Geochemical Analysis of Soil and Lake Deposits As Indicators of Paleoenvironmental Conditions; Examples from Texas, New Mexico, and D.R

Southern Methodist University SMU Scholar

Earth Sciences Theses and Dissertations Earth Sciences

Fall 2019

Geochemical Analysis of Soil and Lake Deposits as Indicators of Paleoenvironmental Conditions; Examples from Texas, New Mexico, and D.R. Congo

Mary Milleson Southern Methodist University, [email protected]

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Recommended Citation Milleson, Mary, "Geochemical Analysis of Soil and Lake Deposits as Indicators of Paleoenvironmental Conditions; Examples from Texas, New Mexico, and D.R. Congo" (2019). Earth Sciences Theses and Dissertations. 12. https://scholar.smu.edu/hum_sci_earthsciences_etds/12

This Dissertation is brought to you for free and open access by the Earth Sciences at SMU Scholar. It has been accepted for inclusion in Earth Sciences Theses and Dissertations by an authorized administrator of SMU Scholar. For more information, please visit http://digitalrepository.smu.edu. GEOCHEMICAL ANALYSIS OF SOIL AND LAKE DEPOSITS AS

INDICATORS OF PALEOENVIRONMENTAL CONDITIONS;

EXAMPLES FROM TEXAS, NEW MEXICO,

AND D.R. CONGO

Approved by:

______Dr. Neil J. Tabor Professor of Sedimentary Geochemistry

______Dr. John Geissman Professor of Geosciences

______Dr. Robert T. Gregory Professor of Geochemistry

______Dr. Dale Winkler Research Professor of Paleontology

______Dr. Crayton Yapp Professor of Geochemistry

GEOCHEMICAL ANALYSIS OF SOIL AND LAKE DEPOSITS AS

INDICATORS OF PALEOENVIRONMENTAL CONDITIONS;

EXAMPLES FROM TEXAS, NEW MEXICO,

AND D.R. CONGO

A Dissertation Presented to the Graduate Faculty of

Dedman College

Southern Methodist University

Partial Fulfillment of the Requirements

for the degree of

Doctor of Philosophy

with a

Major in Geology

Mary Milleson

B.S., Geology, Southern Methodist University

August 6, 2019 ACKNOWLEDGEMENTS

Many individuals have made this work possible. Peter Van Metre and USGS provided the

White Rock Lake core and valuable field notes. I also owe thanks to N. Tabor and Frank Tabor for their field work and sample collection in north central New Mexico. Additionally, I would like to acknowledge N. Tabor, S. Thomas, S. Myers, and the Royal Museum of Central Africa for collection of the Dekese core.

I owe a debt of gratitude to Dr. Kurt Ferguson and Roy Beavers for allowing me to follow them around and pester them until they showed me how to operate equipment. I would like to extend my gratitude to Dr. Robert Gregory for experience I gained as his teaching assistant. I would like to thank my entire family for their unyielding support over the years, and in particular, Billy G. Crouch III, who prodded me over the finish line.

My deepest gratitude is expressed to my advisor, Neil Tabor. Dr. Tabor offered me my first laboratory experience, my first field trip, and my first teaching assistant position. Through his mentorship, I have learned persistence and resourcefulness. I am extraordinarily thankful he possessed the patience to deal with my shenanigans and excuses over the duration of my time at

SMU.

iii Milleson, Mary B.S., Southern Methodist University

Geochemical Analysis of Soil and Lake Deposits as Indicators of Paleoenvironmental Conditions; Examples from Texas, New Mexico, and D.R. Congo

Advisor: Professor Neil Tabor

Doctor of Philosophy conferred August 6, 2019

Dissertation completed August 4, 2019

Geochemical proxies of paleoenvironmental conditions will be presented from studies of modern lacustrine deposits from White Rock Lake, Dallas Texas, Upper Pennsylvanian through

Lower Permian paleosols from Mora County New Mexico, and Upper Paleozoic lacustrine strata from the Congo Basin, D.R. Congo.

The composition of organic matter preserved in the sediments of White Rock Lake will be utilized to establish a temporal relationship between urbanization of the White Rock Creek watershed and lake system. Geochemical proxies of watershed conditions allow for the evaluation of changes in watershed vegetation, sediment input, and lake productivity, which are correlated with historical records of urbanization and changing land use practices over the past century. Since the onset of urbanization, measured sedimentary organic matter δ13C values have shifted by ~9‰, from -16‰ in the oldest sediments to -25‰ in the most recent deposits.

Similarly, C/N ratios have shifted from 21.3 to 8.0 over the same interval of time. Weight percent carbon preserved in lake sediments exhibits a shift from ~2.2% and the base of the core to ~4.5% and the top. The δ13C and C/N values of sedimentary organic matter likely indicate a

iv shift between contributions from predominantly allochtonous C4 photosynthesizers, representative an agricultural landscape, to sediments dominated by primarily C3 photosynthesizing organisms. These trends suggest increased primary productivity and an influx of detrital material into White Rock Lake is resultant from urbanization of the watershed landscape.

Lithostratigraphic, mineralogic, and stable isotope proxies of equatorial paleoclimate from Upper Pennsylvanian (Missourian-Virgilian) through Lower Permian (Wolfcampian) strata of the Taos Trough, north-central New Mexico will be utilized to reconstruct paleoenvironmental conditions during pedogenesis. The Upper Paleozoic sequence includes mixed marine carbonate and terrestrial clastic rocks in Pennsylvanian strata (Madera Fm) which transition to terrestrial- dominated siliciclastic rocks in the Lower Permian (Sangre de Cristo Fm). Paleosol morphologies include eutric Argillisols in Missourian rocks, Calcisols in Virgilian and

Wolfcampian rocks, and gypsic subgroups in upper Wolfcampian strata. X-ray diffraction analysis of pedogenic clay minerals shows a shift from kaolinite, smectite, and illite assemblages representative of seasonal, subhumid environments in the lower part of the stratigraphy to poorly-ordered 2:1 phyllosilicate and illite assemblages, representative of increased aridity and decreased seasonality, in the uppermost part of the stratigraphy. The stratigraphic distribution of these morphologies suggests a stepped change from a subhumid, seasonal climate in Missourian time to a nonseasonal arid climate in end-Wolfcampian time.

There is no significant stratigraphic trend among organic matter δ13C values, which range from ~-22.5 to -24‰. However, paleosol calcite δ13C values are ~-5‰ in Virgilian strata, ~-8 in lower Wolfcampian strata, and ~-5‰ in topmost Wolfcampian strata. These data suggest

v relatively high atmospheric PCO2 (or low biological productivity) during deposition of

Pennsylvanian and upper Wolfcampian strata, and relatively low atmospheric PCO2 (or high biological productivity) during deposition of lower Wolfcampian strata. The δ18O value of pedogenic calcite ranges from ~-4 to -6.5‰ throughout the Missourian, Virgilian and

Wolfcampian strata. However, calcite δ18O values of the top-most Wolfcampian strata range from +1 to +3‰. This stratigraphically abrupt, +5-7‰, shift in calcite δ18O values is consistent with predicted changes in tropical rainfall δ18O values during deglaciation of the ice-sheet in

Gondwanaland.

Lithostratigraphy, mineralogy, and isotopic values of lacustrine sediments from Upper

Carboniferous-Lower Permian strata in the Lukuga Formation, central Democratic Republic of

Congo will be presented. The stratigraphic distribution and internal concentration of strata with outsized clasts delineates a single, large-scale pattern of climatic change from relatively cold to warm conditions through the entire succession. Mineralogical composition of the <2µm fraction of sediments from the Lukuga Formation delineates a more complex, at least two stratigraphically long climate cycles, beginning with cold and then changing to warm conditions.

Finally, the δ18O values of early burial lacustrine calcite cements from the Lukuga Formation delineate at least three, and possibly as many as six climate cycles, each beginning with cold and then changing to warmer conditions. This regional reconstruction of Permo-Carboniferous terrestrial paleoclimate is similar to that interpreted from mixed marine strata in the Karoo and

Kalahari basins (Scheffler et al., 2006), and strongly supports the notion that Gondwanaland was characterized by repeated shifts from glacial to inter- and/or non-glacial conditions (Montañez et al., 2007, Fielding, et al., 2008). Furthermore, the magnitude of climate shifts inferred from this terrestrial study may provide a useful means of correlation to strata from contemporaneous vi marine basins to other near-field terrestrial basins in Gondwanaland, and to far-field, paratropical terrestrial basins in Euramerica.

vii TABLE OF CONTENTS

LIST OF FIGURES ...... x

LIST OF TABLES ...... xiii

Chapter

1. INTRODUCTION ...... 1

2. GEOCHEMICAL ANALYSIS OF LAKE SEDIMENTS AS A PROXY FOR WATERSHED URBANIZATION, WHITE ROCK LAKE, DALLAS, TX ...... 5

Introduction ...... 5

Geologic Setting...... 6

Methods...... 11

Sampling ...... 11

Stable Isotopes ...... 12

C/N Ratios ...... 14

Results ...... 14

Chronology ...... 14

Organic Matter ...... 15

Calcite Abundance and Stable Isotope Values ...... 18

Carbonate Morphology ...... 20

Discussion ...... 26

Preservation of Isotopic and Chemical Ratios ...... 28

Organic Matter Stable Isotope and Elemental Data ...... 29

Carbonates...... 37

Conclusions ...... 41 viii 3. PENNSYLVANIAN-PERMIAN PALEOENVIRONMENTAL TRENDS FROM LITHOSTRATIGRAPHY, CLAY MINERALOGY, AND STABLE ISOTOPE GEOCHEMISTRY OF THE MADERA AND SANGRE DE CRISTO FORMATIONS, MORA COUNTY, NEW MEXICO ...... 43

Introduction ...... 43

Geologic Background ...... 45

Methods...... 49

Fieldwork and Sampling ...... 49

Paleosol Mineralogy ...... 56

Paleosol Carbonates ...... 57

Fossil Organic Matter δ13C ...... 58

Results ...... 59

Pedotypes ...... 59

Paleosol Mineralogy ...... 65

Petrographic Analysis of Paleosol Carbonates ...... 66

Stable Isotope Geochemistry of Paleosol Carbonate and Organic Matter ...... 67

Discussion ...... 78

Pedotypes ...... 78

Petrographic Analysis of Carbonates ...... 81

Paleosol Clay Mineralogy ...... 82

Stable Isotope Compositions of Paleosol Calcite and Organic Matter ...... 86

Oxygen Isotopes...... 86

Carbon Isotopes ...... 87

Conclusions ...... 92

4. PERMO-CARBONIFEROUS PALEOCLIMATE OF THE CONGO BASIN: EVIDENCE FROM LITHOSTRATIGRAPHY, CLAY MINERALOGY, AND STABLE ISOTOPE GEOCHEMISTRY ...... 93

Introduction ...... 93

Basis for Paleoclimate Reconstruction from the Lukuga Formation ...... 95

Geologic Background ...... 96

Tectonics ...... 96

Basin Fill ...... 97

Lukuga Formation ...... 98

Dekese Core ...... 99

Methods...... 101

Sampling ...... 101

U-Th-Pb Isotope Systematics of Detrital Zircons ...... 102

X-Ray Diffraction ...... 104

Isotope Geochemistry ...... 106

Results ...... 107

Lithostratigraphy ...... 107

Mineralogy ...... 109

Bulk X-Ray Powder Diffraction Analyses ...... 109

Clay-size (< 2μm e.s.d) Fraction Analyses ...... 118

Detrital Zircons ...... 119

Calcite ...... 119

Isotope Geochemistry ...... 127

x Discussion ...... 127

Interpretation ...... 130

Lithostratigraphy ...... 130

Clay Mineralogy ...... 131

Calcite Stable Isotope Composition ...... 132

Conclusions ...... 134

5. SUMMARY ...... 137

APPENDIX ...... 139

WRL-1...... 139

MORA-1 ...... 141

MORA-2 ...... 155

MORA-3 ...... 161

DEKESE-1 ...... 165

DEKESE-2 ...... 189

REFERENCES ...... 204

xi LIST OF FIGURES

Figure

2.1 Map of Texas highlighting the White Rock lake location, subdivisions of the White Rock Creek watershed, and county boundaries...... 10

2.2 Dated 137Cs profile of WRL sediments ...... 16

2.3 Age model of White Rock Lake sediment core ...... 17

2.4 Age model showing approximate age of each sample in the White Rock Lake sediment core ...... 17

13 2.5 Plot of C/N ratios and δ C(SOM) values representative of C4, C3, and algal contributions to organic matter preserved in lacustrine sediments ...... 21

13 2.6 Plot of stratigraphic position versus δ C(SOM) values, C/N ratios, weight percent carbon of White Rock Lake sedimentary organic matter...... 22

2.7 Stratigraphic distribution of weight percent N and C values of White Rock Lake sedimentary organic matter ...... 23

13 18 2.8 Plot of carbonate carbon isotope (δ C(CC)) vs carbonate oxygen isotope (δ O(CC)) values of White Rock Lake core samples ...... 24

13 18 2.9 Plot of stratigraphic position versus δ C(CC) values, δ O(CC), and weight percent carbon of White Rock Lake carbonate samples ...... 25

2.10 Land use change over time for Collin County and Dallas County from 1905 to 2005 ...... 32

3.1 Paleogeography of Late Paleozoic basins and uplifts in North America...... 46

3.2 Stratigraphy and age assignments for the upper Paleozoic strata of the Rowe-Mora basin of northern New Mexico ...... 47

3.3 Field photos of exposures along Highway NM518 of the Sangre de Cristo Formation in the Rowe-Mora basin ...... 49

xii

3.4 Representative stratigraphic section of the Madera and Sangre de Cristo formations of the Rowe-Mora basin ...... 50

3.5 Detailed stratigraphic section of the Madera and Sangre de Cristo formations of the Rowe- Mora basin ...... 51

3.6 Diagrams of the seven paleosol type morphologies present in the Madera and Sangre de Cristo formations of the Rowe-Mora basin ...... 62

3.7 Stratigraphic position versus pedotypes, cumulative clay mineral percentages, and 13 Δ C(CC-OM) of paleosols from the Madera and Sangre de Cristo formations ...... 63

3.8 Background-subtracted X-ray diffractograms of oriented aggregates from the clay-size (<2µm) fraction of paleosol matrix sample ...... 67

3.9 Examples of common micritic calcite textures observed in paleosol nodules from the Madera and Sangre de Cristo formations...... 68

3.10 Examples of alteration of primary mineral textures in paleosol carbonate nodules and nodular limestone unit from the Madera and Sangre de Cristo formations ...... 69

3.11 Measured δ18O values versus δ13C values of paleosol carbonates from the Madera and Sangre de Cristo formations of the Rowe-Mora basin...... 71

3.12 Cross plot of stratigraphic position versus δ18O values of carbonates from the measured section of the Madera and Sangre de Cristo formations ...... 72

3.13 Cross plot of the stratigraphic position versus δ13C values of identifiable, strata-bound organic matter, paleosol carbonate-associated organic matter, and carbonate samples of the Madera and Sangre de Cristo formations ...... 73

13 3.14 Plot of carbonate carbon isotope (δ C(cc)) values versus weight-percent calcite estimates from paleosol carbonates of the Madera and Sangre de Cristo formations ...... 72

18 3.15 Plot of carbonate oxygen isotope (δ O(cc)) values versus weight-percent calcite estimates from paleosol carbonates of the Madera and Sangre de Cristo formations ...... 72

3.16 Plot of 1000ln versus temperature of carbonate formation ...... 88

13 3.16 Cross plot of stratigraphic position versus MAP estimates calculated from Δ C(cc-om) of paleosols from the Madera and Sangre de Cristo formations ...... 89

4.1 Map of Africa in its modern continental configuration illustrating positions of Upper Paleozoic glacial deposits and the Dekese borehole ...... 100

xiii

4.2 Relative stratigraphic position versus cumulative clay mineral percent of major constituents in the <2µm phyllosilicate fraction of Lukuga Formation samples ...... 108

4.3 Examples of common sedimentary structures and textures in the Luka Formation in the Dekese core ...... 115

4.4 Examples of polymicts collected from the Lukuga Formation in the Dekese core ...... 116

4.5 Examples of carbonate cements from the Lukuga Formation in the Dekese core ...... 117

4.6 Cross plot of the stratigraphic position versus δ18O values of calcite samples collected from the Lukuga Formation in the Dekese core ...... 121

4.7 Background-subtracted X-ray diffractograms of oriented aggregates from the <2µm fraction of a matrix sample collected from the Lukuga Formation at 961 m depth in the Dekese core ...... 122

4.8 Background-subtracted X-ray diffractograms of oriented aggregates from the <2µm fraction of a matrix sample collected from the Lukuga Formation at 1072 m depth in the Dekese core ...... 123

4.9 Background-subtracted X-ray diffractograms of oriented aggregates from the <2µm fraction of a matrix sample collected from the Lukuga Formation at 1072 m depth in the Dekese core ...... 124

4.10 Plot of relative probability versus U-Pb age of detrital zircons collected from sedimentary rocks of the Lukuga Formation in the Dekese core ...... 125

4.11 Plots of stable isotope composition and weight percent micritic calcite for samples from the Lukuga Formation in the Dekese core...... 126

xiv LIST OF TABLES

Table

2.1 White Rock Lake and White Rock watershed morphometric characteristics ...... 8

2.2 Carbon isotopic composition and weight percent organic carbon of standing watershed vegetation and soil organic matter ...... 19

2.3 Stable carbon isotope values of particulate organic matter (POM) filtered from White Rock Lake water ...... 19

2.4 Ranges of sedimentary organic matter isotopic and elemental characteristics and the relative proportions of carbonate present within four sediment response zones ...... 31

2.5 Calculated values of carbonate precipitated in isotopic equilibrium with lake waters. Water δ18O values ...... 38

2.6 Isotopic values of carbonate associated with White Rock Lake deposits ...... 38

3.1 Carbonate sample stratigraphic positions stable carbon and oxygen isotope values and weight percent calcite of analyzed samples from the Madera and Sangre de Cristo formations ...... 75

3.2 Stratigraphic position of samples and the corresponding stable carbon isotope values of organic matter sampled from the Madera and Sangre de Cristo formations of the Rowe-Mora basin ...... 76

4.1 List of lithologic and mineralogical characteristics of samples collected from the Lukuga Formation in the Dekese core ...... 108

xv DEDICATION

For my parents, who never doubted.

xvi

Chapter 1

INTRODUCTION

The work in this dissertation utilized multiple geochemical proxies to infer information about landscape-, regional-, and global-scale environmental change, and is divided into three parts. The first part in this volume (Chapter 2) identifies and correlates stratigraphic trends in the geochemistry of lake sediments with historical records of land use change and the onset of urbanization across the watershed of White Rock Lake, located in Dallas, TX, USA over approximately the past century. In the following chapters, lithostratigraphic, geochemical and mineralogical investigations of Upper Carboniferous-Upper Permian strata from paleo-equatorial

Cimmaron arch in north central New Mexico (Chapter 3) and strata from high-paleolatitude

Gondwanaland of the Congo basin (Chapter 4) are utilized to establish high-resolution (103-104 yr) deep-time paleoclimate records in order to refine the understanding of the relationship(s) between high- and low-latitude paleoclimate patterns during Permo-Carboniferous time. These three different studies demonstrate the utility of a multi-proxy approach to paleoenvironmental and paleoclimate reconstructions over different temporal and spatial scales.

Chapter 2 assesses the stratigraphic trends in stable carbon and oxygen isotope (13C and

18O) values, carbon:nitrogen (C/N) ratios, and concentration (wt. %) of organic and inorganic carbon contained within a ~2.5 meter-, ~100 year-long sedimentary record extracted from a core

of the shallow 4.4 km2 urban White Rock Lake in Dallas, Texas. Major changes in the 13C and

C/N ratios of organic matter preserved in the White Rock Lake sediments begin in the late

1940s, and are coincident with initiation of widespread watershed urbanization. The concomitant shifts in sedimentary organic matter 13C values, C/N ratios, and weight percent organic carbon and carbonate suggest that lake sediments preserve an increasing proportion of autochthonous organic carbon since urbanization commenced, and may be attributed to 1) a shift from C4- to C3-dominated vegetation across the White Rock Lake watershed, 2) a relative increase in organic matter contributed by primary production within the lake system and/or increased erosion and deposition of poorly developed soils and weathered bedrock from the up- stream watershed.

Chapter 3 presents lithostratigraphic, mineralogic, and stable isotope proxies of equatorial paleoclimate from Upper Pennsylvanian (Missourian-Virgilian) through Lower Permian

(Wolfcampian) strata of the Taos Trough, north-central New Mexico, USA. The Upper

Paleozoic sequence includes mixed marine carbonate and terrestrial clastic rocks in

Pennsylvanian strata (Madera Fm) which transition to terrestrial-dominated siliciclastic rocks in the Lower Permian (Sangre de Cristo Fm). Paleosol morphologies include eutric Argillisols in

Missourian rocks, Calcisols in Virgilian and Wolfcampian rocks, and gypsic subgroups in upper

Wolfcampian strata. X-ray diffraction analyses of pedogenic clay minerals show a shift from kaolinite, smectite, and illite assemblages representative of seasonal, sub-humid environments in the lower part of the stratigraphy to poorly-ordered 2:1 phyllosilicate and illite assemblages, representative of increased aridity and decreased seasonality, in the uppermost part of the stratigraphy. The stratigraphic distribution of these morphologies is suggestive of stepped

change from a sub-humid seasonal climate in Missourian time to a non-seasonal arid climate in end-Wolfcampian time.

There appears to be no significant stratigraphic trend among organic matter δ13C values, which range from ~-22.5 to -24 ‰. However, paleosol calcite δ13C values are ~-5 ‰ in

Virgilian strata, ~-8 in lower Wolfcampian strata, and ~-5 ‰ in topmost Wolfcampian strata.

These data suggest relatively high atmospheric PCO2 (or low biological productivity) during deposition of Pennsylvanian and upper Wolfcampian strata, and relatively low atmospheric

PCO2 (or high biological productivity) during deposition of lower Wolfcampian strata.

Missourian, Virgilian and Wolfcampian strata have pedogenic calcite δ18O values that range from ~-4 to -6.5 ‰. However, calcite δ18O values of the topmost Wolfcampian strata range from

+1 to +3 ‰. This stratigraphically abrupt, +5 to 7 ‰, shift in calcite δ18O values is consistent with predicted changes in tropical rainfall δ18O values during deglaciation of the ice-sheet in

Gondwanaland.

In chapter 4, approximately 1000 m of strata in Upper Paleozoic Lukuga Formation, preserved in the Dekese core taken from the central Congo basin, D.R. Congo provides lithostratigraphic, mineralogical, and isotopic evidence of substantial climatic variation within a long-lived lacustrine basin. Lithostratigraphic indicators of cold climate include polymictic strata and coupled laminations of fine, clay-size material and coarse silt. Dropstones are concentrated in the stratigraphic zones in the lower ~425 m of the Lukuga Formation, and varved strata occur in two broad stratigraphic zones in the lower ~700 m of the formation. These sedimentological indicators suggest that the lower two thirds of the Lukuga Formation was strongly influenced by frigid conditions and periglacial-like processes. The clay-size fraction of

the core samples is dominated by detrital minerals, including quartz, feldspar, chlorite, illite, and poorly-ordered 2:1 expansible phyllosilicates. Based on mineralogical variation within these samples, the Lukuga Formation is divisible into three Clay Mineral Zones (CMZs), numbered in ascending stratigraphic order. CMZ 1 and CMZ 3 include several horizons of expansible 2:1 phyllosilicates that represent warmer/wetter intervals. CMZ 2 is a zone dominated by chlorite and illite, and is interpreted as a continuous cold/frigid interval.

There are numerous calcite-cemented layers, including spar-filled veins, radiaxial fibrous cements, and micrite that is inferred to have crystallized near the time of deposition. Stable isotope analyses of micrite samples yield 13C values that range from -44.6 to -4.1 ‰ and 18O values that range from -20.0 to 5.0 ‰ (VPDB). The carbon data likely reflect a range of local carbon sources derived from bacterial activity and are unrelated to paleoclimatic conditions.

However, the stratigraphic trends defined by the oxygen isotope data suggest five or six intervals of frigid conditions conducive to glacial processes.

Each study contained in this work utilizes lithological, mineralogical, and geochemical indicators to develop a composite understanding of paleonvironmental conditions in a variety of contexts. The application of each indicator to terrestrial environments contributes to furthering our understanding of the timing of mechanisms associated with local, regional, and global-scale climate and environmental evolution.

