11 BL-11 BL-11 BL-11 BL-11 BL-11 Site 1 1 1 1 1 1 1 1 1 1 2 2 2 2 Analysis 210 212 214 215 216 220 221 222 224 227 244 246 251 256 Mineral Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill Ill SiO2 60 62.1 60.2 58.7 60.8 57.8 59.5 60.6 61.8 62.1 60.2 60.6 59.2 59.9 Al2O3 24.1 23.4 22.8 22.5 23.8 21.8 22.2 23.1 21.3 23 25.6 22.4 22.8 23.1 FeO 6.8 5.8 8.2 9.8 6.7 7.7 8.7 7.5 7.8 7.6 6.6 8.3 8.4 8.3 MgO 3.5 3.7 3.9 4.1 4.6 3.2 3.7 3.9 3.5 3 3.8 3.8 3.5 3.8 K2O 4.1 3.8 3.1 3.4 2.8 3.7 3.9 3.3 4 3.3 3.1 3.4 3.9 3.5 CaO 0.7 0.7 0.9 0.9 0.8 1.1 1 1.1 1 0.9 0.6 1.2 1.1 1.2 Na2O 0 0 0 0 0 0 0 0 0 0 0 0 0 0 TiO2 0.6 0.6 0.7 0.7 0 4.6 1.1 0.4 0.5 0 0 0 0.7 0 Total 99.8 100.1 99.8 100.1 99.5 99.9 100.1 99.9 99.9 99.9 99.9 99.7 99.6 99.8

Sample BL-11 BL-11 BL-11 BL-13 BL-14 BL-14 BL-14 BL-14 BL-14 BL-11 BL-11 BL-11 BL-11 BL-11 Site 2 2 2 2 2 3 1 Analysis 257 258 259 91 183 186 187 188 189 217 218 219 223 228 Mineral Ill Ill Ill Kln Kln Kln Kln Kln Kln Kln Kln Kln Kln Kln SiO2 60.8 60.3 59.5 59.4 62.1 55.8 56.2 57.1 56.8 54.3 54.1 54.2 54 55.5 Al2O3 23.2 22.9 25.3 33.1 32 44.2 42.1 38.9 40.2 41.6 41.4 41.8 42.6 40.2 FeO 6.9 8 6.6 3.4 2.6 0 0.6 1.5 1.2 3.5 3.8 3.4 2.7 2.5 MgO 4.1 3.6 3.8 1.4 1.5 0 0.5 1.1 0.7 0.6 0.7 0.6 0 0.7 K2O 3.6 3.3 3.1 2 1.4 0 0.4 1 0.7 0 0 0 0.4 0.6 CaO 0.5 1.2 0.8 0.6 0.4 0 0.3 0.4 0.2 0 0 0 0.3 0.5 Na2O 0 0 0 0 0 0 0 0 0 0 0 0 0 0 TiO2 0.8 0.5 0.5 0 0 0 0 0 0 0 0 0 0 0 Total 99.9 99.8 99.6 99.9 100 100 100.1 100 99.8 100 100 100 100 100 Ill- illite; I/S- illite - smectite; I/SS- illite – smectite rich in smectite; Kln - kaolinite; I M- illitized muscovite

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Texas Tech University, Colton Mallett, May 2021

APPENDIX F. CORE PROFILE

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Texas Tech University, Colton Mallett, May 2021

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