Chapter 2

GEOCHEMICAL ANALYSIS OF LAKE SEDIMENTS AS A PROXY FOR WATERSHED URBANIZATION, WHITE ROCK LAKE, DALLAS, TX

Introduction

Lakes and reservoirs often preserve nearly continuous records of sedimentation, which potentially include periods of time for which comprehensive records of watershed conditions are lacking. They may also provide information about short-term processes that affect organic matter delivery and preservation (Meyers, 2003). Previous studies have demonstrated the utility of lake sediments as indicators of the effects of anthropogenic activity on lake systems, as well as changes in watershed conditions, vegetation, lake productivity, water quality, and regional climate (Buchardt and Fritz, 1980; Davis, 1990; Meyers, 1994; Grimm et al., 2000; Herczeg et al. 2001; Lindstrom and Hakanson, 2001; Sternbeck and Ostlund, 2001; Meyers, 2003; Routh et al. 2004; Diefendorf et al., 2007;). Urban lakes and reservoirs have the potential to yield high- resolution records of watershed conditions and the associated ecosystem response to urbanization. The type and abundance of organic matter preserved in lake sediments is a function of lake morphology, watershed topography, local climate, and the relative abundances of plant types present across the watershed (Meyers and Ishiwatari, 1993). Determination of the type of organic matter present within lake sediments is vital to reconstructing environmental conditions

and identifying changes in community-level sediment contributions and watershed vegetation.

Over the past century, the city of Dallas, the White Rock Creek watershed, and the man- made reservoir named White Rock Lake, have undergone a significant shift from a rural to urban landscape. White Rock Lake serves as a local catchment for watershed sediments that provides a continuous record of changes in watershed conditions over the past one hundred years. Previous studies of White Rock Lake sediment cores utilized major and minor element concentrations, grain size distribution, diatoms, pollen, and organochlorine compounds preserved in the lake sediments as indicators of water quality trends in the influent White Rock Creek, which were attributed to increased watershed urbanization (Van Metre and Callender, 1997; Bradbury and

Van Metre, 1997). This study documents stratigraphic trends in the geochemical composition of organic matter and calcite preserved in the White Rock Lake sedimentary record and the current lacustrine environment that collectively records community-level changes in sediment supply and lake productivity. Carbon isotopic and C/N elemental analyses of bulk sedimentary organic matter from the White Rock Lake core define four episodes of lake sediment response which correlate with historical records of urbanization and changing land use practices across the watershed.

Geologic Setting

White Rock Lake is a 4.4 km2 reservoir located on White Rock Creek, in northeastern

Dallas, TX. The primary tributary to the reservoir is White Rock Creek, a perennial gaining stream with flow derived from ground water and overland flow from a catchment spanning 215

km2 with a maximum relief of 134 meters above lake level. The lake lies within the upper sub- basin of the White Rock Creek watershed (Figure 2.1), which is currently 95% urban, as compared to 13% urban in 1961, and completely rural in 1912, when the reservoir was initially filled (Ogle, 1956; Ollman, 1969; NCTCOG, 2002). White Rock Creek cuts through up to 3 meters of the Upper Cretaceous Austin Group, a sequence of interbedded chalks and marls with several bentonitic strata. The Austin chalk has a maximum thickness in the city of Dallas of 167 meters, and is composed almost entirely of micritic calcite fossils, especially coccoliths, and contains common phosphate and pyrite nodules. Across the White Rock Creek watershed, the

Austin chalk is covered by up to 4.5 meters of sandy clay to clay-rich alluvium and black clay prairie soils (Allen and Flanigan, 1986).

Cultivation of White Rock Lake watershed began around 1860; consequently, agricultural practices contributed to rapid erosion and degradation of the landscape (Marshall and

Brown, 1939). Poor soil conservation practices resulted in significant erosion, reducing the topsoil across the landscape to one-half its pre-cultivation thickness by 1939 (Marshall and

Brown, 1939). The pre-agricultural landscape was characterized by treeless prairies, dominated by C4 tall grasses; big bluestem (Andropogon gerardi) and little bluestem (Schizachyrium scoparium), Texas needlegrass (Nassella leucotricha), and Indian grass (Sorghastrum nutans).

Trees were prevalent near streams and drainage ways, and included Bois d'arc (Maclura pomifera), American elm (Ulmus Americana), cedar elm (Ulmus crassifolia), honey locust

(Gleditsia triacanthos), and pecan (Carya illinoinensis), which all utilize the C3 photosynthetic pathway. Historically, agricultural crops have been a mixture of C3 and C4 plants, including cotton, corn, and other small grains (Marshal and Brown, 1939).

Construction of White Rock Lake began in 1910, and water was first diverted for municipal use in 1912 (Marshall and Brown, 1939). The reservoir served as a municipal water supply from 1912 until 1964; however, the lake has been used solely for recreation since 1964.

At the time of core collection, White Rock Lake had a maximum depth of approximately 6.7 m and an average depth of 3 m (TPWD, 2003). Lake depth has varied over time, being the deepest in 1939 at approximately 10 m (Patterson, 1941). Morphometric and watershed characteristics are listed in Table 2.1 (TWDB, 2003).

Table 2.1. White Rock Lake and White Rock watershed characteristics (TWDB 2003).

Characteristic

Conservation pool elevation (MSL) 139 m

Surface area 4.4 x 106 m2

Mean depth 3 m

Max depth 6.7 m

Volume 1.11 x 107 m3

Watershed area 1.75 x 108 m2

Watershed area/water area 34

Multiple natural and man-made events have occurred over the history of the lake which have had the capacity to affect the lacustrine sedimentary record. The watershed area has

suffered multiple floods and periods of prolonged drought. Extreme flooding took place in 1942,

1949, and 1962 (Gilbert 1963). Prolonged drought conditions persisted from 1950 to 1957, and a minor drought episode occurred from 1995 to 1996.

Five dredging events occurred prior to the sediment core collection; however, none of the dredging activity affected the area of the lake sampled in the core. The first dredging took place from 1936 to 1941 at the northern end of the lake and resulted in the removal of approximately

446,000 m3 of sediment. A subsequent dredging, via dragline, in 1955-1956 removed an estimated 11,740 m3 of material. A third and minor episode of sediment removal took place in

1970 during a severe drought, as area residents removed approximately 31 m3 of exposed sediment from the northern shore of the lake (Van Metre and Callender, 1997). The fourth dredging endeavor occurred over the course of 7 months in 1974 at the northern end of the lake and resulted in the removal of 1.03 x 106 m3 of sediment (Kelley, 1994). The most recent episode of sediment removal occurred from January to September of 1998, resulted in the removal of approximately 2.3 x 106 m3 of material, and brought the lake to a minimum depth of

2.4 m (Ostdick, 2007).

The White Rock Lake sediment core that is focus of this study was collected by the

USGS in September 2003 using a Benthos gravity corer with a 6.5 cm diameter barrel at a location approximately 400 m north of the dam in a water depth of 5 m (Figure 2.1; Van Metre et al., 2004). Sediments predating construction of the reservoir were penetrated at the base of the

250 cm-long core.

Figure 2.1. Map of Texas highlighting the White Rock lake location, subdivisions of the White Rock Creek watershed, and county boundaries. The starred site marks the core sample location.

Methods

Sampling

White Rock Lake sediment core samples were given to SMU by the USGS in 2005. The samples had been freeze-dried, then stored in HTPE containers at USGS soon after collection.

From this White Rock Lake core (WR3), samples were received at 3 cm intervals from the core- top to 30 cm depth, at 5 cm intervals from 30 cm to 150 cm depth, and 10 cm intervals from 150 cm to 250 cm depth. Activities of 137Cs were measured by the USGS following the methods of

Van Metre and others (2004), and are the basis for the age model used herein to date the sediments of the White Rock Lake core. Bulk sediment core samples were separated via sieving into bulk, coarse (>0.5 mm), intermediate (<0.5 - >0.149 mm), and fine (<0.149 mm) grain size fractions.

Plant samples were collected from the lake and creek margins, as well as from various locations across the watershed in order to determine the δ13C values of standing vegetation within the watershed. Water samples from White Rock Lake were collected periodically, from

November 2010 to October 2011, in clean 1 L HTPE bottles in order to isolate, then determine the δ13C values, of particulate organic matter in the water column (POM). Three soil samples were collected from the upper 4 cm of soil profiles at locations proximal to White Rock Lake in order to determine the δ13C values of present-day pedogenic detritus that may be contributed to the White Rock Lake sedimentary record via erosion, transport, and deposition.

Multiple size fractions of lake sediments were imaged using scanning electron microscopy (SEM) at SMU in Dallas, TX. Attention was given to determining the size and

morphology of calcite grains, as well as identifying the occurrence of microfossils such as coccoliths and diatoms. The White Rock Lake sediment core samples were compared to SEM images taken from the underlying Austin Chalk in order to distinguish between authigenic and detrital calcite phases in the White Rock Lake sedimentary record.

Stable isotopes

Sedimentary organic matter (SOM) was isolated from bulk samples of the White Rock

Lake sediment core via chemical pretreatment. Each sample was flushed with successive aliquots of concentrated (12.1N) HCl to remove inorganic carbon, repeatedly rinsed with de- ionized water until reaching a pH of ~5.5, then freeze-dried. Bulk soil samples were chemically pretreated with aliquots of 1.0N HCl in order to remove carbonate and were washed with successive rinses of de-ionized water to remove acidity. Acid-treated soil samples were sieved to isolate the less than 0.125 mm grain size. Samples of current watershed vegetation, collected from the lake margin, the riparian corridor, and several sites across the watershed, were dried overnight at 50 ºC to drive off excess water, then analyzed for stable carbon isotopic composition.

One-liter lake water samples were filtered through 0.45 µm diameter glass fiber filters.

The accumulated residue was treated with 10% HCl to remove carbonate, repeatedly rinsed with de-ionized water until a pH of ~5.5 was reached, then dried at 50 ºC. The sample-coated filter membranes were pyrolyzed and analyzed for stable carbon isotopic composition of particulate organic matter (POM).

Pyrolysis of organic matter samples followed the methods of Boutton and others (1991).

Samples were sealed in individual Vycor tubes with approximately 1 g copper oxide and 0.5 g of copper metal. The sample tubes were incrementally heated, initially to 900 ºC for 2 hours, stepped down 2 º/min, until reaching 650 ºC. Samples dwelled at 650 ºC for two hours, and were then cooled to room temperature in the furnace for a period of ~12 hours. The evolved CO2 was cryogenically purified using high-vacuum extraction lines. Yields of CO2 were determined (+/-

0.3 µmol) by mercury manometry and were used to calculate the weight percent organic carbon amongst the analyzed samples.

An aliquot of each size fraction of the WRL sediment samples was processed for determination of calcite δ13C and δ18O values. Between 50 and 100 mg of sample was reacted overnight with 100 % anhydrous phosphoric acid at 25 °C (McCrea, 1950). Reaction products were cryogenically purified using high-vacuum extraction lines, and CO2 yields were determined via mercury or baratron manometry. All stable isotope analyses were determined at Southern

Methodist University using a Finnigan-MAT 252 IR-mass spectrometer. The isotope values are reported in standard delta notation:

δ = (Rsample/Rstandard -1)*1000 where R is the ratio of heavy-to-light stable isotope present in the sample or standard, and δ- values are reported relative to the Peedee Belemnite (PDB) standard (Craig, 1957).

C/N ratios

Samples of standing vegetation were dried overnight at 50 ºC prior to C/N analysis.

Organic matter samples from lake water and the sediment core were treated with 10% HCl to remove inorganic carbon, repeatedly rinsed with de-ionized water, then dried at 50 ºC.

Approximately 2 to 5 mg of acid-treated samples were loaded into tin capsules and combusted in order to determine the C/N ratios of watershed vegetation, lacustrine sedimentary organic matter

(SOM), and the particulate organic matter (POM) of lake water using a Costech ECS 4010 elemental analyzer attached to a Finnigan MAT 253 IR-mass spectrometer at Southern Methodist

University.

Results

Chronology

The USGS determined the chronology of sediments in the White Rock Lake core based on the preserved 137Cs profile using the methods outlined in Van Metre and others (2004; Figure 2.2).

Sediments predating the construction of the reservoir were penetrated at 250 cm depth. Thus, the base of the core is assigned the reservoir filling date, C.E. 1912. The top of the core is assigned the date of C.E. 2003, the year of core collection. The date of 1952 was assigned to the depth of

165 cm based on the increase of 137Cs above background levels, and the date of 1963 was assigned to the depth of 132.5 cm based on the peak abundance of 137Cs (Figure 2.3). Mass accumulation rates (MAR) were calculated as the mass of dry sediment per unit area of an interval of a core, divided by the time the interval represented, and utilized to interpolate dates of

samples at intervals between depths dated by 137Cs concentrations (Figure 2.4)(Van Metre et al.,2004). For the periods from 2003 to 1963, 1963 to 1952, and 1952 to 1912, mass accumulation rates were calculated as 1.55, 1.68, and 1.44 g/cm2yr, respectively.

Organic matter

13 Comparison of  C(SOM) values with corresponding C/N ratios from 24 samples, 2 per depth from 12 depths, in WR3 is shown in figure 2.5. Sedimentary organic matter 13C values ranged from -24.6 to -15.8‰, and C/N ratios varied from 27.2 to 6.8, with average values between 21.3 and 8.0 (Figure 2.6). Carbon isotopic values record a shift toward more negative values beginning around 1952 (Figure 2.6a). C/N ratios trend towards lower values after approximately 1966. Weight percent organic carbon values increase from 2.0 to 4.5% from the base to the top of the core; organic matter becomes increasingly abundant in sediments deposited after ~1987 (Figure 2.6c). Weight-percent sedimentary organic matter values determined from elemental analysis for nitrogen range from 0.07 % to 0.42 %, and weight percent organic carbon values range from 1.00 % to 3.25 % (Figure 2.7a, b). These values differ from those calculated from offline pyrolyzation and subsequent mercury manometry readings and will not be considered beyond a generalized assessment of the overall trends present in the dataset through the White Rock Lake core. The nitrogen weight percent values are tightly constrained in samples predating 1978 and carbon values exhibit more scatter over this interval (Figure 2.7).

137Cs (pCi/g)

2010

2000

1990

1980

1970

1960

1950

1940

1930

1920 0.0 0.5 1.0 1.5 2.0

Figure 2.2. Dated 137Cs profile of WRL sediments as determined by Van Metre and others (2004). The reservoir filling date, 1912, is assigned to the core base. The depth of 165 cm is assigned a date of 1952, based on increased 137Cs over background levels, and 1963 is assigned to a depth of 132.5 cm based on the peak abundance of 137Cs. The top of the core is assigned the core collection date of 2003.

-25

-50

-75

-100

-125

-150

-175

-200 Depth below sediment Depthbelow sediment (cm) top -225

-250 2010 1990 1970 1950 1930 1910 Year

Figure 2.3. Age model of White Rock Lake sediment core based dates assigned to top (2003) and bottom (1912) of core, peak 137Cs concentration in 1964, and an increase over background levels in 137Cs in 1952.

0 -25 -50 -75 -100 -125 -150 -175 -200

-225 Depth below sediemnt Depthbelow sediemnt (cm) top -250 2006 1998 1990 1982 1974 1966 1958 1950 1942 1934 1926 1918 1910 Year Figure 2.4. Age model showing approximate age of each sample. Approximations are based on mass accumulation rates for intervals between 137Cs dated depths. 17

Samples of standing vegetation from the modern watershed surface record 13C values ranging from -13.8 to -29.7‰, and bulk soil organic matter 13C values are approximately -18‰

(Table 2.2). The weight percent organic carbon values of vegetation samples ranged from 37 to

53%, and soil values ranged from 3 to 5% (Table 2.2). Particulate organic matter samples filtered from White Rock Lake water have 13C values of ~ -21‰ (Table 2.3).

Calcite abundance and stable isotope values

Calcite 13C values range from -2.1 to 0.9‰ and 18O values range from -4.8 to -0. 9‰,

(Figure 2.8). In all calcite samples, the fine (<0.149 mm) fraction 13C values are slightly more negative, by up to 0.5 ‰, than the coarsest (>0.5 mm) fraction, (Figure 2.9a). Calcite 18O values of the coarsest fraction have more negative values than the coarse (<0.5 - >0.149 mm) or fine fractions. While there appears to be no significant stratigraphic trend in calcite 13C values, the 18O values of the fine fraction shift slightly toward more positive values with depth, from -

4.0‰ at the top to -2.4‰ at the base of the core (Figure 2.9 a, b). Weight percent calcite values show a stratigraphic trend of decreasing values with depth in all size fractions; however, in all samples the fine size fraction values are higher than the corresponding bulk and coarse fractions at each depth (Figure 2.9c).

Table 2.2. Carbon isotopic composition and weight percent organic carbon of standing watershed vegetation and soil organic matter.

Vegetation Sample mass (mg) wt% C 13C (PDB)

Southern wood fern leaf 5.00 38.4 -29.5

Southern wood fern leaf (repeat) 4.53 36.9 -29.7

Southern wood fern stem 5.20 44.0 -28.6

Water oak leaf 4.95 46.3 -28.5

Bermudagrass 3.9 41.8 −

Eastern red cedar leaf 3.9 53.0 −

Crabgrass 6.2 41.4 −

Pecan leaf 2.9 45.7 −

Soil OM Sample mass (mg) wt% C 13C (PDB)

A - Tietze Park 22.9 3.2 -18.1

B - WRL Rowing Dock 22.1 3.0 -18.8

C - Moss Park 21.4 5.0 -18.1

Table 2.3. Stable carbon isotope values of particulate organic matter (POM) filtered from White Rock Lake water.

Water POM 13C (PDB) RD4 − RD6 − RD7-2 − RD7-1 −

Carbonate morphology

SEM images of sieved White Rock Lake core samples show that the fine, coarse, and bulk fractions are visually indistinguishable from one another, and that the bulk and coarse fractions are primarily comprised of aggregates of finer-grained material. Each size fraction contains sub-hedral calcite grains, coccolith fragments, and sparse non-calcareous diatom tests.

Based upon these observations no significant distinction could be made between authigenic and detrital calcite amongst the size fractions from the analyzed samples.

13 Figure 2.5. Plot of C/N ratios and δ C(SOM) values representative of C4 (brown), C3 (green), and algal (blue) contributions to organic matter preserved in lacustrine sediments (after Meyers and Lallier-Vergès, 1999).

vascular

alues indicative indicative alues

dominated system with strong non with strong system dominated

values; b) C/N ratios; c) weight percent carbon present of sedimentary present percent sedimentary of weight carbon c) C/N b) ratios; values;

SOM

dominated and terrestrial deposits. Yellow region indicates deposits. v terrestrial region Yellow and dominated



deposition. Blue region represents C represents region Blue deposition.

er within samples analyzed from the White Rock Lake core. Stratigraphic shift from Stratigraphic from shift core. the from White er Rock analyzed withinsamples Lake

dominated organic matter deposited in WRL setting. Brown area represents values values area setting. represents in deposited Brown matter WRL organic dominated

a) a)

to algal to

Figure 2.6. 2.6. Figure matt organic C with C consistent mixedC component.

Figure 2.7. Depth from surface for samples analyzed from the White Rock Lake core is shown on the left-side y-axis, whereas the corresponding age estimates for sedimentation as calculated from the 137Cs age model based upon Van Meter and others (2004) are shown on the right-side y- axis. Tie lines across the y-axes show the boundaries between different depositional episodes within the White Rock Lake core. The x-axis of the left-side graph shows the measured weight- percent nitrogen value from sedimentary organic matter amongst the samples analyzed, whereas the x-axis of the right-side graph shows the corresponding weight-percent carbon value from sedimentary organic matter amongst the samples analyzed. Note that each sampling depth has two data points and these represent the results of two separate analyses from two aliquots of the same treated fractions from the White Rock Lake core. See text for discussion.

0.0

-0.2

-0.4

-0.6

-0.8

Ccc (PDB) Ccc -1.0

13 

-1.2 Bulk -1.4 Coarse Fine -1.6 -5.0 -4.0 -3.0 -2.0 -1.0 0.0 18  Occ (PDB)

13 Figure 2.8. Plot of carbonate carbon isotope (δ C(CC)) vs carbonate oxygen isotope 18 (δ O(cc)) values from sieved fractions taken from the White Rock Lake core. Coarsest- fraction samples (>500 m) are denoted by triangle-shaped data points, Coarse-fraction samples (500 m to >125 m) are denoted by square-shaped data points, and Fine- fraction samples (<125 m) are denoted by circle-shaped data points. See text for discussion.

axis,

percent calcite calcite estimates percent

shaded vertical space corresponds to the corresponds shadedvertical space

Rock Lake core sieved sample fractions. See See fractions. sieved core sample Rock Lake

); the gray the );

(cc)

cles) sample fractions correspond to sieved fractions of of correspond fractions sieved to sample fractions cles)

(cir

2.8. (a) Measured carbon isotope isotope (a) 2.8. compositionMeasured carbon carbonate of

lyzed from the White Rock Lake core is is core the on shown the lyzed y from White Rock Lake

which were reported in reported Weight were Collins which (c) (2013).

samplesana

(squares) and (squares)and Fine

reaction of carbonate within White within carbonate of reaction

O values O

m respectively Fig in as respectively m

carbonate carbonate

(triangles), Coarse (triangles),

m, and <125 <125 and m,

age AustinChalk

Bulk

yields derived from H derived from yields

2.7

). (b)Measured oxygen isotope isotope ( oxygen compositioncarbonate of (b)Measured ).

m, 500 to >125 to 500 m, >125

(cc)

Figure 2.9. Depth from surface for surface Depth 2.9. for from Figure within different episodes between boundaries show Lake White the Rock depositional graphs across lines tie three the whereas in as Figcore >500 ( Cretaceous of range CO upon based textdiscussion. for 25

Discussion

Plant detritus from terrestrial and lacustrine sources can be differentiated into two biochemically distinct categories; vascular and nonvascular photosynthesizers. Vascular plants, such as trees, shrubs, and grasses, have woody tissue and cellulose and grow on land and in shallow waters. Nonvascular photosynthesizers lack woody and cellulosic tissue and include simple algae and mosses. Most vascular plant material in lacustrine settings is allochthonous, while non-vascular matter is autogenic.

C/N ratios and stable carbon isotope compositions of organic matter preserve original information about the types and relative abundances of primary producers that contribute organic carbon to an ecosystem, as well as to the sedimentary record of that ecosystem. Nonvascular aquatic plants have low C/N ratios, typically between 4 and 10 (Meyers, 1994). Vascular land plants have higher C/N ratios, most often greater than 20 (Meyers, 1994). As a result of these differences, C/N ratios have been used to distinguish between algal and terrestrial vascular plant origins of sedimentary organic matter (Ishiwatari and Uzaki, 1987; Jasper and Gagosian, 1990;

Meyers, 1994; Silliman et al., 1996; Kendall et al., 2001; Perdue and Koprivnjak 2007), and changes in C/N ratios of lacustrine organic matter have been interpreted to represent changes between periods dominated by terrestrial and algal organic matter contributions to lake sedimentary records (Guilizzoni, et al., 1996; Kaushal and Binford, 1999).

The stable carbon isotope composition of organic matter in lacustrine sediments can also be used to distinguish between contributions from different types photosynthetic pathways

(Schell and Barnes, 1980; Meyers 1994). The carbon isotope composition of organic matter primarily results from the isotopic fractionation associated with photosynthetic carbon fixation 26

within the organism and may be used to distinguish between organisms that utilize C3, C4, and

CAM photosynthetic pathways (O'Leary, 1988). Most terrestrial vascular plants utilize the C3

13 photosynthetic pathway and preferentially discriminate against C. In vascular land plants, C3 photosynthesis results in ~ -20‰ shift in δ13C values from the inorganic carbon source,

13 atmospheric CO2, resulting in modern δ C values of ~ -27‰ ± 3‰ (PDB) (O’Leary, 1988). C4 photosynthesis results in a δ13C shift of ~ -7‰ from the inorganic carbon source, resulting in modern δ13C values of ~ -14‰ ± 3‰ (PDB) (O’Leary, 1988). Trees, most shrubs, herbs, cool- weather grasses, and aquatic plants utilize C3 photosynthetic pathways; whereas C4 photosynthesis is primarily used by heat-tolerant grasses. Aquatic plants and algae are C3 photosynthesizers that utilize dissolved CO2 in isotopic equilibrium with atmospheric CO2 as

13 their inorganic carbon source. As a result, these organisms have δ C values isotopically indistinguishable from the terrestrial C3 plants of the adjacent watershed (O'Leary, 1988; Meyers and Lallier-Vergès, 1999; Sharp, 2007). Therefore, the C/N ratios and δ13C values of lacustrine organic matter potentially provide a means of identifying the contributions of terrestrial and algal organic components preserved within a lacustrine sedimentary record.

Determination of the type of organic matter preserved in the lake system aids in understanding the stability of the landscape across the watershed and historical changes in lake trophic levels. Larger proportions of allochthonous organic material preserved in lake sediments can indicate increased soil erosion and detrital transport, or unproductive lake ecology, while larger proportions of autogenic algal components may suggest transport of less detrital organic matter into the lake from the surrounding watershed. Furthermore, increases in the relative abundance of autogenic organic matter can suggest periods of increased primary productivity

and/or eutrophic conditions (Schelske and Hodell, 1991; Guilizzoni, et al., 1996; Routh et al.,

2004).

Preservation of isotopic and chemical ratios

Sediment trap studies have indicated that organic matter concentration decreases dramatically between lake surface and bottom sediments, but C/N ratios and 13C values of the source organic matter are well preserved (Ertel and Hedges, 1985; Meyers and Eadie, 1993).

Previous studies have documented the preferential mineralization of nitrogen during early burial diagenesis of organic matter in sediments, and algal organic material has been shown to be more susceptible to partial degradation than vascular plant material (Herczeg, 1988; Meyers, 1997;

Herczeg, et al., 2001). Alteration of organic matter in lake sediments typically results in an increase in C/N ratios as diagenesis progresses, but the C/N ratios are generally preserved well enough so that primary source information can be determined (Meyers, 1997).

Previous stable carbon isotope studies have demonstrated that selective diagenesis in bulk sedimentary organic matter may decrease 13C values by up to ~2‰ (Burchardt and Fritz, 1980;

Meyers and Eadie, 1993). However, Meyers and Ishiwatari (1993) have found no evidence for alteration of original organic matter 13C values in lacustrine sediment samples with less than a few weight percent organic carbon. The average weight percent carbon values from WR3 core range from 2.0 to 4.5%; therefore, the C/N ratios and 13C values of the acid-treated samples from the core are considered to represent primary values reflective of the balance of organic matter contributed to the lacustrine sedimentary record over the past century.

Organic matter stable isotope and elemental data

The proportions of autochthonous and allochthonous material present in the lake sediments can be discerned by comparing the δ13C values and C/N ratios of sedimentary organic matter with the typical values associated with each end member. The average ranges of δ13C values and C/N ratios associated with terrestrial C4 and C3 vascular plants, and non-vascular aquatic photosynthesizers are delineated in Figure 2.5. The values of the WRL samples are presented in the context of C3, C4, and algal components in Figures 2.5, and 2.6. Terrestrial C3 plants sampled from White Rock Lake watershed had average 13C values of -29.4‰ (Table

2.2). Particulate organic matter sampled from the lake water had an average 13C value of -21‰

(Table 2.3), and this value is taken to represent the algal component of the sedimentary organic carbon record.

Measured sedimentary organic carbon 13C values shift by ~9‰, from ~-16‰ in oldest

13 sediments to ~-25 ‰ in 2003. The  C(SOM) values of sediments deposited prior to approximately 1959 (below a depth of 146cm) vary over a narrow range of 1.6‰, from -17.4‰

13 to -15.8‰. In sediments deposited after 1959,  C(SOM) values record a trend toward more

13 negative values with decreasing age, from -18.5‰ in 1962 to -24.6‰ in 2003. These  C(SOM) values suggest significant contributions from C4 vegetation to the total sedimentary organic carbon record prior to 1959, with a steadily increasing proportion of material derived from C3 photosynthesizing organisms (either vascular or algal) after 1959.

Sedimentary organic matter C/N ratios range from 27.2 to 6.8 with average values ranging between 21.3 and 8.0 (Figure 2.6). The C/N ratios of sediments deposited prior to about

1971 have average values ranging from 21.3 to 13.5, and individual sample C/N ratios ranging from 27.2 to 10.5 (Figure 2.6b). The large spread between replicate sample C/N values within this interval likely reflects heterogeneity within each sample (Figure 2.6b). Sediments deposited after 1971 have more tightly constrained C/N ratios ranging from 11.9 to 8.1, and appear to be more homogenous with respect to source material than the underlying deposits (Figure 2.6b).

Additionally, the weight percent nitrogen preserved in the White Rock Lake core sedimentary organic matter increases from pre-1959 values of <0.10% to values ranging up to 0.42% after

1966, with an inflection point in approximately 1992 (Figure 2.7). After 1992, the average weight percent nitrogen increases significantly from 0.15% to 0.38%, more than doubling over a period of eleven years. Similarly, weight percent carbon shifts towards higher values from 1.50% to 3.30% over the same, post-1992, time interval (Figure 2.7). The C/N ratios reflect a shift from sedimentary organic matter that was dominated by detrital, terrestrial-dominated vascular plants prior to 1966, to a more authigenic, lake-derived, algal-influenced system by 1982 (Figure 2.6).

The WRL sedimentary organic matter 13C and C/N ratios indicate a shift from a watershed and catchment characterized by C4 terrestrial vegetation to a system dominated by allochthonous and autogenic C3 photosynthesizers. Four zones of lake sediment response were identified from the WRL core samples based on the 13C values and the C/N ratios of preserved sedimentary organic matter. Zone I is comprised of sediments deposited between 1912 and

1952, collected from -215 cm to -165 cm depth (Figure 2.6). This interval is characterized by

13  C(SOM) values that vary over a narrow range, from -16.7‰ to -15.8‰. Individual sample C/N

ratios range from 19.3 to 26.3, with average values ranging from 17 to 26. The large spread between replicate sample C/N values within this interval likely reflects organic matter heterogeneity within each sample. Weight percent organic carbon values vary from 2.0 to 2.4 percent (Figure 2.7), and weight percent carbonate varies from 13.7 to 15.3% (Figure 2.9). The isotopic and elemental values of Zone I deposits are reflective of lacustrine organic matter dominated by terrestrial C4 plants (Figure 2.5), which coincides with widespread agricultural land use practices in place across the watershed at the time (Figure 2.10).

Table 2.4 Ranges of sedimentary organic matter isotopic and elemental characteristics and the relative proportions of carbonate present within four sediment response zones.

13 Zone δ C(SOM) C/N Wt% C(SOM) Wt% CaCO3

I (1912-1952) -16.7 to 15.8 17 to 26 2.0 to 2.4 13.7 to 15.3

II (1952-1966) -18.6 to 17.4 14 to 21 2.0 to 2.7 19.2 to 23.5

III (1966-1982) -20.3 to -18.2 10 to 18 2.7 to 3.0 23.8 to 30.7

IV (1982-2003) -24.6 to -21.4 8 to 10 2.7 to 4.0 31.4 to 35.7

Comparison of percent landscape utilized as farmland and change in urban population from early 1900s to early in to 1900s early change from population and farmland landscape Comparisonearly urban as utilized percent of

Figure 2.10 Land use change over time. Collin County data in orange, Dallas County data in blue, and average of the two two blue, in the of data andDallas County orange, in average over data Collin change 2.10 time. use County Land Figure gray. in 2000s.

Zone II deposits reflect a transitional period of landscape development across the lower watershed, which lasted from 1952 to 1966. Zone II sediments were deposited between depths

13 of -155 cm to -128 cm (Figure 2.6), and have  C(SOM) values ranging from -18.6 ‰ to -17.4 ‰, individual sample C/N ratios range from 11.5 and 27.1, with average C/N ratios varying from

13.4 to 21.3, and weight percent organic carbon values shifting from 2.0 to 2.7 percent (Figure

2.7). The average weight percent carbonate present in Zone II samples varies between 2.7 and

13 2.0 percent (Figure 2.9). The stratigraphic ~-2 ‰ shift in  C(SOM) values within Zone II (Figure

2.6) reflects an increasing importance of C3 plants on the landscape (Figure 2.5). In Dallas county, the aerial extent of agricultural land usage decreased from 65 percent in 1950 to 45 percent in 1964, while Collin county land usage remained constant (Figure 2.10). Urban development across Dallas county contributed to the proliferation of C3 plants across the lower portion of watershed during the late 1950s and early 1960s, stabilization of the floodplains and uplands which previously were utilized for agriculture, and a source of detrital sedimentary organic matter with C4 signatures, and proliferation of authigenic, algal derived sedimentary organic within White Rock Lake. All of which appear to have contributed as important and separate sources of sedimentary organic matter in the Zone II portion of the White Rock Lake core based upon stable isotope and elemental ratio characteristics of the samples analyzed.

The sediments of Zone III were deposited between 1966 and 1982, from a depth of -128

13 cm to -78 cm (Figure 2.6), and are characterized by  C(SOM) values which decrease sharply from -18.2 ‰ to -20.3 ‰, individual sample C/N ratios that decrease from 16.7 to 9.9, averaged sample C/N ratios which decrease from 13.5 to 10.5, and weight percent organic carbon values that increase from 2.7 to 3.0 percent (Figure 2.7). The average weight percent carbonate present

in Zone III samples varies between 23.8 and 30.7 percent (Figure 2.9). The C/N ratios of replicate samples from Zone III sediments are more tightly constrained than in older deposits and reflect a rise in non-vascular autogenic organic matter contributions to the catchment (Figures

2.5, 2.6). Beginning around 1970, the northernmost portion of the watershed experienced accelerated urbanization, as agricultural land usage decreased from 88 percent in 1969 to 65 percent in 1982. Similarly, the lower watershed continued to experience increased urbanization as Dallas county farmland usage decreased from 45 percent in 1969 to 34 percent by 1982

13 (Figure 2.10a). The  C(SOM) values and C/N ratios of corresponding lake sediments indicate an increased relative contribution of terrestrial C3 plants and non-vascular C3 photosynthesizers, and decreased detrital C4-derived organic matter from the uplands to the sedimentary organic matter within the White Rock Lake core sediments during Zone III deposition (Figure 2.5).

Zone IV deposits date from 1982 to 2003, were collected from -78 cm to the top of the

White Rock Lake core (Figure 2.6) and suggest lake sediments are dominated by a C3-

13 photosynthesizing algal component. Zone IV sediments have  C(SOM) values which range from

-21.4 ‰ to -24.6 ‰ (Figure 2.6), which are consistent with the average carbon isotopic values for C3 plants. The individual sample C/N ratios of Zone IV are tightly constrained between 6.8 and 11, with sample averaged values between 8 and 10 (Figure 2.6). The weight percent of organic carbon increases stratigraphically from 2.7 percent at the base of the zone to 4.0 percent at the top (Figure 2.7). All three indicators, are suggestive of an increasingly important non- vascular primary production component present in the lake sediments (Figure 2.5). Both Collin and Dallas counties saw a decline in the rate of urbanization between 1982 and 2003 (Figure

2.10). Lake sediments from Zone IV are dominated by aquatic autotrophs, with a decreased

contribution from C4 plants as compared to the lower zones. The weight percent organic carbon present in Zone IV sediments may be attributed to a rapid increase in primary productivity and algal-derived sedimentary organic matter within White Rock Lake, and also as a result of decreased post-depositional organic matter degradation in these younger sediments.

13 Figure 2.6 demonstrates the relationship between  C(SOM), C/N ratios, and weight

13 percent C(SOM) of the lake sediments. Sediments predating 1952 have  C(SOM) values indicative

13 of C4-dominated deposition while sediments deposited between 1952 and 1982 have  C(SOM)

13 values suggestive of mixed contributions from C3 and C4 photosynthesizers. The  C(SOM) values of sediments deposited after 1982 are indicative of a C3 dominated system. The C/N ratios of

WRL organic matter reflect a similar trend. Sediments deposited prior to 1982 (-75 cm) have

C/N values consistent with mixed algal and vascular plant material, and post 1982, the C/N values reflect a dominantly algal contribution. The increase in the percentage of nitrogen preserved in the organic lake deposits is coincident with population and land use changes in the upper watershed and is likely an additional direct response to watershed urbanization and activities associated with changes.

Weight percent organic carbon increases from ~2.2 % at the base to ~4.5 % at the top of the core (Figure 2.7). The stratigraphic increase in weight percent organic carbon may reflect that recently deposited material has not undergone microbial degradation, or it may represent increased contributions of autochthonous organic carbon to the lake system during more recent times. The latter scenario appears to be supported by corresponding changes in C/N ratios suggesting increased algal contributions of sedimentary organic matter toward the top of the

White Rock Lake core sediment samples (Figure 2.6). 35

13 When considered together, the stratigraphic shifts in δ C(SOM), C/N ratios, weight percent nitrogen, and weight percent carbon suggest a relative increase in organic matter contributed by

C3 photosynthesizing aquatic autotrophs beginning around 1952 and a more significant shift starting in 1966 (Figure 2.6). An increase in weight percent organic carbon is observed at each of these transition points (Figures 2.6, 2.7). The stratigraphic increase in weight percent nitrogen near the top of core is an additional indicator of increased primary productivity.

A further indication of the transition from terrestrial C4-dominated to algal-dominated deposits can be discerned from the pollen record of White Rock Lake. Prior to 1940, lake deposits were dominated by agricultural weeds, with a significant contribution from

Chenopodiacea + Amaranthus, which utilizes the C4 photosynthetic pathway (Bradbury and Van

Metre, 1997). After approximately 1948 (upper Zone I, this study), the pollen record is dominated by C3 photosynthesizing trees and shrubs. The abrupt transition in the pollen record corresponds to the onset of a shift across the lower watershed from an agricultural landscape characterized by native C4 vegetation and agricultural weeds to an urban landscape dominated by

C3 vegetation (Figure 2.6). Additionally, the initial presence of diatoms in 1986 (Bradbury and

Van Metre, 1996) correlates with accelerated urbanization and population of the upper White

Rock Lake watershed and coincides with the transition from sediment response Zone III to Zone

IV as previously described (Figures 2.6, 2.10).

Carbonates

Additional responses to urbanization may be deduced from changes in the weight percent

18 values of carbonate in samples from the White Rock Lake core. The measured  Occ values of the fine fraction range from -1.8 ‰ to -4.0 ‰, the coarse fraction from, -0.9 ‰ to -3.1 ‰, and the bulk fraction from -2.8 ‰ to -4.5 ‰ (Figure 2.8). The isotopic signature of the coarsest

18 fraction is dominated by values similar to the fine fraction. The  Occ values of the coarse fraction record heavier values than the bulk and fine fraction by up to 2.4‰. The isotopic composition of waters from White Rock Creek and Lake was monitored for a period of twelve months from 1994 to 1995 (Robertson, 1995; Table 2.4). Those White Rock Lake water 18O values have been used to calculate 18O values of calcite that would have formed in stable oxygen isotopic equilibrium with lake water in conjunction with the corresponding average monthly high temperature at the time of collection (NOAA, 2012; Table 2.4), using the equation of O'Neil and others, (1969),

6 2 1000lncc-w = 2.78*10 /T – 2.89.

Measured 18O values of the Austin Chalk range from -3.7‰ to -4.4‰ (Collins, 2013). The

18O value of calcite formed on algal crusts sampled from the White Rock Lake margin is -2.0

‰, whereas watershed soil carbonate has a 18O value of -3.3‰, and authigenic calcite collected from White Rock Creek has a 18O value of -14‰ (Table 2.5; K. Ferguson, personal communication, 2008).

Table 2.4. Calculated values of carbonate precipitated in isotopic equilibrium with lake waters. 18 18 Water δ O values ( Ow) from Robertson (1995). Temperature data from NOAA (2012).

18 18 Avg High  OW  OCC Date (C) (SMOW) 1000ln (PDB) 29-Jun-94 32.9 -1.6 26.781 -5.6 18-Jul-94 35.9 -3.6 26.218 -8.1 8-Aug-94 35.6 -3.2 26.271 -7.6 26-Aug-94 35.6 -2.7 26.271 -7.2 13-Sep-94 31.3 -4.5 27.096 -8.1 21-Sep-94 31.3 -3.7 27.096 -7.3 4-Oct-94 25.7 -2.6 28.232 -5.1 10-Oct-94 25.7 -4.3 28.232 -6.8 28-Oct-94 25.7 -5.0 28.232 -7.4 11-Nov-94 19.4 -4.9 29.595 -6.0 18-Nov-94 19.4 -4.3 29.595 -5.5 11-Jan-95 13.6 -4.3 30.930 -4.2

Table 2.4. Isotopic values of carbonate associated with White Rock Lake deposits.

18 Sample  OCC (PDB) WRL cc fine -1.8 to -4.0 WRL cc coarse -0.9 to -3.1 WRL cc bulk -2.8 to -4.5 WRL Water to CC -4.2 to -8.1 Austin chalk -3.7 to -4.4 Algae -2.0 Soil -3.3 Authigenic -14

18 The  Occ values of the coarse WRL sediments are isotopically heavier than the calculated and measured authigenic calcite values, and the values of the underlying Austin

Chalk. As such, they cannot be attributed to authigenic precipitation from lake water, or to major contributions of detrital Austin Chalk to the sediment record. Possible mechanisms for the anomalously high WRL 18Occ values may be the precipitation of authigenic calcite at cooler temperatures from waters evaporatively enriched in 18O or from biogenic production of carbonate by lacustrine algae. Further investigation is necessary to elucidate the mechanisms responsible for the measured 18O values of the coarse fraction of the WRL calcite. The bulk and fine fractions of the WRL sediments have 18O values that closely resemble those of the underlying Austin Chalk, and most likely represent a detrital carbonate component transported to the White Rock Lake core sediment record from erosion of chalk in the upper watershed.

Weight percent calcite increases stratigraphically within the WRL core. The bulk fraction weight percent calcite increases from 13.7% at the base of the core to 36.2% to the top.

The overall weight percentage of bulk carbonate continually increases from 1948 onward, with a distinct inflection point in approximately 1971. The increased rate of carbonate deposition coincides with a marked decrease in farmland cultivation in the upper watershed, beginning in

18 1971. When taken into consideration with the  Occ values of the bulk fraction of Zone IV, which overlap with average 18O values of the Austin Chalk, the increase in carbonate abundance likely reflects an increased influx of detrital carbonate into the lake system during more recent times. This is consistent with greater runoff and resultant downcutting in the White

Rock Creek watershed due to continuing urbanization (Vicars-Groening and Williams, 2006) and

associated storm drainage that results in greater volumes of channelized water, erosive power, and down-cutting through Austin Chalk above White Rock Lake.

Primary organic matter production within a lake is controlled by the influx of limiting nutrients (N, P) into the system and by photic conditions. Soil erosion across the watershed is a dominant control on nutrient influx, and a primary source of biologically available phosphorus in most natural lake systems (Sonzongi et al, 1982). However, the influx of material into a lake is also affected by natural environmental or anthropogenic changes. For example, increased soil erosion can result from natural disasters such as fires and floods, as well as changes in land use practices associated with agriculture and urbanization. Landscape disturbance generally results in greater amounts of detritus, commonly nutrient-rich soil material, entering the water body and consequently, increased lake primary productivity (Meyers and Ishiwatari, 1993). As a watershed reaches build out, the natural landscape vegetation is replaced by paved surfaces, which contribute to increased intensity and erosive capacity of runoff, and a potentially greater influx of allochtonous material transported into the catchment.

Suburban development across White Rock Creek watershed replaced a vegetated landscape characterized by fertile, organic-rich soils with a landscape marked by a high percentage of impervious surfaces such as concrete and asphalt. The proliferation of impervious surfaces across the watershed has promoted intense runoff to storm drainage networks, which in turn rapidly transport runoff to White Rock Creek and its tributaries above White Rock Lake

(Figure 2.1C). This dramatic shift in percent and intensity of runoff has contributed to observed increases in storm flow volumes and peak flow levels (Vicars-Groening and Williams, 2006).

Hydrograph models characterizing changes in storm response for a hypothetical 10-year storm

event suggest that infiltration capacity of the watershed has decreased by approximately 60% between the 1960s and 2000s, and flooding events now occur at lower precipitation levels and higher intensities than in the 1960s (Vicars-Groening and Williams, 2006). The increased intensity of storm runoff to White Rock Creek and its tributaries has contributed up to several meters of fluvial-channel incision above White Rock Lake, which has subsequently increased exposure of the Austin Chalk in creek channels. The resultant creek margin instability provides additional sources of detrital carbonate and overlying, nutrient-rich soil material deposited in

White Rock Lake is likely a driving factor in the observed stratigraphic increase in weight percent calcite within the sediment core, and proliferation of primary productivity since the mid-

1960s (Figure 2.9).

Conclusions

Watershed urbanization has played a significant role in controlling the type and amount of organic material and the abundance of carbonate deposited into the White Rock Lake over the past century as deduced from various analyses of sediments within the White Rock Lake core.

13 18 The C/N ratios, δ C(SOM), δ Occ values (Figures 2.6, 2.9) of the White Rock Lake sediments correlate well with historical records of White Rock Lake watershed development (Figure 2.10) and reflect the impact of urbanization on the physical and geochemical characteristics of the landscape and sedimentary system. It is proposed that the concomitant shifts in the White Rock

Lake organic matter and carbonate geochemical record indicate four episodes, labeled Zones I-

IV, of lake sediment response to watershed development. Zone I deposits are reflective of lacustrine organic matter dominated by terrestrial C4 plants, and coincide with widespread 41

agricultural land use practices in place across the watershed at the time. Zone II deposits reflect a transitional period of land usage in the lower watershed and the increasing importance of C3 plants on the landscape during the late 1950s and early 1960s. Zone III deposits are representative of a later transitional period, marked by increasing lake primary productivity and detrital carbonate accumulation, as the northernmost reaches of the watershed experienced an

13 increased rate of urbanization from 1966 to 1982. In Zone IV,  C(SOM) values and C/N ratios suggest that the youngest lake sediments are dominated by algal C3 photosynthesizers and detrital carbonate deposition, resultant from the hydromorphological characteristics of a completely urbanized landscape.

13 18 The C/N ratios, δ C(SOM), δ Occ values of the White Rock Lake sediments correlate well with historical records of White Rock Lake watershed development, and reflect that urbanization has impacted the physical and geochemical characteristics of the watershed. Utilization of these geochemical proxies provides a potential means to assess the primary productivity conditions of lacustrine deposits for which no historical record of watershed conditions exists. The results shown here should be increasingly important for purposes of landscape utilization, urban landscape planning, and water quality characteristics of municipal water resources as urban centers continue to expand and increase in population across all continents except Antarctica.

Chapter 3

PENNSYLVANIAN-PERMIAN PALEOENVIRONMENTAL TRENDS FROM LITHOSTRATIGRAPHY, CLAY MINERALOGY, AND STABLE ISOTOPE GEOCHEMISTRY OF THE MADERA AND SANGRE DE CRISTO FORMATIONS, MORA COUNTY, NEW MEXICO

Introduction

This study presents lithostratigraphic, mineralogic, and stable isotope data from Upper

Pennsylvanian (Missourian-Virgilian) through Lower Permian (Wolfcampian) strata of the

Rowe-Mora basin, north-central New Mexico, USA. The sequence includes mixed marine carbonate and terrestrial clastic rocks in Pennsylvanian strata (Madera Formation) which transition to terrestrial-dominated siliciclastic rocks in the Lower Permian (Sangre de Cristo

Formation). Paleosol morphologies include eutric Argillisols in Missourian rocks, Calcisols in

Virgilian and Wolfcampian rocks, and gypsic subgroups in upper Wolfcampian strata. X-ray diffraction analysis of pedogenic clay minerals shows a shift from kaolinite, expansible 2:1 phyllosilicate, and illite-dominated assemblages representative of seasonal, subhumid environments in the lower part of the stratigraphy to illite-dominated assemblages, representative of increased aridity and decreased seasonality, in the uppermost part of the stratigraphy. The stratigraphic distribution of these morphologies is suggestive of stepped change from a subhumid seasonal climate in Missourian time to a nonseasonal arid climate in end-Wolfcampian time.

Fossilized vascular plant organic matter δ13C values range from ~-22.5 to -24 ‰ through the succession. However, paleosol calcite δ13C values range from ~-5 ‰ in Virgilian strata, ~-8

‰ in lower Wolfcampian strata, and ~-5 ‰ in topmost Wolfcampian strata. These data suggest relatively high atmospheric PCO2 (or low biological productivity) during deposition of

Wolfcampian strata. However, paleosol calcite δ18O values of the top-most Wolfcampian strata range from +1 to +3‰. This stratigraphically abrupt, +5 to +7 ‰, shift in calcite δ18O values is consistent with predicted changes in tropical rainfall δ18O values during deglaciation of the ice- sheet in Gondwanaland and aridification of western tropical Pangea.

Late Pennsylvanian to Early Permian time was an interval of significant global paleoclimatic change which included early phases of construction of the supercontinent Pangea

(Scotese and Golonka, 1992), numerous ice-building and ice-waning phases in high-latitude

Gondwanaland (Isbell et al., 2003), and complete reorganization of low-latitude tropical climates, floras, and faunas toward aridity and pronounced seasonality on the Pangean landmass

(DiMichele et al., 2001; DiMichele et al., 2006; Tabor and Montañez, 2002; Montañez et al.,

2007; Tabor et al., 2013a; Michel et al., 2015). Construction of the supercontinent is hypothesized to have reorganized global atmospheric circulation systems, which led to aridification over Pangaea (Parrish, 1993; Tabor and Poulsen, 2008). Lithological and paleobotanical evidence indicate a significant and rapid shift from humid Late Pennsylvanian to semi-arid Early Permian climate in the western tropics of Pangea (Tabor and Montañez, 2004;

DiMichele et al., 2004, 2006; Tabor et al., 2008, 2013a). Chemical-based paleoclimate proxies 44

suggest significant (~10°C) tropical warming (Tabor and Montañez, 2005; Tabor, 2007; Tabor et al., 2013a), and possibly large variations in paleoatmospheric PCO2 (Montañez et al., 2007) during this time, which appear to be correlated with a global climate change from icehouse to greenhouse states.

This work documents stratigraphic trends defined by the morphology, mineralogy, and geochemistry of Permo-Pennsylvanian paleosol profiles from north-central New Mexico in order to evaluate environmental and climatic conditions through this important transition in Earth’s history.

Geologic Background

During the Late Paleozoic, New Mexico occupied a position in western equatorial

Pangea, a region which migrated northward from ~0º to ~10ºN from Late Pennsylvanian through

Early Permian time (Figure 3.1; Scotese, et al., 1999; Blakey, 2007; Algeo et al., 2008). During this time, the region was subjected to several periods of tectonic instability associated with the uplift of the Ancestral Rocky Mountains, which resulted in the development of narrow, north- south oriented basins separated by basement-cored uplifts (Sutherland, 1963, Kottlowski and

Stewart, 1970; Kues and Giles, 2004; Figure 3.1). Tectonics, eustatic sea-level changes, and climate variations all impacted deposition of the late Paleozoic strata in this region.

Pennsylvanian-aged strata of this region are typically dominated by thick marine sequences in the southeast, which thin and become intercalated with continental sequences towards the northwestern part of the state. The study area is characterized by mixed marine carbonate and terrestrial clastic rocks in Upper Pennsylvanian strata of the Madera (Read and Wood,

1947)/Alamitos (Sutherland, 1963) Formation, which grade upward into terrestrial, fluvial- 45

dominated siliciclastic strata in the uppermost Carboniferous and Lower Permian (Figure 3.2).

The transition from marine to terrestrial deposition in the Taos Trough suggests emergence of the basin in response to either (1) rejuvenated vertical motions along the boundaries of uplifts that originated during the Pennsylvanian (Otte, 1959, Sutherland and Harlow, 1973), and/or (2) a response to eustastic sea-level fall associated with the waxing of continental glaciers during glacial phases (Isbell, et al., 2003; Jones et al., 2008).

The Rowe-Mora basin, located in the Southern Sangre de Cristo Mountains, was a Late

Paleozoic equatorial intraplate basin that developed in response to ancestral Rocky Mountain orogenesis, and was bordered by the Uncompahgre Uplift to the west, the Cimarron Arch to the north, and the Sierra Grande Arch to the east (Figure 3.1, b). The Pennsylvanian rocks of north- central New Mexico belong to the Madera and the Lower Sangre de Cristo formations, and a continuous record of sedimentation from Morrowan- through Wolfcampian-time is preserved in strata from Mora County, New Mexico. Early to mid-Pennsylvanian sedimentation in the Rowe-

Mora basin was almost exclusively marine (Soegaard and Caldwell, 1990, Soegaard, 1990).

Biostratigraphic control places deposition of the intercalated shallow marine and terrestrial deposits of the upper Madera Formation in Morrowan- through Virgilian-time and deposition of the terrestrial-dominated Sangre de Cristo Formation from Missourian to Wolfcampian-time

(Sutherland, 1963, Sutherland and Harlow, 1973, Soegaard and Caldwell, 1990).

Figure 3.1 Paleogeography of Late Paleozoic basins and uplifts in North America. (A) Paleogeographic reconstruction of the extent of the Late Paleozoic Midcontinent Sea (white region), the proximal land surfaces (gray regions), and paleolatitude are indicated are shown. Modified from Algeo et al. (2008). The yellow star marks the position of the Rowe-Mora basin. The green rectangle indicates the approximate boundary of (B) the detailed paleogeography of northern New Mexico during the Late Pennsylvanian. Modern latitude and longitude are indicated on the perimeter of the figure. The field location, near Mora, New Mexico is indicated by the red star. Modified from Casey (1980)

Figure 3.2 Stratigraphy and age assignments for the upper Paleozoic strata of the Rowe-Mora basin of northern New Mexico.

Methods Fieldwork and sampling

The Late Pennsylvanian to Early Permian Madera and Sangre de Cristo formations are exposed along cliff lines and road cuts near Highway NM518 in the southern Sangre de Cristo

Mountains near Mora, New Mexico (Figures 3.2, 3.3). Fieldwork included the identification, description, and collection of paleosols and associated lithologies throughout the Upper

Pennsylvanian-Lower Permian stratigraphy of the study area. Field descriptions were used to classify the paleosol profiles based on the dominant characteristics of each profile. These characteristics included horizonation, upper boundary type, color, reactivity with dilute HCl, and the presence of carbonate bearing horizons, translocated clay minerals, and vertic features (sensu

Mack et al., 1993). Each paleosol horizon was sampled for bulk matrix, whereas paleosol carbonate nodules and strata-bound fossil organic matter were sampled where present. All paleosol descriptions and samples were placed within a previously defined biostratigraphic framework (Kottlowski and Stewart, 1970; Sutherland and Harlow, 1973; Soegaard, 1990; Lucas and Krainer, 2005).

Figure 3.3 Field photos of exposures along Highway NM518 of the Madera and Sangre de Cristo formations in the Rowe-Mora basin. (A) Exposure of the Madera Formation along Highway NM518 near Mora, New Mexico. (B) Sequence of paleosol-bearing strata punctuated by arkosic and micaceous trough cross-bedded sandstones in the Sangre de Cristo Formation. (C) In situ arrangement of carbonate nodules from argillic Calcisol of the Madera Formation, described from a stratigraphic position of 583.3 m.(D) Stacked carbonate nodules in a platy to medium angular blocky mudstone are observed in a vertic Calcisol described from the Sangre de Cristo Formation at stratigraphic position of 1253.1 m. The paleosol profile is overlain by a coarse- grained arkosic sandstone.

Figure 3.4. Representative stratigraphic section of the Madera and Sangre de Cristo formations of the Rowe-Mora basin. Paleosols (n=41, stars), identifiable strata-bound organic matter (n=9, leaf symbol), carbonate samples (n=38, oblong symbol) were collected at the indicated stratigraphic positions.

Figure 3.5 Detailed stratigraphic section of the Madera and Sangre de Cristo formations of the Rowe-Mora basin. Paleosol horizons, limestones (shell symbol), identifiable strata-bound organic matter (leaf symbol), were collected at the indicated stratigraphic positions. Stratigraphic height is indicated in meters. Shown above: 0 to 425 m.

Figure 3.5 (continued) Detailed stratigraphic section of the Madera and Sangre de Cristo formations of the Rowe-Mora basin. Shown above: 425 to 850 m.

Figure 3.5 (continued) Detailed stratigraphic section of the Madera and Sangre de Cristo formations of the Rowe-Mora basin. Shown above: 850 to 1305 m.

Figure 3.5 (continued) Detailed stratigraphic section of the Madera and Sangre de Cristo formations of the Rowe-Mora basin. Shown above: 1305 to 1472 m.

Paleosol mineralogy

X-ray diffraction techniques were used to identify the mineralogy of bulk and clay-size

(<2 µm equivalent spherical diameter; e.s.d.) fractions of paleosol B-horizon matrix samples

(n=43). Samples were ground to powder in a ceramic ball mill and divided into two fractions.

One fraction was analyzed via X-ray diffraction of bulk powders. The second fraction was disaggregated via sonication in de-ionzed water, and the clay-size fraction (<2 μm e.s.d.) was isolated via centrifugation. Each clay-size fraction was divided into four aliquots and mounted as an oriented aggregate via the filter-membrane technique (Drever, 1973; Tabor et al., 2002) as follows: (1) K-saturated with 1.0 M KCl solution at room temperature (~23°C), (2) K-saturated with 1.0 M KCl solution and heating to 500°C for 2-4 hours, (3) Mg-saturated with 0.5 M with

MgCl2 solution, (4) Mg-saturation with 0.5 M MgCl2 solution and glycerol solvation with a 1:4

Glycerol: H2O solution. X-ray diffraction analyses of powder mounts included scans from 2 to

70° 2θ, with a scan speed of 3° 2θ/minute. X-ray diffraction analyses of oriented aggregates included scans from 2 to 30 2 and a scan speed of 2° 2θ /min. All X-ray diffraction analyses were conducted using a Rigaku Ultima III X-ray diffraction system using CuK radiation and scan window of 0.04° 2θ. Mineral identification followed the methods outlined in Moore and

Reynolds (1996) and Brown and Brindley (1984). The relative abundance of phyllosilicate minerals within each <2 µm sample was calculated using the mean area under the peak of its first order basal (001) reflection. The area under each XRD spectral peak was approximated by multiplying the peak height by the full peak width at half-height (FWHM). Each area was taken as a proxy for the integrated peak intensity (Moore and Reynolds, 1996), and corresponding

mineral abundance within the clay size fraction relative to (001) peaks of other phyllosilicate minerals present in the clay size fraction.

Paleosol carbonates

Paleosol carbonates (n=36) were identified using macroscopic inspection and reaction with dilute HCl, then prepared for petrographic and isotopic analysis (Figure 3.4). Staining with

Alizrin Red-S and potassium ferricyanide was utilized as a means of differentiating calcite from high- and low-Mg dolomite (Friedman, 1959; Dickson, 1965). The carbonate textures and fabrics associated with each paleosol were described and documented in order to discriminate between primary pedogenic calcite and secondary mineral phases that occluded pore space or replaced pre-existing textures during burial (Deutz et al., 2002). Carbonate mineral identifications were confirmed by x-ray diffraction of the stained thin sections. Calcite was identified based on XRD spectra peaks at ~3.86 Å and ~3.03 Å, whereas dolomite was identified based on the presence of peaks at ~2.89 Å and ~2.19 Å. Primary pedogenic calcites that were interpreted to have formed in well-drained conditions under open-system gaseous diffusion with the paleo-troposphere were analyzed for isotopic composition (Ekart et al., 1999; Sheldon and

Tabor, 2009). Micritic (<2 µm crystal size) calcite was sampled from matching billets from petrographic thin sections using a handheld Dremel rotary tool with a diamond carbide-tipped bit in order to collect 10 to 100 mg of sample for isotopic analysis. Samples were reacted overnight with 100% anhydrous phosphoric acid at 25°C (McCrea, 1950). Reaction products were cryogenically purified using high-vacuum extraction lines, and CO2 yields were determined via

13 18 mercury manometry. The  C and  O values of extracted CO2 samples were determined at

Southern Methodist University using a Finnigan-MAT 252 isotope ratio mass spectrometer.

Stable isotope values are reported in standard delta notation,

18 13  O (or C) = (Rsample/Rstandard -1)*1000, where R is the ratio of heavy-to-light stable isotope present in the sample or standard, and - values are reported relative to the Peedee Belemnite standard (PDB) for carbon and Standard

Mean Ocean Water (SMOW) for oxygen isotope values.

Fossil organic matter δ13C

Identifiable fossil vascular plant material of 9 samples collected from the Madera

Formation was chemically pretreated for isotopic analysis (Figure 3.4). Samples were treated with successive aliquots of concentrated HCl in order to eliminate inorganic carbon, repeatedly rinsed with deionized water until reaching a pH of ~5.5, and then dried at 50 ºC. Pyrolysis of organic matter samples followed the method of Boutton and others (1991). Individual samples were sealed in a vycor tube with approximately 1 g copper oxide and 0.5 g of copper metal. The sample tubes were incrementally heated, initially to 900ºC for 2 hours, then stepped down

2º/min., until reaching 650 ºC. Samples were allowed to dwell at 650 ºC for two hours, and then cooled to room temperature. The evolved CO2 was cryogenically purified using high-vacuum extraction lines, and the δ13C values determined at Southern Methodist University using a

Finnigan-MAT 252 isotope ratio mass spectrometer.

Carbonate-associated organic material was isolated from 16 paleosol carbonate samples using a SPEX SamplePrep Mixer/Mill 8000D. Approximately 6 to 10 grams of pedogenic carbonate and an alumina ceramic grinding ball were loaded into alumina ceramic vial. The mill

was operated for 6 minutes per sample, or until the carbonate nodule was completely powdered.

The powdered sample was treated with successive aliquots of dilute (1.0N) and concentrated

(12.1N) HCl to remove any carbonate, then rinsed repeatedly with deionized water until a pH of

~5.5 was reached. Samples were dried at 50 ºC and processed for isotopic analysis via combustion as described above.

Results

Pedotypes

Permo-Carboniferous paleosols in the Madera and Sangre de Cristo formations of the

Rowe-Mora basin are classified into 7 different morphologies (Mack et al., 1993), based upon morphological attributes among 43 paleosol profiles (Figures 3.6, 3.7). Paleosol profiles were divided into diagnostic horizons based on color, ped size and structure, mottling, and accumulations of translocated clays, carbonate nodules, and rhizoliths (Retallack, 1990; Tabor and Montañez, 2002; Tabor et al, 2006; Rosenau et al., 2013). Described below is the range of characteristics of these paleosol types through the stratigraphy. Each description is subsequently followed by a paleoenvironmental interpretation that is based on the modern environments in which analogous morphological features develop.

Argillisols (n=3) are characterized by the presence of argillans, or translocated clays which have accumulated upon and coat ped surfaces. Clay-coated peds range from about 1 to 5 cm in long dimension and define both subangular blocky and angular blocky ped structure.

These paleosols are present at 683, 693, and 1338 m within the stratigraphic succession, in the middle and upper portions of the stratigraphy (Figures 3.6, 3.7). Argillisols are interpreted to represent humid, well-drained conditions of pedogenesis. This assumes that vertically flowing 59

meteoric waters drained through the soils at the time of pedogenesis in a sufficient amount to leach Ca2+ from clay mineral exchange sites, which is necessary to form soil profiles with argillic horizons (Markewich and Pavich, 1991).

Calcic Argillisol (n=4)/Argillic Calcisols (n=2) are characterized by the presence of argillans, translocated clays which coat ped surfaces, and calcic horizons which include nodular and tubular carbonate accumulation (Figures 3.6). Clay-coated peds range from about 1 to 7 cm in diameter and define mostly angular but also occasional subangular-blocky structure.

Carbonate nodules range from <1cm to ~10 cm in diameter and generally become larger with depth in a paleosol profile. Tubular carbonate features are typically narrow, ranging from a few mm to ~2 cm wide and can be traced vertically as much as 35 cm through parts of the profiles.

Calcic Argillisols occur in the middle part of the measured stratigraphic section, at 750, 1189,

1190, 1196 m and argillic Calcisols occur at 583 and 587 m in the stratigraphic succession

(Figure 3.7). Calcic Argillisols and argillic Calcisols are interpreted to represent sub-humid, well-drained conditions of pedogenesis, during which evaporation periodically exceeded precipitation. This interpretation assumes that soils must have been well drained and received enough meteoric rainfall to thoroughly leach the upper parts of the paleosol profiles in order to remove Ca2+ ion and facilitate translocation of clay-sized materials (Markewich and Pavich,

1991), but also received insufficient rainfall to remove all Ca2+ ion from the profile and thus permitted the accumulation of CaCO3 in lower subsurface horizons.

Vertic Calcisols (n= 1) and calcic Vertisols (n=2) are characterized by features associated with shrink-swell processes within the paleosol profile (Figure 3.6). These shrink-swell features include wedge-shaped peds, slickensides, and clastic dikes. Wedge-shaped peds range from a few cm to ~30 cm across and become larger downward in profiles generally. Slickensides consist 60

of highly oriented clay and fine silt materials which are concentrated upon and against wedge- shape aggregate surfaces. Slickensides occur in horizons ranging from the upper surface of paleosols to ~1.5 meters beneath the surface horizon. Clastic dikes may be sandstone- or mudstone-filled: they typically taper downward with tops usually around 3 cm wide and also extend downward to as much as 75 cm. Calcic Vertisols and vertic Calcisols occur at 1260,

1334, and 1428 m; near the top of the stratigraphic succession (Figure 3.7). These so-called

“vertic” features derive from a presence of 2:1 expansible clay minerals which are preserved through seasonal precipitation patterns that are incapable of leaching base cations in the soil profile (Wilson, 1999). Therefore, the vertic features in this stratigraphic succession are interpreted to form in environments with distinct, seasonally wet and dry periods. The calcic features are interpreted to form under well-drained conditions, when evaporation exceeded precipitation.

Calcisols (n=23) are characterized by red, blocky to massive, mudstones with prominent calcic horizons, which contain varying amounts of carbonate that occurs as nodules, tubules, or indurated horizons (Figure 3.6; Mack, et al., 1993). Most Calcisols are blocky mudstones with pedogenic structure ranging from fine to coarse medium angular blocky. However, a few

Calcisols in the upper ~1/3 of the Sangre de Cristo Formation have horizons characterized by massive structure. Nodules range from <1 cm to ~20 cm in diameter; larger nodules tend to be lenticular and occasionally are cemented to adjacent nodules and their abundance typically increases toward the base of profiles. Tubules range from ~2 to 4 cm in diameter and are oriented with their long dimensions vertical to subvertical in the paleosol profiles. These soils form in well-drained environments under arid conditions, in which evaporation exceeds precipitation.

Calcisols are prevalent throughout the stratigraphy, increasing in occurrence near the middle to 61

top of the measured stratigraphic section. Calcisols occur at 518, 716, 717, 723, 914, 1009,

1061, 1133, 1203, 1236, 1248, 1288, 1323, 1330, 1445, 1447, 1451, 1452, 1456, 1460, 1462,

1467, 1470, 1479 m in the stratigraphy (Figure 3.7). Calcisols represent pedogenesis during stable periods of relatively dry conditions. Based upon the occurrence of modern soils with carbonate accumulations, these paleosols formed in climates that received precipitation <750 mm/yr (Royer, 1999).

Calcic Gypsisols (n=2) are characterized by horizons that include carbonate and gypsum nodules and lenses up to 1 meter in length and 12 cm thick. The calcic Gypsisols in the uppermost Sangre de Cristo Fm have horizons with massive to weak angular blocky structures.

Dissolution structures are present on the interior of nodules and lenses, suggesting carbonate replacement of primary gypsic satin-spar texture (Figure 3.6, 3.9). These paleosols occur near the top of the stratigraphy, near 1458 m (Figure 3.7). Gypsic nodules include satin-spar and swallow-tail tabular crystals and range from ~2 to 10 cm in diameter. Considering that gypsic soil horizons in modern soil-forming environments occur in climates characterized by low rainfall, <300 mm/yr, (Watson, 1992 ), the calcic Gypsisols represent soil formation during the most arid climates recorded within the Madera and Sangre de Cristo formations.

Figure 3.6 Diagrams of the seven paleosol type morphologies present in the Madera and Sangre de Cristo formations of the Rowe-Mora basin. Drawn here are: Argillisol from stratigraphic position of 683.17 m, calcic Argillisol from stratigraphic position of 1188.54 m, argillic Calcisol from stratigraphic position of 583.34 m, argillic Vertisol from stratigraphic position of 510.79 m, Vertisol from stratigraphic position of 1470.30 m, vertic Calcisol from stratigraphic position of 1253.05 m, Calcisol from stratigraphic position of 1202.53 m, calcic Gypsisol from stratigraphic position of 1458.55 m. Key explains symbols and nomenclature used to represent sedimentary and pedogenic structures, lithology, and paleosol horizons. See text for discussion.

Figure 3.7 Stratigraphic position versus (a) pedotypes, as described from field-based indicators, delineated in the Mack and others (1993) classification scheme; (b) cumulative clay mineral percent of major constituents in the <2 µm phyllosilicate fraction of paleosol matrix samples; (c) 13 Δ C(CC-OM) values calculated from paleosol carbonates and organic matter occluded within pedogenic carbonate nodules and tubules from the Madera and Sangre de Cristo formations in the Rowe-Mora basin. See text for discussion.

Paleosol mineralogy

Powder diffraction spectra of bulk samples exhibit intense X-ray diffraction peaks at 6.4,

4.26, 4.04, 3.34, 3.1 – 3.2, 3.7, 3.85, and 3.00 Å. The Peaks at 4.26 and 3.34 Å correspond to d(100) and d(101) Miller indices, respectively, of quartz (SiO2). Peaks at ~6.4Å, 4.04Å, and 3.1

– 3.2Å likely correspond to d(001), d(201), d(202) and d(040) Miller indices, respectively, of calcian albite ((Na, Ca)(Si, Al)4O8). Peaks at ~3.85Å and ~3.03Å correspond to d(102) and d(104) Miller indices, respectively, of calcite. All samples also exhibit relatively low and broad peaks near 14, 10 and 7 Å, which correspond to d(hkl) spacings of 2:1 and 1:1 phyllosilicates.

X-ray diffraction analyses of <2µm fraction oriented aggregates of paleosol B-horizons indicate the presence of several minerals, which were identified based on the position and intensity of low-order basal reflections (Brown and Brindley, 1984). The clay mineral fraction contains kaolinite, illite, and hydroxy-interlayered minerals (HIM), as well as feldspar and quartz (Figure

3.8). Koalinite was identified based on the presence of basal reflections at approximately 7.1Å, and 3.5Å, which correspond to the d(001) and d(002) Miller indices, respectively. Illite was identified based on the presence of d(001) and d(002) basal reflection peaks at ~10.0Å and

~5.0Å, respectively (Moore and Reynolds, 1996). The breadth and symmetry of illite d(001) peaks varies among, and within, samples according to chemical pre-treatment. Specifically, some samples exhibit increased symmetry and peakedness of the illite d(001) peak after Mg-saturation and glycerol solvation compared to K-saturation. Such behavior is indicative of the presence of poorly sorted 2:1 expansible phyllosilicates in addition to illite (Moore and Reynolds, 1996).

HIM peaks exhibit asymmetrical patterns between 10.5-11.5Å upon KCl, 25°C treatment, which collapse to 10Å upon heating to 500°C (Figure 3.8). Broad, low d(001) peaks occur near 14Å upon treatment with MgCl, whereas these peaks shift slightly toward 14.5 Å upon glycerol 65

solvation (Figure 3.8). This behavior is characteristic of poorly ordered 2:1 expansible phyllosilicates (HIM: hydroxy-interlayered mineral; Moore and Reynolds, 1996). Based upon

FWHM measurements of kaolinite, HIM, and illite d(001) peaks, these minerals are present in varying abundances in the clay mineral fractions of paleosol profiles throughout the stratigraphy

(Figure 3.7). From ~500 to 700 meters in the stratigraphic succession, paleosols contain nearly equal amounts of kaolinite and illite. Illite dominates the mineralogy of the paleosol profile clay- size fractions from ~725 to 1100 m, and again from ~1300 to 1450 m in the stratigraphic succession. The abundance of kaolinite and HIM decreases dramatically near the top of the section.

Petrographic Analysis of Paleosol Carbonates

Four main calcite textures are recognized from Bk horizons of paleosol bearing strata of the Madera and Sangre de Cristo formations of the Rowe-Mora basin (Figure 3.9). Clotted micritic calcite is common in paleosol carbonate nodules sampled from the Bk horizons of argillic Calcisols, vertic Calcisols, and Calcisols. The clotted micrite texture is characterized by amalgamations of red or gray microcrystalline calcite, often in association with clay-lined fractures, opaque iron oxide minerals with diffuse boundaries, and blebs of non-oriented clay minerals. Clotted micrite occurs is observed in argillic Calcisols at 518.0m, 587.4, Calcisols at

716.3 mm, 1236.4 m, 1247.5 m, 1253.6m, and 1470.2m (Figure 3.9, A). Clotted micrite in association with calcite-spar filled septarian fracture features occurs in argillic Calcisols at

583.3m, 586.6m, Calcisols at 519.1 m, 1061.4 m, 1172.8 m, and a vertic Calcisol at 1259.9m

(Figure 3.9, B). Massive red micritic calcite comprises the bulk of carbonate nodules in Calcisols at 1323.5 m and 1330.0 m (Figure 3.9, C). Coarse and fine clay- and micritic calcite-filled 66

tubules (Figure 3.9, D) ranging in size from 0.25 to 3.0 mm in width and up to 5.0 mm in length occur in argillic Calcisols are encountered at a stratigraphic height of 519.1m and 587.7 m within the stratigraphic section.

Several anomalous carbonate textures are observed from paleosol carbonate horizons and a singular nodular limestone (Figure 3.10). Nodules comprised of granules of microcrystalline calcite matrix, with fractures lined by dolomite and further infilled by sparry calcite are observed in Bk horizons of Calcisols at 1451.6 m, 1451.9 m, 1462.1m, 1470.15m, and 1470.2 m.

Swallowtail pseudomorphs of calcite after gypsum were observed in a nodular limestone encountered at a stratigraphic height of 1102.2 m. Nodules of intermingled gypsum, dolomite, and micritic calcite, cut by micritic-calcite filled fractures were observed in a Calcisol at

1435.83m. Clotted micrite with common irregularly-shaped voids bound by discrete calcite prisms occurs in lenticular carbonate nodules at 1202.2 m in the stratigraphic section. This texture also occurs in the Bk horizon of a Calcisol at 1247.5 m, and the Bk horizon of a vertic

Calcisol at 1253.8m.

Stable Isotope Geochemistry of Paleosol Carbonate and Organic Matter

Petrographic analysis of paleosol carbonate nodules and tubules reveals that most samples are composed of clotted red micrite to microspar calcite with coarse spar-crystal filled septarian cracks. Samples of micritic calcite were drilled and subsequently analyzed for stable carbon and oxygen isotope values of calcite as well as carbon isotope values of organic matter occluded within the nodules and tubules.

Figure 3.8 Background-subtracted X-ray diffractograms of oriented aggregates from the clay- size (<2 µm) fraction of a paleosol matrix sample collected from the Madera Formation at a stratigraphic position 683.17 m. The clay-size fraction of the sample was K-saturated and analyzed at room temperature (light blue spectrum), K-saturated and heated at 500°C for 2 h prior to analysis (dark blue spectrum). The collapsible 7.2 Å peak in the K-saturated spectrum is a diagnostic indicator of the presence of kaolinite. Hydroxy interlayered minerals (HIM) identified by increased peakedness and symmetry of the 10.1 Å peak after Mg saturation plus glycerol solvation. Kaolinite-smectite interlayered mineral 34 Å superstructure, smectite identified from ~8.3 Å peak, kaolinite identified from 7 Å peak that collapses upon heating to 500 °C.

Figure 3.9 Examples of common micritic calcite textures observed in paleosol nodules from the Madera and Sangre de Cristo formations of the Rowe-Mora basin imaged under crossed Nichols. (A) Clotted gray and red micrite sampled from a Madera Formation Calcisol at 519.4 m. (B) Clotted red micrite with fine Fe-oxides and septarian fractures, sampled from a Madera Formation argillic Calcisol at 583.3 m. (C) Massive micritic calcite, sampled from Sangre de Cristo Formation Calcisol at 1330.0 m. (D) Cross section of a micritic calcite and clay-filled tubule from a Madera Formation argillic Calcisol at 587.7 m. All meter-level references refer to stratigraphic positions in Figure 3.2. See text for discussion.

Figure 3.10 Examples of alteration of primary mineral textures by calcite and dolomite in paleosol carbonate nodules and a nodular limestone unit from the Sangre de Cristo Formation in the Rowe-Mora basin. All images captured under crossed Nichols. (A) Calcite matrix, after primary gypsum, with a dolomitic fracture infilled by sparry calcite observed in paleosol carbonate nodule from a Calcisol at 1462.1 m. (B) Swallowtail pseudomorph of calcite after primary gypsum, sampled from a nodular limestone bed at 1102.2 m. (C) Intermingled gypsum, dolomite, and micritic calcite, cut by micritic calcite fractures, sampled from a Calcisol at 1435.83m. (D) Clotted micrite with common irregularly shaped voids bound by calcitic prisms, similar in texture to structures described as Microcodium, sampled from a rhizolith in a vertic Calcisol at 1253.8m. All meter-level references refer to stratigraphic positions in Figure 3.2. See text for discussion.

Pedogenic calcite (n=37) samples exhibit a range of δ13C values from -8.0 to -5.1‰

(PDB), and a range of 18O values from -8.5 to 2.9‰ (PDB) (Table 3.1; Figure 3.11). There appears to be no strong correlation between weight percent calcite values and corresponding

18O or δ13C values (Table 3.1). While there appears to be no significant stratigraphic trend in paleosol calcite δ13C values, paleosol calcite 18O values show an abrupt +7‰ shift in the upper

Sangre de Cristo Formation samples (Table 3.1, Figures 3.12, 3.13). No significant correlation is

13 18 observed between carbonate carbon isotope (δ C(CC)) and oxygen isotope (δ O(CC)) values versus weight-percent calcite estimates based upon CO2 yields derived from H3PO4 reaction of carbonates from the Madera and Sangre de Cristo formations (Figure 3.14, 3.15)

Identifiable fossil vascular-plant organic matter δ13C values range from -24.1 to -22.4‰

(n=9). Carbonate-associated organic matter δ13C values range from -25.1 to -21.1‰ (n=18;

Table 3.2, Figure 3.13). With the exception of larger calcite δ13C values toward the top of the stratigraphic section (Table 3.1, figure 3.13), there appears to be no significant stratigraphic trend in paleosol calcite δ13C values. However, paleosol calcite 18O values show an abrupt +7‰ shift in the upper Sangre de Cristo Formation samples (Table 3.1, figure 3.12). The stable carbon isotope difference between paleosol calcite and co-existing carbonate-associated organic matter

13 13 ( C(CC-OM)) ranges from 13.1 to 19.5‰ (Figure 3.7).  C(CC-OM) values range from ~15 to 17‰ in the upper Madera Formation, from stratigraphic positions of 519.4 m to 750.3 m. The single

13  C(CC-OM) value from a stratigraphic position near the base of the Sangre de Cristo Formation is

13 13.1‰.  C(CC-OM) values in the upper 400 m of the Sangre de Cristo Formation range from ~16 to 19.5‰ and exhibit a trend toward larger values upward through the top of the formation.

-2

-4

O O (PDB)

18 

-6

-8

-10 -8 -7 -6 -5 13C (PDB)

Figure 3.11 Measured δ18O values versus δ13C values of paleosol carbonates from the Madera and Sangre de Cristo formations of the Rowe-Mora basin.

1500

1400

1300

1200

1100

1000

900 Stratigraphic Stratigraphic Height (m)

800

700

600

500 -10.0 -8.0 -6.0 -4.0 -2.0 0.0 2.0 4.0 δ18O(cc) (PDB)

Figure 3.12 Cross plot of stratigraphic position versus δ18O values of carbonates from the measured section of the Madera and Sangre de Cristo formations of the Rowe-Mora basin.

1600

1400

1200

1000

800

Stratigraphic Stratigraphic Height (meters) 600

400 Identifiable OM

Carbonates 200 Carbonate- associated OM

0 -26 -24 -22 -20 -18 -16 -14 -12 -10 -8 -6 -4 13C (PDB)

Figure 3.13 Cross plot of the stratigraphic position versus δ13C values of identifiable, strata- bound organic matter (triangles), paleosol carbonate-associated organic matter (circles) and carbonate samples from the measured section of the Madera and Sangre de Cristo formations in the Rowe-Mora basin. Solid black line denotes the boundary between the Madera and Sangre de Cristo formations at a stratigraphic position of 805 m.

-4.0

-5.0

-6.0

-7.0

-8.0

Ccc Ccc (PDB) 13

-9.0

-10.0 0 20 40 60 80 100 Weight % Calcite

13 Figure 3.14 Plot of carbonate carbon isotope (δ C(CC)) values versus weight-percent calcite estimates based upon CO2 yields dereived from H3PO4 reaction of carbonates from the Madera and Sangre de Cristo formations of the Rowe-Mora basin.

4.0

2.0

0.0

-2.0

-4.0

O O (PDB) 18  -6.0

-8.0

-10.0 0 20 40 60 80 100 Weight % Calcite

18 Figure 3.15 Plot of carbonate oxygen isotope (δ O(cc)) values versus weight-percent calcite estimates based upon CO2 yields dereived from H3PO4 reaction of carbonates from the Madera and Sangre de Cristo formations of the Rowe-Mora basin.

Table 3.1 Carbonate sample stratigraphic positions stable carbon and oxygen isotope values and weight percent calcite of analyzed samples from the Madera and Sangre de Cristo formations.

Sample Height δ13C δ18O δ18O wt% ID (m) Description (PDB) (PDB) (SMOW) calcite Mora35 510.79 Ck - Horizon -6.0 -4.0 26.8 79.2 Mora40 519.44 Bk with Rhizoliths -6.1 -4.8 26.0 --- Mora42 583.34 Bk -6.9 -5.1 25.7 59.6 Mora43 583.54 Bk -7.3 -5.4 25.3 54.8 Mora44 583.99 Bk -6.1 -4.7 26.1 59.8 Mora45 586.54 Bk -7.0 -5.1 25.6 45.2 Mora46 587.44 Bk -6.3 -5.9 24.9 68.6 Mora47 587.69 Btk -5.1 -8.5 22.1 --- Mora65 717.07 Stage II mud rich calcisol -6.1 -4.5 26.3 89.5 Mora66 722.67 Groundwater carbonate -6.9 -5.2 25.5 79.3 Mora67 727.32 BkIII -7.1 -4.9 25.9 76.4 Mora70 750.28 Btk horizon -7.1 -5.4 25.4 30.7 Mora71 750.65 Btk horizon -7.0 -4.5 26.3 58.9 Mora74 913.69 Bk horizon -8.0 -3.8 27.0 51.3 Mora76 1009.30 Bk -7.7 -3.9 26.9 74.3 Mora77 1061.40 Bk horizon -6.8 -4.8 26.0 56.9 Mora83 1132.60 Bk horizon -6.6 -5.1 25.7 82.5 Mora86 1188.79 Bk nodules -6.7 -6.4 24.4 53.3 Mora88 1189.10 Btk -6.3 -5.8 24.9 55.1 Mora90 1189.90 Btk mm-size CaCO3 nodules -7.1 -4.1 26.7 9.4 Mora 90 1189.90 Red laminated mud with micrite veins -7.2 -6.6 24.1 46.4 Mora93 1197.20 Bkss -5.8 -5.5 25.2 91.0 Mora97 1202.80 Bk -7.8 -4.8 26.0 86.7 Mora99 1226.80 Bk nodules -6.4 -4.8 26.0 77.3 Mora103 1247.50 BkIII -6.5 -5 25.8 85.4 Mora106 1253.10 Bk in ang blky fine sandstone -6.3 -5.2 25.5 72.9 Mora107 1253.60 Bk -6.0 -4.8 26.0 73.6 Mora108 1253.80 Bk nodules and rhizoliths -6.4 -4.8 26.0 69.0 Mora110 1288.20 Bk -6.6 -5.4 25.3 79.2 Mora111 1323.50 Bk -7.5 -5.1 25.7 69.1 Mora112 1330.10 CaCO3 nodules, recrystallized -6.1 -4.5 26.3 66.7 Mora115 1354.85 lenticular microbial carbonate -5.9 2.8 33.8 --- Mora118 1447.80 Bk -5.9 2.5 33.5 82.2 Mora120 1451.70 Bk -5.9 0.8 31.7 21.8 Mora124 1455.70 Bk nodules -5.6 2.9 33.9 37.8 Mora125 1457.50 CaCO3 after gypsum -5.9 1.8 32.8 --- Mora129 1462.10 BkII nodules and rhizoliths -5.9 1.8 32.7 37.1 Mora132 1470.20 BkII -5.6 2.6 33.6 58.5

Table 3.2 Stratigraphic position of samples and the corresponding stable carbon isotope values of organic matter sampled from the Madera and Sangre de Cristo formations of the Rowe-Mora basin. Identifiable organic matter was sampled from hand-specimens of samples collected from 12.30 m to 414.57 m in the stratigraphic section. Organic matter was isolated from paleosol carbonate nodules via acid digestion for samples collected from 519.44 m to 1455.70 m.

Height Sample ID (m) Description δ13C (org) δ13C(cc) Δ13C Mora3 12.3 Organics -23.863 Mora10 71.05 Organics -24.123 Mora11 77.75 Organics -22.608 Mora12 92.05 Organics -22.419 Mora13 120.55 Organics -22.85 Mora15 129.55 Organics -23.917 Mora16 137.30 Organics -22.933 Mora24 182.74 Organics -22.775 Mora29 414.57 Organics -22.895 Mora40 519.44 cc associated om -23.373 -6.1 17.3 Mora42 (2) 583.34 cc associated om -22.14 -6.9 15.2 Mora46 587.44 cc associated om -22.658 -6.3 16.4 Mora46 (2) 587.44 cc associated om -22.184 -6.3 15.9 Mora65 717.07 cc associated om -21.328 -6.1 15.2 Mora70 750.28 cc associated om -23.544 -7.1 16.5 Mora71 750.65 cc associated om -22.466 -7.0 15.5 Mora74 913.69 cc associated om -21.106 -8.0 13.1 Mora77 1061.39 cc associated om -23.84 -6.8 17.0 Mora86 1188.79 cc associated om -23.059 -6.7 16.3 Mora90 1189.90 cc associated om -23.573 -7.2 16.4 Mora99 1226.80 cc associated om -22.544 -6.4 16.1 Mora103 1247.50 cc associated om -23.498 -6.5 17.0 Mora108 1253.80 cc associated om -25.238 -6.4 18.8 Mora110 1288.20 cc associated om -23.751 -6.6 17.2 Mora121 1451.90 cc associated om -24.948 -5.9 19.0 Mora124 1455.70 cc associated om -25.13 -5.6 19.5

Discussion

Pedotypes

Paleosols in the Madera and Sangre de Cristo formations of the Rowe-Mora basin exhibit variations in characteristics such as texture, color, structure, and occurrence and concentration of soil-forming minerals. Based upon field observations of these characteristics, there are seven different paleosol types, or pedotypes, which reflect differences in four important soil forming processes distributed upon paleolandscapes and through time in the Rowe-Mora basin. These pedotypes are Argillisols, calcic Argillisols, argillic Calcisols, calcic Vertisols, vertic Calcisols,

Calcisols, and calcic Gypsisols (Figures 3.6, 3.7; Mack et al., 1993). Protosols, or very weakly developed paleosol profiles, occur within the stratigraphic succession of the Madera and Sangre de Cristo formations also, but are not useful for paleoenvironmental/paleoclimate reconstructions

(Mack et al., 1993; Buck and Mack, 1995) and are not considered further herein.

The dominant pedogenic process in the formation of Argillisols is the accumulation of clay-sized material in sub-surface layers through lessivage, or downward washing of fine-grained material. Calcic Argillisols in the succession have carbonate nodules and tubules in horizons beneath argillic horizons. Modern soil profiles accumulate subsurface argillic horizons upon stable (normally requiring at least a few thousand years) and well-drained land surfaces (Soil

Survey Staff, 1975). Furthermore, climate also appears to be a factor in the development of these horizons because they do not appear to form in regions with perhumid climates, and therefore require climates where the soils become thoroughly or partially dry during some seasons (Soil

Survey Staff, 1975). Yet, these soils must also receive sufficient rainfall to thoroughly leach calcium ions from the exchange complexes of clay minerals the upper horizons of the soil, as it is only then that it appears clays will disperse and translocate downward through the profile 78

(Markewich and Pavich, 1991). Therefore, the presence of paleosols with argillic horizons are taken to represent relatively long-term stability of paleo-land surfaces in environments characterized by relatively high rainfall, but with a distinct dry season.

The dominant pedogenic process in Vertisols is episodic shrink-swell (also called pedoturbation) that results from wet-dry cycles, whereas calcic Vertisols also include an important process of carbonate accumulation (Mack et al., 1993). Shrink-swell processes in

Vertisols are commonly a result of the presence of expansible 2:1 clay minerals within the profile that swell during wet periods and contract during dry periods (Buol et al., 1997; Schaetzl and Anderson, 2005; Southard et al., 2012). The expansion and contraction of the profiles leads to characteristic morphological features such as slickensides, pressure faces, clastic dikes, and gilgai surface and mukkara subsurface expression. Additionally, shrink-swell processes lead to the formation of wedge-shaped peds. These soil types commonly develop in climates characterized by strongly seasonal rainfall such as monsoonal- or Mediterranean-type rainfall patterns (Southard et al., 2012). Not all Vertisols form from seasonal rainfall patterns; some result from episodic fluctuation of the water table (Mintz et al., 2011). Therefore, Vertisols in the

Sangre de Cristo Formation represent episodic wetting and drying of parts of the paleolandscape, perhaps resulting from extreme seasonal rainfall patterns over the region. Calcic Vertisols also represent episodic wetting and drying, seasonal rainfall, as well as incomplete leaching which possibly corresponds to drier environmental conditions or greater time of pedogenesis in order to accumulate substantial amounts of soil carbonate (Marriott and Wright, 1993).

The dominant pedogenic process in the formation of Calcisols is the accumulation of carbonate in subsurface soil horizons (Mack et al., 1993). Modern soil profiles that accumulate calcium carbonate in subsurface horizons are most commonly found in climates where 79

evapotranspiration exceeds precipitation for most months of the year (Buol et al., 1997).

Specifically, in well-drained soils that receive >760 mm/yr, carbonate will not usually be retained within the subsurface horizons of the soil (Royer, 1999). Therefore, Calcisols in the

Madera and Sangre de Cristo formations are interpreted to represent relatively dry climatic conditions compared to those associated with clay mineral accumulation in Argillisols in other parts of the stratigraphy. Furthermore, Calcisols represent relatively dry conditions without extreme seasonality in precipitation that likely was an important soil forming factor at other times, as indicated in the stratigraphy that preserves Vertisols and calcic Vertisols.

The dominant process involved in the creation of calcic Gypsisols is the accumulation of shallow subsurface gypsum nodules and veins interfingering with and below horizons of carbonate nodules. Gypsum is commonly present in subsurface horizons of modern soils that are characterized by evapotransporation far in excess of precipitation (Eswaran and Zi-Tong, 1991;

Watson, 1992). Specifically, gypsum is leached from the soil in regions that receive precipitation in excess of 250 mm/year (Watson, 1992). Therefore, calcic Gypsisols in the upper

Sangre de Cristo Formation represent the most extreme examples of dry conditions through the stratigraphic succession, characterized by high evapotranspiration rates and relatively low precipitation that resulted in incompletely leached soil profiles which accumulated both gypsum and calcite. Collectively, paleosol morphologies in upper Madera and Sangre de Cristo formations in the Rowe-Mora Basin define a long-term trend from relatively humid climates in latest Carboniferous time to increasingly seasonal and drier climates through the Permian Sangre de Cristo Formation recorded by the presence of Vertisols, calcic Vertisols, Calcisols, and calcic

Gypsisols (Figure 3.7). However, calcic Argillisol profiles in the mid-Sangre de Cristo

Formation record evidence for an episode of increased humidity/ moisture on the long-term trend toward peak aridity near the top of the Sangre de Cristo Formation.

Petrographic analysis of carbonates

Petrographic analysis of paleosol carbonate nodules indicates that a wide variety of conditions persisted during the formation and subsequent burial of strata preserved in the Madera and Sangre de Cristo formations in the Rowe-Mora basin. The four main carbonate textures are described from Bk horizons of paleosol bearing strata of the Madera and Sangre de Cristo formations of the Rowe-Mora basin and include clotted micrite, clotted micrite with septarian features, massive micrite, and micritic tubules (Figure 3.9). These Bk horizon features are taken to represent primary pedogenic carbonate formation in argillic Calcisol, vertic Calcisol, and

Calcisol paleosol profiles.

Several anomalous carbonate textures occur in paleosol carbonate horizons and a singular nodular limestone that collectively are interpreted as examples of alteration of primary mineral textures and replacement by calcite and dolomite (Figure 3.10). Clotted micrite with common irregularly shaped voids bound by individual calcitic prisms, similar in appearance to the texture often referred to as Microcodium was observed in a rhizolith from a vertic Calcisol at 1253.8m.

According to Košir (2004) the occurrence of Microcodium suggests pedogenic deposition of microbial via biogenic means (Klappa 1978; Kabanov et al., 2008) (Figure 3.10, D). As depicted in figure 3.10, a granular texture of calcite matrix is interpreted to have formed as a pseudomorph after primary gypsum, and a secondary a dolomitic fracture is infilled by a later generation of sparry calcite. Swallowtail pseudomorphs of calcite after gypsum, may indicate the an original bedded gypsiferous layer/horizon that was once present at 1102.2 m in the 81

stratigraphic section that has since been recrystallized as calcite (Figure 3.10, B). The occurrence of the swallowtail crystal habit indicates the primary precipitation of gypsum in the burrowed, nodular limestone bed at 1102.2 m. This once gypsiferous unit provides evidence of deposition under highly evaporative conditions. Additionally, the occurrences of intermingled gypsum, dolomite, and micritic calcite, cut by micritic calcite fractures, sampled from the uppermost paleosols of the Sangre de Cristo Formation suggest multiple generations of calcite precipitation and provide evidence of primary gypsum formation. Furthermore, petrographic observations of original gypsum and gypsiferous textures warrants the inclusion of a gypsic modifier on the

Calcisol classification of four paleosol profiles soils at stratigraphic positions of 1447.7 m,

1451.6 m, 1451.9 m, 1467.1 m, even though no original gypsum remains (Mack et al., 1993;

Michel et al., 2015). As such, petrographic observations of paleosols from the uppermost Sangre de Cristo reinforces the paleosol morphological interpretation that these soils formed under arid conditions within the stratigraphic sequence.

Paleosol clay mineralogy

The clay mineralogy of surficial deposits and soils is strongly correlated with regional climate (Yemane et al., 1996; Chamley, 1989; Srodon, 1999; Wilson, 1999). As a result, stratigraphic studies of clay mineralogical variation in sedimentary strata may serve as qualitative paleoclimate proxies (Yemane et al., 1996; Stern et al., 1997; Hartmann et al., 1999;

Tabor et al., 2002; Milleson et al., 2015). This is particularly useful in soils, as it reflects that the clay-size fraction may be generated either by (1) physical breakdown of parent material to a fine grain size or (2) chemical weathering that is facilitated by availability of aqueous solution (i.e., rainfall and groundwater) and energy (heat) that results in either transformation or neoformation 82

of a fine size fraction (Wilson, 1999; Rasmussen and Tabor, 2007). Cold (or hot) and dry climates are dominated by physical weathering, and therefore the clay-size fraction is approximately equivalent to that of the underlying parent material or surrounding country rock.

As a result, soils formed, cold and dry climates are typically dominated by chlorite, illite (or mica-like minerals), and quartz (Keller 1970; Wilson, 1999; Srodon, 1999). In environments characterized by greater amounts of moisture and higher temperatures, there is increasing hydrolytic reaction of the parent material, which results in progressive loss of base cations, de- silication, and neoformation of clay-sized minerals. As a result, weathering in climates with seasonal fluctuations of temperature and precipitation, such as the modern mid-latitudes, results in the formation of expansible and mixed-layer expansible 2:1 phyllosilicate clay minerals, whereas open-system weathering during periods of fresh water flushing, often in humid and warm climates, such as the modern tropics and sub-tropics, results in the formation of 1:1 phyllosilicates such as kaolinite or halloysite (Hancock and Taylor, 1978; Sommer, 1978; Tabor et al., 2002; Kormali and Abtahi, 2003; Yemane, et al., 2006).

Kaolinite, hydroxy-interlayered 2:1 expansible minerals (HIM), and illite comprise the dominant constituents of the clay-size fraction of paleosol matrix in soils from the Madera and

Sangre de Cristo formations. The presence and relatively high abundance of kaolinite in the upper Madera Formation at 518.0m up to 693.2 m, comprising ~18 % to ~48 % of the clay-size fraction, and in the middle interval of the Sangre de Cristo Formation from 1062.0 m to 1211.0 m (~21 % to ~54 %) suggests two periods of soil formation under warm, humid conditions in a well-drained soil profile (Figure 3.7). Hydroxy-interlayered minerals (HIM) are an equally important component of the paleosol clay mineral assemblages present in the Madera and Sangre

de Cristo formations. HIM formation in soils is associated with incipient chemical weathering under oxidizing conditions and in conjunction with frequent wetting and drying cycles (Wilson,

1999). HIM are found in varying abundances throughout the Madera and Sangre de Cristo formations, ranging from 4.0 % to 41.7 % of the clay-size fraction (Figure 3.7). The highest proportions of HIM are coincident with the highest proportions of kaolinite and provide additional evidence for pedogensis of soils in the upper Madera Formation (519 m to 693.2 m) and middle Sangre de Cristo Formation (1062 to 1211.0 m) portion of the stratigraphy under warm humid, well-drained conditions. Illite dominates the mineralogy of the clay-sized fraction of calcic paleosols in the middle and uppermost portion of the Madera Formation, and lower

Sangre de Cristo Formation, from 716.9 m to 1062.0 m and from 1211.0m to 1467.0m (Figure

3.7). The relative proportion of illite in the middle of the stratigraphy ranges from 39.1 % to

92.7 % of the clay-size fraction. The lowest proportions of HIM are coincident with the presence of gypsic features in Calcisols in the uppermost portion of the stratigraphy. From 1288.2 m to

1467.0m, the amount of HIM present in paleosol matrix samples is negligible, averaging less than 6.0%. Taken into consideration with the dominance of illite, and the low abundance of kaolinite, in the clay-size fraction, it is likely that the soils of the uppermost Sangre de Cristo

Formation developed under arid conditions, dominated by physical weathering.

Light Stable Isotope Compositions of Paleosol Calcite and Organic Matter

Oxygen Isotopes

Stable oxygen isotope values among the thirty-nine paleosol carbonates analyzed from the upper Madera and Sangre de Cristo formations range from -8.5 to 2.9‰ (PDB) (Figure 3.11;

Table 3.1). The majority of the samples in the lower ~830 m of the Sangre de Cristo Formation have carbonate 18O values of ~-5±1‰. Over a very short stratigraphic interval (~20 m) in the uppermost Sangre de Cristo Formation paleosol carbonate 18O values become substantially more positive, ranging from ~+1 to +3 ‰ (Figure 3.12).

The stable oxygen isotope composition of a particular carbonate mineral is determined by (1) the temperature of crystallization and (2) the stable oxygen isotope composition of the water from which the mineral crystallized. As mentioned, calcite is the dominant carbonate mineral in the lower ~830 m of paleosol nodules of the Sangre de Cristo Fm., whereas calcite and dolomite are mixed carbonate minerals in paleosol nodules in the upper ~170 m of paleosol nodules of the

Sangre de Cristo Formation (Table 3.X). This stratigraphic shift in the mineralogical composition of paleosol carbonate is coincident with the important change toward higher 18O values of paleosol carbonate values (e.g., Figure 3.12). Care was taken to select only calcite for stable isotope analysis, and the behavior of the reaction of carbonate powders with orthophosphoric acid was consistent with only the presence of calcite during CO2 extraction and purification (i.e., there was no sustained effervescence of CO2 after 24 hours of reaction of

25°C). However, estimations of weight percent calcite, determined from CO2 yields of carbonate powders reacted with orthophosphoric acid, range from 9.4 to 91.0 %. Weight percent calcite values of mixed carbonate samples of less than 50 % also suggests that some fraction of dolomite participated in the reaction. The inclusion of some dolomite within the analyzed powders of paleosol carbonate nodules and tubules that were collected from the upper ~170 m of the Sangre de Cristo Fm. may be an important aspect of the more positive 18O values in that part of the

stratigraphic section, because the stable oxygen isotope fractionation factors for calcite-water and dolomite-water are, respectively (for both mineral and water 18O on the SMOW scale):

3 (X) 10 ln(calcite-H2O) = 18.03*103/T(°K) – 32.42 (Kim and O’Neil, 1997), and

3 (X) 10 ln(dolomite-H2O) = 2.73*106/T(°K) + 0.26 (Vasconcelos et al., 2005).

3 The 10 ln values for calcite-H2O and dolomite-H2O are plotted in Figure 3.16 for temperatures ranging from 0° to 100° C (i.e. the temperatures of liquid water at earth-surface pressures). At isothermal temperatures of crystallization from water with the same stable oxygen isotope values, dolomite 18O values will be 2.9 to 4.0‰ more positive than co-precipitated calcite 18O values (Figure 3.16). In this respect, the mineralogical shift from calcite to dolomite alone could make up for about one-half of the per-mil measured shift toward higher 18O values of paleosol carbonate near ~880 m in the Sangre de Cristo Fm. However, the additional 2 to 4‰ of oxygen isotope enrichment would require either an ~10° to 20° C cooling or a 2 to 4‰ shift toward isotopically heavier waters within the soil carbonate nodules during deposition of the upper Sangre de Cristo Fm. A Permian-Carboniferous cooling of this magnitude seems unlikely considering the evidence for a substantial warming near the Permo-Carboniferous boundary and sustained warming during Early Permian time (Tabor and Montanez, 2005; Rosenau et al.,

2013). Therefore, a shift toward waters with isotopically more positive 18O values seems most likely. Furthermore, the 18O values of soil waters at typical earth-surface temperatures for

Permo-Carboniferous time (~25-35°C) would be near 2 to 4‰. Such positive soil moisture 18O values ordinarily aren’t encountered in modern soil systems except for those soils that are characterized by substantial amounts of soil moisture deficiency, evaporation, and evapotranspiration that leads to oxygen isotope enrichment of residual soil waters. Such soils are

often characterized by aridity. Significantly, this mineralogical shift toward dolomitic and calcitic compositions of paleosol carbonates, in conjunction with evidence for an isotopic shift toward enriched values consistent with net soil moisture deficiency and associated aridity is consistent with stratigraphic trends toward gypsic Calcisols and clay mineralogical trends of away from kaolinite and toward illite in the upper ~130 m that also is consistent with a trend toward aridity.

Carbon isotope values

Among the 39 samples of paleosol carbonate analyzed from the Sangre de Cristo Fm., 13C values exhibit a relatively narrow and negative range of values, -8.0 to -5.1‰ through the stratigraphic succession (Figure 3.13). These carbonate 13C values are relatively negative and exhibit relatively minor stratigraphic variability compared with values reported from penecontemporaneously-formed Lower Permian paleosol carbonates in western equatorial

Pangea (e.g., Tabor et al., 2005; Montañez et al., 2007).

It is notable that fossilized, vascular plant, carbon compressions are found only in the Madera

Formation, with only rare tree branch and trunk petrifactions and molds in the Sangre de Cristo

Formation. Nevertheless, fossil plant compressions are common in the Madera Formation and exhibit a narrow range of 13C values, from -24.1 to -22.4‰ (Figure 3.13, Table 3.2). While these 13C values are less negative than the carbon isotope values that typically are reported for

C3 photosynthesizing vascular plants on modern landscapes (e.g., Cerling, 1991), they correspond closely with stable carbon isotope values from fossil vascular plant compressions from penecontemporaneously-deposited Upper Carboniferous strata in other tropical and

paratropical basins (e.g., Montañez et al., 2007). Furthermore, there appears to be no clear stratigraphic trend or pattern to the observed variations in the 13C values in the Madera

Formation.

40 Dolomite Calcite 35

30 

1000ln 25

15 0 20 40 60 80 100 Temperature (°C)

Figure 3.16 Cross plot of 1000ln values versus temperature (°C) for 18O of water and dolomite (blue) and water and calcite (red). For isothermal reactions dolomite will be enriched in 18O by 2.9 to 4.1 per mil. See text for discussion.

1500

1400

1300

1200

1100

1000

900 Stratigraphic Stratigraphic Height (m)

800

700

600

500 15 20 25 30 35 40 45 Mean Annual Precipitation (MAP; cm)

13 Figure 3.17 Cross plot of stratigraphic position versus MAP estimates calculated from Δ C(CC- OM) calculated from paleosol carbonates and organic matter occluded within pedogenic carbonate nodules and tubules from the Madera and Sangre de Cristo formations in the Rowe-Mora basin.

Carbonate-associated organic matter samples from paleosol nodules and tubules in the upper

Madera and the Sangre de Cristo formations have 13C values ranging from -25.1 to -21.1‰, which is within the range of vascular plant organic matter stable carbon isotope values reported from penecontemporaneously deposited strata from tropical and paratropical Permo-

Carboniferous basins (e.g., Tabor et al., 2005; Montañez et al., 2007). While the stratigraphic pattern of carbonate-associated organic matter samples shows a rough trend from relatively negative, to more positive, and a return to more negative 13C values through the stratigraphic succession of the Sangre de Cristo Formation there is a lot of variability, up to ~2.5‰, from one sample to its nearest neighbor within that broad stratigraphic trend (Table 3.2; Fig. 3.13).

For the seventeen samples where calcite and co-existing carbonate-associated organic matter

13 13  C values have been determined, there is a range of  C(CC-OM) values ranging from 13.1 to

13 19.5‰ (Table 3.2). There is a stratigraphic pattern toward lower  C(CC-OM) values moving upward through the upper ~285 m of the Madera Formation and the lower ~ 110 m of the Sangre

13 de Cristo Fm., and stable carbon isotope values shift toward progressively larger  C(CC-OM) values in the upper 600 m of the Sangre de Cristo Formation. Fundamentally, measurements

13  C(CC-OM) values from paleosol carbonates have been the principle input parameter for determination f paleoatmospheric pCO2 estimates using the paleosol-carbonate paleobarometer model (Cerling, 1984; Yapp and Poths, 1996; Ekart et al., 1999; Tabor et al., 2005; 2013; 2017;

Montañez et al., 2007; Myers et al., 2012). In a very qualitative sense, the stratigraphic trends in

13  C(CC-OM) values are suggestive of declining paleoatmospheric pCO2 values during deposition of the lower 400 m, and increasing paleoatmospheric pCO2 values during deposition of the upper

13 600m, of the Permo-Carboniferous Sangre de Cristo Fm. However, a recent study of  C(CC-OM)

values from modern soils and paleosols of Permian and Triassic age suggests that carbon isotopic differences between paleosol calcite and co-existing organic matter is strongly correlated with rainfall, and mean annual precipitation (MAP) in the various locales in which soil carbonate

13 crystallizes (Tabor et al., 2017). Therefore, we pursue below this  C(CC-OM) proxy of MAP during Early Permian deposition of the Sangre de Cristo Fm.

13 As mentioned,  C(CC-OM) values in the Sangre de Cristo Range vary between 13.1 and 19.5

13 ‰. Tabor and others (2017) observed that  Ccc-om proxy values in modern C3-dominated soils from California exhibit the following relationship with MAP (in climates with MAP ranging from ~7 to 40 cm/yr precipitation):

13 (Eqn. 3.1) MAP = ( C(CC-OM) – 25.4)/0.31

13 Eqn. 3.1 in conjunction with the stratigraphic trends in paleosol carbonate  C(CC-OM) values correspond to paleo-MAP estimates that generally increase from ~26 to 40 cm/yr in the lower

~400 m, to paleo-MAP estimates that generally decrease from ~40 to 19 cm/yr in the upper 600 m of the Sangre de Cristo Formation. This pattern is generally consistent with an overall trend toward progressively dryer/more arid soils through the course of Early Permian time across the paleo-equatorial tropics of Pangea (e.g., Michel et al., 2015), and in western equatorial Pangea in particular (e.g., Parish et al., 1995; Tabor and Montañez, 2002; 2005; Tabor et al., 2002; 2008;

Tabor et al., 2017). Furthermore, this trend toward progressively dryer, more arid paleoclimate toward the top of the Sangre de Cristo Formation. that is based upon paleo-MAP proxies from

13 paleosol carbonate  C(CC-OM) values is consistent with other indicators of a similar trend toward aridity that is presented herein. These other indicators include stratigraphic trends toward (1) paleosol morphologies like Calcisols and gypsic Calcisols, (2) paleosol clay mineralogy away

from chemically weathered materials such as kaolinite and HIM toward illite-dominated materials and (3) a large stable oxygen isotope shift toward more positive values that is indicative of increased evaporation and aridity toward the top of the Sangre de Cristo Formation.

Conclusions

The occurrence and distribution of climate-sensitive geochemical indicators preserved in paleosols of the Madera and Sangre de Cristo formations of the Rowe-Mora basin provide evidence of a long-term aridification trend, punctuated by a period of increased seasonality, during deposition of the upper Pennsylvanian and lower Permian strata. The stratigraphic shift from argillic pedotypes, containing high relative proportions of kaolinite and HIM in the clay-

18 size fraction of paleosol matrix samples, relative negative δ O(CC) values of paleosol carbonates,

13 and MAP estimates calculated from Δ C(CC-OM) of paleosol carbonates and organic matter occluded within pedogenic carbonate nodules and tubules from the upper Madera and lower

Sangre de Cristo formations in the Rowe-Mora basin all suggest seasonal, humid conditions of pedogenesis during the Pennsylvanian-Permian transition. The prevalence of paleosols characterized by calcic and gypsic features, relatively high proportions of illite, paleosol carbonates significantly enriched with respect to 18O, and MAP estimates calculated from

13 Δ C(CC-OM) of paleosol carbonates and organic matter occluded paleosol nodules, ranging from

29.0 to 19.0 cm/year in the upper Sangre de Cristo Formation supports the inference that the

Rowe-Mora basin was subject to a period of significant aridification during Early Permian time.

Chapter 4

PERMO-CARBONIFEROUS PALEOCLIMATE OF THE CONGO BASIN: EVIDENCE FROM LITHOSTRATIGRAPHY, CLAY MINERALOGY, AND STABLE ISOTOPE GEOCHEMISTRY

Introduction

Over the past two decades, mounting sedimentological, mineralogical, geochemical, and chronologic evidence from various parts of Gondwana indicates that Permo-Carboniferous paleoclimate fluctuated between frigid (~0-8 °C), mesic (~8-15 °C), and possibly hyperthermic

(>22 °C), temperature regimes (e.g., Isbell et al., 2003; Scheffler et al., 2006; Fielding et al.,

2008; Tabor et al., 2011). Recent studies delineate a complex paleoclimatic history between the

Late Pennsylvanian (Gzhelian) and Middle Permian (Guadalupian) that includes several 106 yr cycles, alternating between large-scale glaciation events and periods with little to no continental ice. (e.g., Isbell et al., 2003; Montañez et al., 2007; Fielding et al., 2008; Montañez and Poulsen,

2013). Although these studies have been critically important in resolving the timing and magnitude of the Gondwanan glaciation, they are based entirely on marine and mixed marine- terrestrial successions. To date, relatively little is known about the evolution of Late

Pennsylvanian-Middle Permian paleoclimate in continental Gondwanan settings. This lack of

knowledge is especially true for sub-Saharan Africa, despite its vast record of Late Paleozoic glaciation (e.g., Visser, 1997).

This study presents lithostratigraphic, petrographic, mineralogical, and stable isotope data from a continuous sequence of Upper Carboniferous-Lower Permian lacustrine strata in the

Lukuga Formation, central Democratic Republic of Congo. The stratigraphic distribution and internal concentration of strata with outsized clasts delineates a single, large-scale pattern of climatic change from relatively cold to warm conditions through the entire succession.

Mineralogical composition of the <2µm fraction of sediments from the Lukuga Formation delineates a more complex pattern, with at least two stratigraphically long climate cycles, beginning with cold and then changing to warm conditions. Finally, the δ18O values of early burial lacustrine calcite cements from the Lukuga Formation delineate at least three, and possibly as many as six, climate cycles; each beginning with cold and then changing to warmer conditions. This regional reconstruction of Permo-Carboniferous terrestrial paleoclimate is similar to that interpreted from mixed marine strata in the Karoo and Kalahari basins (Scheffler et al., 2006), and strongly supports the notion that Gondwanaland was characterized by repeated shifts from glacial to inter- and/or non-glacial conditions (Montañez et al., 2007, Fielding, et al.,

2008). Furthermore, the magnitude of climate shifts inferred from this terrestrial study may provide a useful means of correlation with strata from contemporaneous marine basins in other near-field terrestrial basins in Gondwanaland, and to far-field, para-tropical terrestrial basins in

Euramerica.

Basis for Paleoclimate Reconstruction from the Lukuga Formation

The clay minerals in lacustrine basins are products of chemical weathering at Earth’s surface (e.g., Chamley, 1989). The type of clay mineral that is deposited in basinal settings reflects the intensity of chemical weathering in the hinterlands and sites proximal to deposition, which is strongly correlated with regional climate (Loughnan, 1975; Gill and Yemane, 1996;

Yemane et al., 1996; Jackson, 1964; Chamley, 1989; Srodon, 1999; Wilson, 1999). As a result, studies of stratigraphic variation in clay mineralogy may serve as qualitative paleoclimate proxies and provide a semi-quantitative indication of paleoclimate (Yemane et al., 1996; Stern et al., 1997; Hartmann et al., 1999; Tabor et al., 2002). In addition, lacustrine calcite 18O values are strongly correlated with 18O values of meteoric water and temperature. Furthermore, meteoric water 18O values show a strong correlation with mean annual surface atmospheric temperatures ~ ≤ 15°C (MAAT; Dansgaard, 1964; Rozanski, 1993; Ferguson et al., 1999), and therefore 18O values of ground- and lake-water calcite cements may provide a quantitative proxy of MAAT (Welhan, 1987; O’Neil, 1987; Gregory et al., 1989; Hays and Grossman,

1991;Yemane and Kelts, 1996; Ferguson et al., 1999; Bhattacharya et al., 2002; Tabor et al.,

2006). This work documents stratigraphic trends in mineralogy of Pennsylvanian-Permian

(~300-250 million years ago) strata, and the oxygen isotope composition of finely-disseminated, early-burial, calcite cements from a continuous, 1860 m long, core (named Dekese) near the center of the Congo Basin, West Africa.

Mineralogy of bulk powders and clay-size fractions (<2μm e.s.d.; equivalent spherical diameter) of the Lukuga Formation in the Dekese core are divided into three ascending units, named Zones 1-3. Zones 1 and 3 contain clay minerals that are indicative of limited chemical 95

weathering, whereas zone 2 contains clay minerals that are indicative of virtually no chemical weathering.

Calcite cement 18O values are relatively negative, and suggest a frigid climate (MAAT ~

0 °C) during deposition in zones 1 and 2, whereas zone 3 calcite 18O values suggest significantly warmer climate (MAAT ≥ 15 °C). The results of this study are compared with paleoclimate proxy data from contemporaneous high-latitude basins in Gondwana (Malawi,

Tanzania, Antarctica, India and Oman). Clay mineralogy and calcite 18O values from the

Dekese core strata, when considered in conjunction with data from other basins in

Gondwanaland, provide a basis for evaluation of the evolution of paleoclimate, and its possible relation to dynamics of glacial ice, during Permo-Carboniferous time.

Geologic Background

Tectonics

The Congo Basin covers approximately 1.2 million km2 in central Africa, and contains up to 9 km of sedimentary strata. Congo Basin deposits range from Early Cambrian to Recent, and are bound by the eastern Mitumbe (Kibaran fold belt), and western Mayombe (West Congolian fold belt), mountain ranges. The region has been subjected to two major deformational events related to marginal tectonics. The first coincides with the Pan-African Collision in western

Africa at the end of the Cambrian (Giresse, 2005). The later deformation event occurred in the

Late Paleozoic, which led to tectonic inversion and widespread erosion of basinal strata and development of a major regional unconformity (Giresse, 2005). The Congo Basin is considered

to be a thermal sag basin that subsided in response to crustal cooling following each deformational event (Giresse, 2005). Regionally, the basin is surrounded by a variety of tectonic influences: Proterozoic thrust belts, a Mesozoic passive margin, and a Cenozoic rift system (Daly et al., 1992).

Basin Fill

The most comprehensive tectonic/structural description of the Congo Basin is that of

Daly and others (1992). The basement sequence includes Archean to Proterozoic crystalline basement rocks. The overlying lithostratigraphy is divided into 5 sequences, starting above the basement rocks and ending with the most recent deposits. Sequence 1 is comprised of stromatolitic carbonates and evaporites, which have been correlated with carbonates of the Ituri

Group in NE-Democratic Republic of the Congo. Sequence 2 is characterized by fining upward marine clastic sediments which prograde into the basin from the west. Both sequence 1 and 2 experienced contractional deformation during the Early Cambrian Pan African orogeny

(Kennedy, 1964), which led to a period of increased erosion, and development of a regional unconformity beneath the overlying sequence. Sequence 3 is composed of Late Ordovician to

Devonian marine shales and arkosic sandstones. These deposits correlate with Aruwimi Group strata in NE-Democratic Republic of the Congo. Sequence 4 conformably overlies sequence 3 and is made up of Carboniferous-Permian glacial deposits, which correlate with the Dwyka and

Ecca Groups of the Main Karoo Basin in South Africa (Catuneanu et al., 2005). The second major contractional deformation event occurred after the deposition of sequence 4, again

resulting in increased erosion and a major regional unconformity. Sequence 5 ranges in age from

Triassic to Recent and includes deposits corresponding to Upper Karoo aged sediments.

Lukuga Formation

The Congo Basin migrated northward, from 65°S to 40°N, between Early Carboniferous

(~340 Ma) and Late Permian (~250 Ma) time. Permo-Carboniferous strata (Sequence 4) in the

Congo Basin are named the Lukuga Formation (Cahen and Lepersonne, 1980), and they are among the most northwestern rocks of the Karoo Supergroup. Lukuga Formation strata are known from (1) exposures along the basin margin, and (2) from a continuous, 1860 m-long, core taken near Dekese, Democratic Republic of Congo (Figure 4.1).

Lukuga Formation outcrops are preserved within U-shaped, glacial valleys in the extreme eastern margin of the Congo Basin (Boutakoff, 1948). In ascending order, these strata are divided into five lithostratigraphic units: (I) conglomerates and sandy polymicts interpreted to be wash-out tillites, (II) Polymictitc argillites interpreted as seasonal, varved, lacustrine strata with dropstones, (III) dark-colored, pyrite- and kaolinite-rich, mudstones with common vascular-plant fossils, (IV) coal measures and (V) “Transition Beds” (Boutakoff, 1948; Cahen and Lepersonne,

1980). Lithostratigraphic units I-III range from 140-160 m thick in the eastern Congo Basin outcrops, and they are collectively called Kiralo ya Mungu beds (Catuneanu et al., 2005; Lower

Subgroup of Cahen and Lepersonne, 1980); they are interpreted to represent glacial, peri-glacial and post-glacial paleoenvironments/paleoclimates (Boutakoff, 1948; Cahen and Lepersonne,

1980; Wopfner, 1999). In particular, Unit III is similar in composition to coeval strata in other

basins throughout central and eastern Gondwanaland (Wopfner, 1999), and they are collectively interpreted to represent a single, contemporaneous, deglaciation event across most of

Gondwanaland near the Asselian-Sakmarian boundary (end of Glacial III of Isbell et al., 2003).

Dekese Core

The Dekese core was collected from the deepest part of the Congo Basin in the 1950’s by the Syndicate pour l’Etude Géologique et Miniere de la Cuvette Congolaise, and is currently archived at the Central African Museum in Tervuren, Belgium. Lukuga Formation strata are over

960 m thick in the Dekese core, and are divided into four informal ascending members: Couche

G (1677.15 – 1553.77 m below surface), Couche F (1553.77 – 860.63 m below surface), Couche

E (860.63 – 753.04 below surface) and Couche D (714.78-753.04 m below surface (Cahen et al.,

1960). Couche G is an upwards-fining unit with brown-red tillite at the base that is overlain by dominantly brown to red sandstone, with or without disseminated pebbles and conglomerate.

Figure 4.1. Map of Africa in its modern continental configuration with the paleogeographic positions of South America, Madagascar, and East Antarctica projected to the west and to the east. The thick, dark lines depict current political boundaries in Africa, whereas the thin vertical and horizontal lines approximate 20°E longitude and 1°N latitude. Outcrops of Upper Paleozoic glacial deposits are depicted in black and subsurface extensions of the strata are shown in orange. Red arrows indicate approximate directions of glacial ice advance in the Late Paleozoic with respect to contemporary sedimentary basins. The blue circle represents the approximate position of the Dekese borehole, which provided the core sampled in this study. 100

There is an undulatory contact between the topmost argillaceous strata of Couche G and the basal conglomeratic layer of Couche F. Couche F is dominantly gray to black, finely laminated claystone interbedded with clay-rich to sand-rich polymictites. The laminated claystones have been interpreted as glacial varveites, whereas the polymictites are interpreted to be glacial dropstones. The upper contact of Couche F is scoured. Couche E is dominantly a sandstone-rich lithology, grading upward from poorly sorted brown to brownish-red clayey sand, with or without pebbles, to yellow-brown, clayey to fine-grained sand. The rocks at the base of Couche

E share similarities with those of the Couche F. Couche D represents the uppermost portion of the Lukuga Formation present in the core. Couche D alternates between brown-yellow sandy clays and very fine-grained mudstones. Note that although the Lukuga Formation strata from the

Dekese core are coeval with outcrops of the Kiralo ya Mungu beds in extreme eastern Congo

Basin, Dekese core strata do not preserve the characteristic “deglaciation” facies of that observed in the Kiralo ya Mungu beds in extreme eastern Congo Basin and in many other basins across

Gondwanaland (Wopfner, 1999, Catuneanu et al., 2005).

Methods

Sampling

Cuttings from the Dekese core Lukuga Formation strata (n = 100) were collected by N.

Tabor and S. Myers at the Royal Central African Museum, Tervuren, Belgium. Generalized lithological descriptions were made as a compliment to previous work done with this core material (e.g., Cahen and others, 1960). Although examples of all lithologies were sampled for

101

mineralogical analysis, particular attention was given to identification of those parts of the stratigraphy that include calcareous cements. The presence of calcareous cement was determined by applying 1 drop each of 1N HCl and 5N HCl about every 10-20 cm of core length.

Effervescence resulting from either acid treatment was noted and marked as a stratigraphic position for sampling and subsequent petrographic, mineralogical, and stable isotope analyses.

Samples were collected by dry cutting ~1/3 to 1/4 of the core using a hand-held diamond saw.

Thin sections for petrographic analysis were either borrowed (n = 75) from the Royal

Central African Museum or prepared from cuttings (n = 25). Samples were analyzed using transmitted plane and cross-polarized light microscopy to describe sedimentary fabrics and cement textures.

U-Th-Pb isotope systematics of detrital zircons

Detrital zircon samples were isolated and analyzed for radioisotopic ages at Boise State

University, using bulk matrix samples collected at 1680.0, 1578.1, 1568.5, 1040.0, 1010.25,

874.0, and 830.0 m within the Dekese core. In situ analysis of zircon U-Th-Pb isotope systematics and trace element compositions used a ThermoElectron X-Series II quadruple

ICPMS and New Wave Research UP-213 Nd:YAG UV (213 nm) laser ablation system. Analysis utilized in-house analytical protocols, standard materials, and data reduction software for simultaneous acquisition and real-time calibration of U-Th-Pb ages and a suite of HFSE and

REE elements using the high sensitivity and unique properties of the interface (Xs cones), extraction lens, and quadruple analyzer of the X-Series II. Zircons are ablated with a laser

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diameter of 25 µm using fluence and pulse rates of 10 – 15 J/cm2 and 10 Hz, respectively, during a 60 s analysis (15 s gas blank, 45 s ablation), which excavates a pit approximately 25 µm deep.

Ablated aerosols are carried by a 1 L/min He gas stream to the plasma. Dwell times are 5 ms for

Si and Zr, 40 ms for 49 Ti, 238U, 232Th, 202Hg, and all Pb isotopes; and 10 ms for all other HFSE and REE elements. Background count rates for each analyte are obtained prior to each spot analysis and subtracted from the raw count rate for each analyte.

For concentration calculations, background-subtracted count rates for each analyte are internally normalized to 29Si, and calibrated with respect to NIST 610 and 612, and USGS BCR-

2, and BIR-1 glasses as the primary standards. For U-Th-Pb age analysis, instrumental fractionation of the background-subtracted 206Pb/238U, 207Pb/206Pb, and 208Pb/232Th ratios is corrected, and ages calibrated with respect to interspersed measurements of the Plešovice zircon standard (Sláma et al., 2008). The FC-1 Duluth Complex Anorthite Series zircon is analyzed in every experiment as a secondary check standard. Signals at mass 204 are indistinguishable from zero following subtraction of mercury backgrounds measured during the gas blank (<1000 cps

202Hg); thus, ages are reported without common Pb correction. However, age error propagation includes an uncertainty contribution due to common Pb using the absolute value of the measured

206Pb/204Pb ratio. Radiogenic isotope ratio and age error propagation for each detrital grain analysis includes uncertainty contributions from counting statistics, background subtraction, common Pb correction, and standard calibration (based on the standard deviation of the isotope ratio measurements of the standard over the course of the experiment). Error propagation for non-detrital spot analyses exclude standard calibration uncertainty, which is instead propagated in quadrature following group statistics (e.g. weighted mean calculations); standard calibration

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uncertainties range from 1% to 2% (2σ) for a given experiment. Spot ages for the zircon standard

R33 run as an unknown (calibrated using Plešovice as primary standard) and indicate a 3 – 4%

(2σ) external reproducibility for the method.

X-ray Diffraction

Cuttings (n = 100) were powdered in a corundum mortar and pestle and divided into two aliquots. One aliquot was analyzed via X-ray diffraction of bulk powders. The second aliquot was disaggregated by shaking overnight in a dilute sodium carbonate solution (pH = 9.2) and the clay-size fraction (<2 μm equivalent spherical diameter; e.s.d.) was isolated via gravity- suspension settling. Each clay-size fraction was divided into four aliquots and mounted as oriented aggregates via the filter-membrane technique of Drever (1973) as follows: (1) K- saturated with 1.0 M KCl solution at room temperature (~23 °C), (2) K-saturated with 1.0 M KCl solution and heating to 500 °C for 2-4 hours, (3) Mg-saturated with 0.5 M with MgCl2 solution,

(4) Mg-saturation with 0.5 M MgCl2 solution and glycerol solvation with a 1:4 Glycerol: H2O solution. X-ray diffraction analyses of powder mounts include scans between 2-70° 2θ, with a scan speed of 3° 2θ/minute. X-ray diffraction analyses of oriented aggregates include scans between 2-302 and a scan speed of 2° 2θ /min. All X-ray diffraction analyses were made with a Rigaku Ultima III X-ray diffraction system using CuK radiation and scan window of 0.04°

2θ. Mineral identification follows the methods outline in Moore and Reynolds (1996) and Brown and Brindley (1984). The relative abundance of minerals within each <2 m sample was calculated using the mean area under the peak of its first order basal (001) reflection. The area

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under each peak was approximated by multiplying the peak height by the full peak width at half- height (FWHM). Each area was taken as a proxy for the integrated peak intensity (Moore and

Reynolds, 1996). . A/tot = % Chl, B/tot = %Ill, where A= the area under the 7Å peak, B = the area under the 10Å peak, and tot = A + B. This calculation does not include minor constituents present in some samples, such as 2:1 expansible phyllosilicates, kaolinite, and non-phyllosilicate minerals. Occurrences of rare phyllosilicate minerals are reported in conjunction with the proportional abundance of major phyllosilicate constituents throughout the stratigraphy (Figure

4.2).

A hydrazine-DMSO intercalation procedure was used as a definitive test for distinguishing between kaolinite and chlorite in oriented aggregates of the <2 m fraction (n =

24). The intercalation procedure used herein is adapted from those of Calvert (1984) and Lim and others (1981). Specifically, ~50 mg of dried <2 m e.s.d. sample was mixed with 125 mg of

CsCl for 3 minutes, followed by mixing of an additional 125 mg of CsCl for an additional 3 minutes, in an agate mortar and pestle. The mixture was transferred to a 50 ml Teflon centrifuge tube, and combined with 5 ml of 85% (by volume) hydrazine monohydrate that was maintained at 80 °C in a hot water bath for 25 minutes. The suspension was centrifuged at 1500 rpm for 25 minutes, then decanted to remove excess hydrazine. Immediately, 5 ml of DMSO solution was added and kept at 90 °C for 20 minutes. The suspension was centrifuged at 1500 rpm for 25 minutes and the supernatant DMSO was decanted. An additional 5 ml of DMSO was added to each sample and kept at 90 °C for 12 hours. Samples were then removed from the oven, centrifuged at 1500 rpm for 25 minutes, decanted, and washed with methanol via centrifugation and decanting of the supernatant. Finally, a slurry of the sample and water was applied to a glass

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slide and allowed to dry in an oven at 60 °C. The samples were immediately analyzed by x-ray diffraction analysis after drying, and again 24 hours later. Clay mineral fractions treated by

DMSO and analyzed by X-ray diffraction were taken from the following depths, given in meters, in the Dekese core: 820.45, 837.9, 867.5, 867, 868.5, 884, 903, 911.7, 945.25, 961, 975.2,

1035.5, 1190, 1286.7, 1293, 1305.9, 1365.5, 1419, 1432, 1471.5, 1479.8, 1492, 1496, 1539.8.

Isotope geochemistry

The presence of calcite in Lukuga Formation strata from the Dekese core samples (n =

34) was determined by macroscopic and petrographic observation, and these samples were processed for isotopic analysis. Vein calcite was isolated from core cuttings using a handheld

Dremel rotary tool with a 30 µm diamond-tipped drill bit, and 5 to 30mg of sample was processed for isotopic analysis. For samples containing disseminated calcite, between 10 and

100 mg of powdered bulk sample was used for analysis. Samples were reacted overnight with

100% anhydrous phosphoric acid at 25 °C (McCrea, 1950). Reaction products were cryogenically purified within a manual, high-vacuum extraction system. CO2 yields were

13 18 determined via mercury manometry, and  C and  O values of cryogenically-purified CO2 gas from the samples were measured at Southern Methodist University using a Finnigan-MAT 252 isotope ratio mass spectrometer. Isotope values are reported in standard delta notation:

18 13  O (or C) = (Rsample/Rstandard -1)*1000,

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where R is the ratio of heavy-to-light stable isotope present in the sample or standard.  values are reported relative to the Peedee Belemnite standard (PDB; Craig, 1957) for carbon and oxygen isotope values.

Results

Lithostratigraphy

Cahen and others (1960; Cantuneanu et al., 2005) provide a very detailed description of the Lukuga Formation in the Dekese core strata, which includes angular to brecciated quartz and feldspar sand and silt grains (Figure 4.3a), mm-scale sandstone-mudstone couplets interpreted as glacial varves (Figure 4.3b-d, 4.4a-c), and outsized clasts interpreted as ice-rafted material

(Figure 4.3c,d, 4.4), and rare calcareous cements (Figure 4.3d, 4.5). This study provides a description of outsized clast and carbonate textures of the Lukuga Formation lithostratigraphy that is pertinent to the interpretation of mineralogical and geochemical data from the Dekese core. A detailed list of samples collected from the Dekese core and the results of analyses used in this study are presented in Table 1.

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Figure 4.2. Relative stratigraphic position versus cumulative clay mineral percent of major constituents in the <2µm phyllosilicate fraction of samples from the Lukuga Formation. The y- axis does not follow a uniform scale. Rather, these data are plotted with equal space between all samples analyzed, and the position of key stratigraphic depths to inform the density of information with respect to depth in the core. The red area represents the relative area beneath the 14 Å chlorite peak, the blue area represents the relative area beneath the 7 Å chlorite peak, and the yellow area corresponds to the area beneath the 10 Å illite peak. Opposing changes in the relative areas of the 7 Å and the 14 Å peaks indicate the presence of kaolinite. Horizontal black lines mark the stratigraphic position of samples that contain some expansible 2:1 phyllosilicates. The white line represents the single occurrence of kaolinite. 108

Outsized clasts, observed in hand specimen and by petrographic inspection of thin sections, exhibit virtually no internal compaction, yet surrounding sedimentary laminations exhibit evidence of significant deformation around the outsized clasts (Figure 4.3c, Figure 4.4).

The distribution of outsized grains occurs sporadically throughout the Lukuga Formation but are particularly concentrated within several zones in the stratigraphy, ranging from -1573 to -1539 m, -1511 to -1503 m, and -1429 to -1276 m (Figures 4.3, 4.4). Sedimentary strata with alternating laminae of clay-sized material and silt or very fine sand occur in two broad zones from -1525 to -1426 m and -1270 to -924 m (Figures 4.3, 4.4, 4.5). Calcite cements occur sporadically through the Lukuga Formation stratigraphy of the Dekese core, but are generally more abundant and stratigraphically more extensive than polymictic strata or varved strata

(Figure 4.5).

Mineralogy

Bulk X-Ray Powder Diffraction Analyses

Powder diffraction spectra of bulk samples exhibit intense peaks at 6.4, 4.26, 4.04, 3.34,

3.1 – 3.2, 3.7, 3.85, and 3.00 Å. Peaks at 4.26 and 3.34 Å correspond to d(100) and d(101) Miller indices, respectively, of quartz (SiO2). Peaks at ~6.4 Å, 4.04 Å, and 3.1 – 3.2 Å likely correspond to d(001), d(201), d(202) and d(040) Miller indices, respectively, of calcian albite ((Na, Ca)(Si,

Al)4O8). Peaks at ~3.85 Å and ~3.00 Å correspond to d(102) and d(104) Miller indices, respectively, of calcite. All samples also exhibit relatively low and broad peaks near 14, 10 and 7

Å, which correspond to d(hkl) spacings of 2:1 and 1:1 phyllosilicates.

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Table 1. List of lithologic and mineralogical characteristics of samples collected from the Lukuga Formation in the Dekese core. Polymict and varve features identified from hand- specimen examination and petrographic microscopy. Presence of chlorite (7 Å and 14 Å) and illite (10 Å) peaks verified using X-ray diffraction (XRD) of the clay-sized (<2 µm) fraction of matrix samples. Presence of expandable peaks refers to expansion of the 10 Å peak after Mg saturation and glycerol solvation. Kaolinite peak identified by shift from 7 Å to ~8 Å after DMSO treatment. Occurrence of calcite was determined from hand-specimens (reaction with dilute HCl), petrographic microscopy, and XRD analysis of bulk powders and clay-size sample fractions.

Depth (m) Polymicts Varves 7Å 14Å 10Å Expandable Kaolinite Calcite peak peak peak peak peak

714.8 X 721.5 X 725.3 X 727.5 X 730.2 X 731.7 X 734.8 X 746 X 748.5 X 762.9 X 763 X 764 X 770 X 782.25 X 785.25 X 801.9 X 808.1 X 816.7 X 820.45 X X X X 824.5 X 837.9 X X X X 839.9 X 852.5 X 859 X 867 X X X X X 868.5 X X X 874 X X X X 880.8 X X X X 884 X X X X

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Depth (m) Polymicts Varves 7Å 14Å 10Å Expandable Kaolinite Calcite peak peak peak peak peak

886.25 X X X X X 893 X X X 901.3 X 901.9 X X X X 903 X X X 905 X X X X 911.5 X X X X X 911.7 X X X X X 912 X X X X X 916 X X X 923.3 X X X X 924.5 X X 930 X 935 X 941.75 X 945.25 X X X X 952 X 954 X X X X 961 X X X X X X 968 X 973.7 X 975.2 X X X X 977.3 X 984 X 985.3 X 988.3 X 1010.25 X 1026.25 X 1035.5 X X X X 1048 X 1054 (97- X X X X 1058.85) X X X X 1062.5 X 1071.25 X X X X 1074.5 X 1093 X 1098 X X X 1107.25 X X X X 1118.3 X X X 1129.5 X X X 111

Depth (m) Polymicts Varves 7Å 14Å 10Å Expandable Kaolinite Calcite peak peak peak peak peak

1140 X X X X 1180.4 X X X X 1184.3 X X X X 1190 X X X X 1202 X X X 1227 X X X X 1243 X X X 1265.5 X X X 1270.3 X X X X 1275.75 X X X X X 1280 X X X X 1283.8 X X X X X 1286.7 X X X X X 1293 X X X X 1294.25 X X X X 1296 X 1298 X X X X 1305.9 X X X X X 1310 X X X X X 1313.6 X X X X 1316.2 X X X X 1322 X X X X 1329.85 X X X X 1334.25 X X X X 1341 X X X X X 1343.25 X X X X 1348.3 X X X X 1350.5 X X X X 1356.5 X X X X 1361 X X X X 1365.5 X X X X 1369.75 X X X X 1375.75 X X X X 1379.8 X X X X X 1385 X X X X 1386 X X X X 1390.5 X X X X 1394 X X X 1396 X 1399.7 X X 112

Depth (m) Polymicts Varves 7Å 14Å 10Å Expandable Kaolinite Calcite peak peak peak peak peak

1405.5 X X X X 1409 X X X X 1419 X X X X X 1426.35 X X X 1432 X X X 1437 X X X 1442 X X X X 1448 X X X 1453 X X X 1454.5 X X X 1458.25 X X X X 1458.8 X X X X 1471.5 X X X X X 1476 X X X X 1478.3 X X X X 1479.8 X X X X 1482.7 X 1485 X X X X 1492 X X X X 1496 X X X X X 1501 X X X X 1503 X X X X X 1506.5 X X X X 1511.5 X X X X 1513 X X X X X 1517.3 X X X X 1521 X X X X X 1524.5 X X X X 1525 X X X X 1527 X X X X 1531 X X X X 1535 X 1536.7 X 1537 X X X X X 1538.5 X 1539.8 X X X X 1542 X 1545 X 1545.5 X 1555.2 X 113

Depth (m) Polymicts Varves 7Å 14Å 10Å Expandable Kaolinite Calcite peak peak peak peak peak

1557 X 1561.2 X 1566.6 X 1568.5 X 1569 X 1573.5 X X 1576.4 X 1579 X 1600 X X X 1662 X 1668 X

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Figure 4.3 Examples of common sedimentary structures and textures in the Luka Formation in the Dekese core. (A) Very fine sand- and silt-size grains of detrital igneous quartz (white and yellow), chert (pale yellow and dark bladed crystal), and feldspar grains (light gray/white and dark twins) in a dark, organic-rich clay matrix sampled from 1051 m depth, showed under crossed Nichols. (B) Plane light image of alternating and coupled (i.e., varved) lamina of banded quartz siltstone (white layers) and claystone (dark layers) from 1426 m depth. (C) Plane light image of poorly developed, varved laminae with very fine sand- to coarse silt-size grains in clay- rich layers from 1512 m. (D) Plane light image of varved clay-rich and fine to medium siltstone laminae, some of which are deformed by compaction around microcrystalline calcite (micrite) cemented nodules. Sample collect at 1093 m depth.

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Figure 4.4 Examples of polymicts collected from the Lukuga Formation in the Dekese core. (A) Polycrystalline outsized detrital grains in varved and slightly deformed argillaceous laminae collected at 1270 m depth, imaged under crossed Nichols. (B) Crossed Nichols image of polycrystalline outsized grain in varved and deformed laminae sampled at 1426 m. (C) Angular fine sand- and coarse silt-size detrital grains and detrital polycrystalline medium sand-size grain in organic-rich claystone from 1500 m, imaged under crossed Nichols. (D) Plane light image of laminated clay and silt varveites with coarse sand-size drop stone layer of polycrystalline quartz from 1098 m.

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Figure 4.5. Examples of carbonate cements collected from the Lukuga Formation in the Dekese core and imaged under crossed Nichols. (A) Irregular microcrystalline calcite nodules around which varved and laminated strata of argillaceous and coarse silt-size sediments are slightly deformed. Sample collected at 1227 m depth. (B) Microcrystalline calcite zones cementing coarse silt-size grains within an organic-rich and opaque argillaceous matrix at 1500m. (C) Granular calcite composed of radiaxial cements that displace and occlude detrital sediments from 1253 m. (D) Radiaxial fibrous calcite cement that mostly occludes detrital sedimentary grains. Sampled at 924 m depth.

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Clay-size (< 2μm e.s.d) fraction analyses

All samples contain chlorite, as indicated by basal reflections at approximately 14.00,

7.00, 4.65, and 3.50 Å. Although the 7 Å peaks remain after heating to 500 °C, the peak intensities significantly decrease (Figure 6A). Ordinarily, collapse of the 7 Å peak indicates that kaolinite is co-mingled with chlorite (Brown and Brindley 1984), but this indicator is not entirely conclusive, because Fe-rich chlorite may exhibit only partial collapse of its 7 Å peak after heating to 500 °C (Moore and Reynolds, 1996). DMSO intercalation expands the kaolinite d(001) peak from ~7 Å to ~11 Å, and therefore provides a diagnostic indication for the presence of kaolinite when it is mixed with chlorite (Lim et al., 1981; Calvert et al., 1984). Application of the DMSO intercalation method to clay fraction samples from the Lukuga Formation indicates that only one sample (at 962 m in Figure 4.6) contains kaolinite. All other samples that exhibit a

~7 Å peak are attributed to the presence of Fe-Chlorite.

All samples contain illite, as indicated by the presence of d(001) and d(002) basal reflection peaks at ~10.0 Å and ~5.0 Å, respectively (Moore and Reynolds, 1996). The breadth and symmetry of illite d(001) peaks varies among, and within, samples according to the chemical pretreatment used. Specifically, some XRD spectra of the clay samples exhibit increased symmetry and kurtosis of the illite d(001) peak after Mg-saturation and glycerol solvation compared to K-saturation (Figures 4.7, 4.8). This behavior, which is indicative of the presence of poorly sorted 2:1 expansible phyllosilicates (Moore and Reynolds, 1996), occurs in samples collected from -1537 to -1442 m and -961 to -837 m. Calcite was identified in clay samples by the presence of XRD peaks at ~3.03 Å in K-saturated samples, which were not pretreated with

HCl. These peaks are absent in Mg-saturated samples, in which the HCl pretreatment dissolved

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the calcite (Figure 4.9). Figure 2 shows the stratigraphic variation in the relative abundances of the clay minerals.

Detrital Zircons

Detrital Zircons were collected from samples at 1680, 1578, 1568.5, 1260, 1040, 1010,

874, and 838 m below the top of the core. All samples have a population of zircons with age estimates from ~2.5 to 3.1 Ga, from ~1.9 to 2.2 Ga, from ~1.0 to 1.4 Ga, and from ~0.5 to 0.7 Ga

(Figure 4.10). All but the two lowest/oldest samples (1680 and 1578 m) also have a population of zircons with age estimates ranging from ~0.8 to 0.9 Ga, whereas the two shallowest/youngest samples at 838 and 874 have a small population of detrital zircons near 375 and 305 Ma, respectively (Figure 4.10).

Calcite

Macroscopic, transmitted light microscopic, and scanning electron microscopic inspection of calcareous samples indicate that three different calcite textures exist within the

Dekese core. Texture 1 comprises coarsely crystalline sparry calcite that is deposited as fracture- fills. Texture 2 is composed of micrite- and microspar-size crystals that occur either as indurated nodules within mudrock layers, or as finely disseminated, pore-filling cements that occlude grains in siltstones and sandstones. Texture 3 is composed of radial, fibrous calcite cements organized into pseudospherulitic bundles ~0.5 mm across (Figure 4.5D). These cements with

Texture 3 form stratiform features up to 1 cm thick. 119

Coarsely crystalline Texture 1, most abundant from 923 to 802 m, is consistent with crystallization during burial, and is often associated with geopressured waters (Bathurst, 1975).

Texture 2 consists of euhedral crystals with sharp crystal terminations, suggests that calcite crystals grew in-situ, and are not detrital grains. Furthermore, very early crystallization of these cements is implied by differential compaction of surrounding, uncemented sedimentary laminations (Figure 4.5). Finely disseminated micritic calcite cements of Texture 2 comprise between ~2 to 91% of the total sample weight (Table 4.2), and are concentrated in three stratigraphic zones from 1579 to 1537 m, from 1419 to 1275 m, and from 1093 to 715 m below the top of the core (Figure 4.6).

The relative purity of radial fibrous calcite cements that characterize Texture 3 (Figure

4.5D) may be attributable to displacive growth within the surrounding sediments, but these cements also partially occlude small pyrite crystals, and approximately 3 to 5% of the thin- section areas of these cements contain silt-size quartz grains (Figure 4.5d). Radial fibrous cement is typically associated with early replacement of aragonite and high-Mg calcites in marine environments (Mazzullo, 1980). However, radial fibrous calcite also occurs as primary, early groundwater cement in terrestrial environments (Mozely and Davis, 2005) and in benthic lacustrine environments characterized by long intervals of non-deposition (Rossi and Canaveras,

1999).

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Figure 4.6. Cross plot of the stratigraphic position versus δ18O values of calcite samples collected from the Lukuga Formation in the Dekese core. Black circles represent δ18O values of disseminated calcite, and red square represent δ18O values of calcite in displacive veins. Vertical dark lines are isotherms that reflect mean annual temperature (MAT) inferred from δ18O values of calcite. Blue shading denotes stratigraphic intervals where calcite cement δ18O values indicate frigid conditions. Horizontal green lines show the stratigraphic positions of polymictic strata, and orange lines indicate stratigraphic positions of varved layers. 121

Figure 4.7. Background-subtracted X-ray diffractograms of oriented aggregates from the clay- size (<2µm) fraction of a matrix sample collected from the Lukuga Formation at 961 m depth in the Dekese core. The clay-size fraction of the sample was K-saturated and analyzed at room temperature (middle spectrum), K-saturated and heated at 500°C for 2 h prior to analysis (lower spectrum). The origin of the collapsible 7.15 Å peak in the K-saturated spectrum is ambiguous and may be attributed to either Fe-chlorite or kaolinite. DMSO treatment expands part of the kaolinite d(001) peak to 11.21 Å, providing a diagnostic text of the presence of kaolinite.

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Figure 4.8. Background-subtracted X-ray diffractograms of oriented aggregates from the clay- size fraction (<2µm) fraction of a matrix sample collected from the Lukuga Formation at 1072 m depth in the Dekese core. The lower spectrum reflects Mg saturation only, whereas the upper spectrum illustrates the effects of Mg saturation and glycerol solvation. The broad-shouldered, illite-dominated 10.05 Å peak in the Mg-saturated spectrum becomes sharper and more symmetrical under Mg saturation and glycerol solvation. This behavior is indicative of poorly amorphous, expansible 2:1 phyllosilicates.

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Figure 4.9. Background-subtracted X-ray diffractograms of oriented aggregates from the clay- size (<2µm) fraction of samples from the Lukuga Formation at 905 m depth in the Dekese core. The lower spectrum reflects K saturation of the clay-size fraction, and the upper spectrum shows the results of Mg saturation. The Mg-saturated sample was treated with 1N HCl solution prior to Mg saturation, dissolving calcite and removing the 3.03 Å peak originally present in the K- saturated sample.

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ed at 838 m (red), 874 m (black), m 874 m m 838 1010 (red), at ed

Pb age of detrital zircons collected from sedimentary rocks of rocks the of zircons from sedimentary detrital collected Pbage

Figure 4.10. Plot of relative probability versus U Plot 4.10. probability relative of Figure (A) Zirconat populations core sampl Dekese Formationthe in Lukuga (red), m 1680 1578 1568 (blue), m m 1260 (black), at populations (orange). m 1040 sampled (B) Zircon and (blue), (orange). depth m

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Figure 4.11 Plots of stable isotope composition and weight percent micritic calcite for samples from the Lukuga Formation in the Dekese core. (A) Measured δ18O values versus weight percent calcite. (B)Measured δ13C values versus weight percent calcite. (C) Measured δ13C versus δ18O values of calcite samples.

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Isotope geochemistry

Micritic calcite cement (Texture 2, n = 81) 13C values, range from -44.6‰ to -4.1‰, whereas 18O values range from –19.3‰ to 7.0‰. There is neither a clear correlation between weight percent calcite and measured 18O and 13C values (Figure 4.11a,b), nor is a good parametric correlation between measured 18O and 13C values (Figure 4.11c). This lack of correlation is considered to indicate that (1) there is no significant contribution of either detrital or diagenetic calcite that is isotopically different from the lacustrine calcite, (2) the calcite is predominantly authigenic and (3) the measured isotope values reflect Permo-Carboniferous paleoenvironmental conditions.

Discussion

The zircon population ranging from ~2.5 to 3.1 Ga likely corresponds to an early crust-forming event associated with the Namibian craton (van Schijndel et al., 2011), adjacent to the Congo craton, and represents uplift, erosion, transport, and depositional cycles from the hinterlands to the Congo Basin during Carboniferous-Permian time. Zircon populations of ~1.9 to 2.2 Ga are thought to reflect crustal formation in the West African craton, (Egal et al., 2002), the so-called

Eburan orogeny, and effectively continuous transport into the Congo Basin during

Carboniferous-Permian. Zircon populations with age estimates of ~1.0 to 1.4 Ga correlate with crust-forming events associated the Kibaran orogeny along the northern boundary of the Congo

Basin (DeWaele et al., 2003). The detrital zircon population with an approximate age of ~0.85

Ga likely corresponds to volcanism and episodic plume events associated with the breakup of

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Rodinia along the southern margin of the Congo Basin (Bogdanova et al., 2009), whereas the abundant population of 0.5 to 0.7 Ga detrital zircons are related to sediments derived from weathering of rocks uplifted by the Pan-African (~Damaran) orogenic event (Grantham et al.,

2003). Finally, the very small population of detrital zircons with age of ~306 to 375 Ma in the two youngest (and stratigraphically shallowest) samples from the Lukuga Formation may be related to crustal formation associated with the assembly of Gondwana and Pangea. These detrital zircon age-date estimates confirm that the Lukuga Formation in the Dekese core is dominated by detrital siliciclastics, which were transported into the basin, and that there was little, if any, direct volcanic input. These results suggest that Permian explosive volcanism played a minor role in the sedimentary fill of the Congo Basin and did not affect the clay mineralogy of the deposits, as it did in other Carboniferous-Permian basins (e.g., Bangert et al.,

1999; Fildani et al., 2009; Wilson and Guiraud, 1998). Therefore the clay mineralogy reported here of the Lukuga Formation likely represents the products of weathering within the Congo

Basin watershed prior to deposition in the lacustrine environment.

The clay-size fraction in siliciclastic lacustrine strata is primarily represented by a mixture of detrital minerals that is washed into the lake system from soils, regolith, and country rock within the watershed (Grim, 1958; Picard and High; 1972; Kelts, 1988; Yemane et al.,

1996). The mineralogy of the detrital clay fraction can be strongly correlated with regional climate over the watershed (Keller, 1970; Srodon et al., 1999; Wilson, 1999). This reflects that the clay-size fraction may be generated either by (1) physical breakdown to a fine grain size or

(2) chemical weathering that is facilitated by availability of aqueous solution (i.e., rainfall and groundwater) and energy (heat) that results in either transformation or neoformation of a fine

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size fraction (Rasmussen and Tabor, 2007). Cold (or hot) and dry climates are dominated by physical weathering, and therefore the clay-size fraction is approximately equivalent to that of the surrounding country rock. As a result, cold and dry climates are typically dominated by chlorite, illite (or mica-like minerals), and quartz (Keller 1970; Wilson, 1999; Srodon, 1999). In environments characterized by greater amounts of moisture and higher temperatures, there is increasing hydrolytic reaction of the country rock, which results in progressive loss of base cations, de-silication, and neoformation of clay-sized minerals. As a result, weathering in climates with seasonal fluctuations of temperature and precipitation, such as the modern mid- latitudes, results in the formation of expansible and mixed-layer expansible 2:1 phyllosilicate clay minerals, whereas weathering in humid and warm climates, such as the modern tropics and sub-tropics, results in the formation of 1:1 phyllosilicates such as kaolinite or halloysite

(Yemane, et al., 2006).

All samples analyzed here contain a significant fraction of quartz and feldspar in bulk powders and clay-size fractions. These minerals are all also observed macroscopically from dropstones and out-sized clasts within the Lukuga Formation (e.g., Cahen et al., 1960), and therefore are considered to be detrital. The dominance of Fe-rich chlorite and illite throughout the clay-size fraction is thought to be derived from a detrital, chemically unweathered sediment, which was physically weathered in the paleo-watershed, then transported into the lacustrine environment.

The composition of the clay-size fraction may be used to delineate three Clay

Mineralogical Zones (CMZs) within the Lukuga Formation in the Dekese core (Figure 4.2). The lower CMZ 1, which ranges from 1600m to 1442 m below the top of the core, consists of 13

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horizons characterized by the presence of chlorite, illite, and poorly sorted 2:1 expansible phyllosilicates. CMZ 1 also includes four different occurrences of authigenic calcite near the top of the sequence. CMZ 2, which extends from 1442 to 961 m, is composed primarily of chlorite and illite. CMZ 2 also preserves 17 different occurrences of finely disseminated authigenic calcite. CMZ 3, extending from 961 to 820.45 m, is comprised of nine horizons containing chlorite, illite, and poorly sorted 2:1 expansible phyllosilicates, as well as one horizon with kaolinite at 961 m. Zone 3 also includes 16 different occurrences of authigenic calcite that is concentrated between 924.5 and 837.9 m.

Interpretation

Lithostratigraphy

The distribution of outsized clasts, interpreted to here as an indicator of ice-rafted transport of coarse materials, provides lithostratigraphic evidence of cold conditions during deposition of the lower Lukuga Formation. The presence of this facies suggests conditions cold enough to freeze solid at the surface of the lacustrine environment and permit transport of detrital clastics far into the center of the basin via ice-rafting or fluvial transport, and subsequent gravity fall during spring thaw. The concentration of outsized clasts in three zones at 1573-1539 m,

1511-1503 m, and 1429-1276 m (Figure 4.6) suggests that at least three episodes of relatively cold climate during deposition of the Lukuga Formation, with the coldest conditions in the lower

2/5 of the formation.

130

Coupled laminations of fine-grained clay-sized, organic rich sediment, and coarser silt and very fine sand-sized sediments are interpreted as glacial varves. These sedimentary structures represent seasonal alteration between cold winters during which lake surfaces were covered in ice and clay- and organic-rich materials settled to the lake bottom and warmer seasons when high fluvial input delivered coarser materials to the catchment. This facies, which is indicative of traction transport, likely represents distal deltaic environments and so also may reflect the proximity of clastic sedimentary inputs to the lake basin. Varved deposits represent seasonal freeze-thaw cycles and imply somewhat warmer conditions than the horizons with outsized clasts. Varves occur in two broad stratigraphic zones in the Dekese core at 1525-1426 m and 1270-924 m (Figure 4.6), that are intercalated with, and extend above, the stratigraphic zones containing outsized clasts. This implies an episodic climate oscillation between relatively cold paleoclimates in the lower part of the Lukuga Formation and slightly warmer conditions, permissive of seasonal freeze-thaw in the middle of the formation. The lack of outsized clasts and varved deposits in the upper Lukuga Formation, between 945 and 800 m depth, suggests even warmer conditions that were not conducive to development of these sedimentary structures.

Clay mineralogy

The relative abundances of chlorite and illite vary substantially in CMZ1 and CMZ3, but are consistent in CMZ2 (Figure 4.2). The stratigraphic layers in CMZ1 and CMZ3 that include substantial decreases in chlorite content are the same layers in which expansible 2:1 phyllosilicate minerals are also present. Specifically, CMZ1 includes 13 stratigraphic units with expansible phyllosilicates, and CMZ3 includes nine horizons with 2:1 phyllosilicates. Based on 131

the environmental occurrences of these minerals, CMZ1 and CMZ3 record episodic environmental shifts between (1) cold climates where only detrital chlorite and illite are deposited and (2) relatively warmer, but still cool, environments in which some fraction of chemically weathered detrital 2:1 expansible phyllosilicates are also deposited with chlorite and illite. Kaolinite present near the base of CMZ3 may represent some of the warmest conditions during deposition of Permo-Carboniferous strata in the Congo Basin. The lack of expansible 2:1 phyllosilicates and kaolinite in CMZ2 reflects a long period of relatively cold conditions during which chemical weathering within the Congo Basin watershed was minimal.

Calcite stable isotope composition

Sparry calcite cement forming veins within the Lukuga Formation likely reflects crystallization in deeper burial environments influenced by geopressured fluids. These crystallization conditions may not reflect environmental conditions near the surface, and therefore are not considered further in this work. In contrast, the micritic calcite cements that are both occlusive and displacive probably formed during, or soon after burial, and may reflect near- surface environmental conditions. Therefore, the stable isotope geochemistry of these cement textures will be the focus of the discussion presented.

Stable carbon isotope values of the micritic calcite cements record extremely large variations, ranging from –44 to –4‰. This range of carbon isotope values likely reflects bicarbonate originating from several different sources, including (1) CO2 in the atmosphere, (2)

CO2 derived from oxidation of organic matter, and (3) CO2 derived from oxidation of CH4

132

generated by methanogenic bacteria (e.g., Hoefs, 1997). These different sources of CO2 are not necessarily related to regional environmental or climatic conditions at the time of crystallization and are not relevant to this study.

The oxygen isotopic compositions of micritic calcite cements in the Lukuga Formation are assumed to represent chemical equilibrium with meteoric waters at the time of mineral crystallization. There is a correlation between meteoric precipitation and mean annual surface temperature recorded by International Atomic Energy Association (IAEA) stations at localities where mean annual temperatures are less than 15 °C; however, there are different correlations associated with 18O-enriched coastal maritime precipitation and 18O-poor continental interior precipitation at isothermal sites (Ferguson et al., 1999). Localities within the IAEA dataset may be segregated into two groups with unique mathematical relationships between precipitation δ18O and temperature: (1) maritime sites <200 km from the coast and (2) continental sites >200 km inland. The Congo Basin contained a permanent, large lake for many millions of years (e.g.,

Cahen et al., 1960), which developed deep in the interior of Gondawana, several thousand kilometers from the nearest marine basins in the Permo-Carboniferous (Blakey, 2007).

Therefore, the mathematical correlation presented by Ferguson et al. (1999) for continental environments will be used to infer surface air temperatures from the δ18O values of micritic calcite cements in the Lukuga Formation:

18 δ OH2O(SMOW) = [MAAT(°C) -18.1]/1.05

This equation, which includes an analytical uncertainty of +/- 2 °C, may be combined with the oxygen isotope fractionation equation between calcite and water:

3 18 10 ln αcalcite-H2O(SMOW) = (18.03 * 1000)/T(°K) – 32.42

133

(Kim and O’Neill, 1997) and the equation used to convert oxygen isotope values from PDB to

SMOW standard:

18 18 δ OSMOW = 30.92 + 1.03092 * δ OPDB to plot the isotherms in Figure 4.6. If these isotherms, representing oxygen isotopic compositions of calcite formed in equilibrium with meteoric waters, are reasonable representations of Late

Paleozoic environments, the stratigraphic pattern of micritic calcite δ18O values can be used to infer crystallization temperatures. The δ18O results suggest substantial variation in surface temperature during deposition of the Lukuga Formation, including six stratigraphic intervals associated with cooling events. Of these six cooling events, designated isotope cold phases

(ICPs), five are characterized by near-freezing temperatures. These ICPs are intercalated with strata that yield micritic calcite δ18O values consistent with warmer temperatures well above freezing. These oxygen isotope data indicate cold conditions for the interval at 1668 – 930 m.

Above this level, the oxygen isotope data indicate a rapid, significant warming trend in the upper

Lukuga Formation, with the exception of a stratigraphically short (and presumably brief) episode of cooling between 825 m and 800 m.

Conclusions

The Carboniferous-Permian stratigraphy of the Dekese core in the central Congo Basin preserves evidence of environmental and climatic change at several different scales, depending on which proxies are considered. The stratigraphic distribution of deposits with outsized clasts delineates a broad, long-term climate shift from relatively cold to warm conditions through the entire Permo-Carboniferous succession. Detrital zircon geochronology reveals several separate age populations ranging from ~2.5 to 3.1 Ga, from ~1.9 to 2.2 Ga, from ~1.0 to 1.4 Ga, and from 134

~0.5 to 0.7 Ga, with a small population of ~350 Ma zircons in the youngest samples from the

Lukuga Formation. These results indicate that (1) detrital siliciclastic sediments dominate the stratigraphy in this part of the Congo Basin, (2) volcanic input was minimal, and (3) the mineralogy of the clay-size fraction reflects weathering conditions of detrital minerals within the

Congo Basin watershed. Clay mineralogy of bulk powders and clay-size fractions divide the

Lukuga Formation into three Clay Mineral Zones (CMZs). CMZ1 and CMZ3 contain clay minerals indicative of limited chemical weathering, whereas CMZ2 contains clay minerals that are more indicative of virtually no chemical weathering. Finally, δ18O values of early burial lacustrine calcite cements delineate at least three, and possibly as many as six, climate cycles oscillating between cold and relatively warm conditions. Clay mineralogy and calcite δ18O values from the Upper Paleozoic portion of the Dekese core, considered in conjunction with data from other contemporaneous Gondwanan basins, provide a basis to evaluate the evolution of

Permo-Carboniferous paleoclimate, and its relation to glacial ice dynamics. This reconstruction of Permo-Carboniferous terrestrial paleoclimate in the Congo Basin is similar to paleoclimate records inferred from mixed marine successions in the Karoo and Kalahari basins (Scheffler et al., 2006), and strongly supports the notion that Gondwanan climate was characterized by repeated oscillations between glacial and interglacial conditions (e.g., Fielding et al., 2008;

Montañez et al., 2007). Furthermore, the magnitude of climate shifts inferred from this study may provide a useful means of correlation to strata from contemporaneous marine basins, to other near-field terrestrial basins in Gondwana, and to far-field para-tropical terrestrial basins in

Euramerica. However, a lack of chronostratigraphic constraints with the Carboniferous-Permian strata of the Congo Basin does not permit direct comparisons with results from other studies which document episodic glacial phases in the Karoo basin in South Africa (Scheffler et al., 135

2006), Australia (Jones and Fielding, 2008), Antarctica (Isbell et al., 2003), and South America

(Gulbranson et al., 2008, 2010). Nevertheless, this study provides a paleoenvironmental context for the Carboniferous-Permian strata of the Lukuga Formation of the Congo basin when a chronostratigraphy is finally determined for these strata.

136

Chapter 5

SUMMARY

The work presented in this dissertation highlights the importance and robustness of geochemical proxies as indicators of past environmental and paleoclimate conditions preserved in terrestrial records throughout geologic time and space. The isotopic record of organic carbon preserved in modern lake sediments tracks historical records of environmental change through a period of well-documented changes in land usage and elucidates relationships between preserved organic carbon isotopes and contributions to lacustrine systems from a changing watershed environment. In Chapter 3, organic carbon isotopes preserved in paleosol carbonates, paleosol clay mineralogical composition, and the morphology of paleosols themselves are utilized to elucidate long-term paleoenvironmental perturbations over timescales of interest in terms of

Geologic time. The work put forth in Chapter 4 further emphasizes the utility of clay mineralogy and stable isotope geochemistry of primary carbonate textures as an invaluable tool in reconstructing environmental changes in Earth’s deep past. Further work is needed to refine the timing of the paleoenvironmental signature preserved in the Lukaga Formation deposits; however, they still hold highly pertinent information about deglaciation on a vegetated landscape. The coupled record of environmental change at the paleo-subtropics and environmental shifts observed from the paleo-polar regions provide insight into the timing and

137

mechanics of continental deglaciation which can aid in understanding and predicting modern global climate pattern changes.

138

APPENDIX WRL-1

Table of stable carbon and oxygen isotope data of carbonate samples from the White Rock Lake sediment core.

Weight % 13 18 Sample Depth Size Fraction CaCO3 δ C δ O WR3 0-3 -1.5 bulk 21.0 -0.29 -3.54 WR3 12-15 -13.5 bulk 35.7 -0.24 -3.86 WR3 24-27 -25.5 bulk 36.2 -0.10 -3.91 WR3 24-27 -25.5 bulk 34.9 -0.09 -3.92 WR3 40-45 -42.5 bulk 32.6 -0.22 -3.86 WR3 60-65 -62.5 bulk 31.4 -0.35 -4.01 WR3 80-85 -82.5 bulk 24.2 -0.82 -4.54 WR3 85-90 -87.5 bulk 30.7 -0.21 -3.14 WR3 100-105 -102.5 bulk 23.8 -0.26 -3.57 WR3 105-110 -112.5 bulk 28.6 -0.25 -3.29 WR3 120-125 -122.5 bulk 23.5 -0.29 -3.41 WR3 120-125 -122.5 bulk 23.6 -0.28 -3.34 WR3 125-130 -127.5 bulk 23.8 -0.43 -3.32 WR3 140-145 -142.5 bulk 19.2 -1.41 -4.36 WR3 145-150 -147.5 bulk 20.1 -0.45 -3.06 WR3 170-180 -175 bulk 13.7 -0.50 -3.63 WR3 180-190 -185 bulk 15.3 -0.41 -2.87 WR3 180-190 -185 bulk 14.9 -0.42 -3.99 WR3 12-15 -13.5 coarse 23.0 -0.24 -2.66 WR 27-30 -28.5 coarse 18.4 -0.27 -1.86 WR3 40-50 -42.5 coarse 22.5 -0.28 -2.38 WR3 80-85 -82.5 coarse 12.8 -0.23 -1.12 WR 125-130 -127.5 coarse 12.1 -0.51 -2.73 WR3 150-160 -155 coarse 5.1 -0.10 -0.89 WR3 170-180 -175 coarse 5.2 -0.40 -3.09 WR3 210-220 -215 coarse 4.6 -0.24 -1.39 WR3 12-15 -13.5 fine 41.6 -0.24 -3.75 WR3 12-15 -13.5 fine 40.1 -0.20 -3.56 WR 27-30 -28.5 fine 40.6 -0.22 -3.93 WR 27-30 -28.5 fine 39.8 -0.23 -3.70 WR 45-50 -47.5 fine 37.0 -0.23 -3.85 WR3 80-85 -82.5 fine 34.7 -0.18 -3.64 WR3 80-85 -82.5 fine 35.4 -0.16 -3.41 WR 125-130 -127.5 fine 24.8 -0.47 -3.18 139

Table A 1 continued

Weight % 13 18 Sample Depth Size Fraction CaCO3 δ C δ O WR 125-130 -127.5 fine 25.6 -0.47 -3.24 WR3 150-160 -155 fine 24.7 -0.41 -3.35 WR3 150-160 -155 fine 25.5 -0.40 -2.83 WR3 170-180 -175 fine 24.9 -0.54 -2.82 WR3 170-180 -175 fine 25.5 -0.50 -2.86 WR3 210-220 -215 fine 25.8 -0.51 -2.83

140

APPENDIX MORA-1

X-ray diffraction spectra of oriented aggregate mounts of the < 2 µm size fraction of paleosol matrix samples from Mora County, New Mexico, as discussed in chapter 3.

K500 Mora1 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

141

K500 Mora23 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora36 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

142

Mora39K K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora41 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

143

K500 Mora48 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora49 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

144

Mora50K K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora51 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

145

K500 Mora52 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora53 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

146

K500 Mora54 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora55 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ 147

K500 Mora63 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

Mora70 K

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ 148

Mora75 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

Mora78 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ 149

Mora82 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora85 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

150

K500 Mora92 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora95 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

151

K500 Mora98 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora109 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

152

Mora110 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora126 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

153

K500 Mora130 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 Mora133 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

154

APPENDIX MORA-2

X-ray diffraction spectra of paleosol thin sections from Mora County, New Mexico. Each thin section was x-rayed with the following parameters: scans from 20° to 55° 2θ, with a scan speed of 3° 2θ/minute.

Mora39_TS Relative Intensity

20 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50 52 54 °2Θ

155

Mora102_TS Relative Intensity

20 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50 52 54 °2Θ

Mora117_TS Relative Intensity

20 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50 52 54 °2Θ

156

Mora121_TS Relative Intensity

20 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50 52 54 56 58 60 °2Θ

157

Mora122_TS Relative Intensity

20 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50 52 54 °2Θ

Mora123_TS Relative Imtensity

20 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50 52 54 °2Θ

158

Mora128_TS Relative Intensity

20 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50 52 54 °2Θ

Mora131_TS Relative Intensity

20 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50 52 54 °2Θ

159

Mora134_TS Relative Intensity

20 22 24 26 28 30 32 34 36 38 40 42 44 46 48 50 52 54 °2Θ

160

APPENDIX MORA-3

Petrographic observations of carbonate textures sampled from the Madera and Sangre de Cristo formations of Rowe-Mora basin in Mora County, New Mexico.

Sample Strat. Primary grains Petrographic observations ID Height (m) Mora17 138.3 calcite black micritic calcite, few thin (up to 0.05mm diameter) microspar veins Mora32 416.74 clay, siliclastics sub-angular siliclastic grains in clay matrix, common angular quartz, and micritic calcite Mora35 510.79 micrite clotted micritic calcite, with rare angular quartz grains

Mora 37 517.99 micrite, qtz, lithics micritic calcite matrix with common black-dark brown lithic fragments, common spicules, up to 0.375mm in length Mora38 519.09 micrite, clay, qtz micritic calcite matrix, clay coatings along fractures, microspar filled veins, tubules up to 2mm in length, .25mm wide are filled with soil matrix material Mora39 519.29 clay, qtz clay matrix with rare angular quartz grains and spicules

Mora40 519.44 micrite, clay, discrete regions of clotted micrite, bounded by clay minerals oxides and iron staining Mora42 583.34 micrite, microspar, clotted micritic calcite, with microspar filled septarian clay fractures, clay Mora43 583.54 micrite, spar clotted micrite blebs, bound by sparry calcite veins that comprise up to 75% of sample Mora44 583.99 micrite, microspar, clotted micritic calcite, with microspar filled septarian clay fractures, clay Mora45 586.54 micrite, quartz clotted micrite with common (~5%) angular to sub-angular quartz grains Mora46 587.44 micrite, microspar, clotted micritic calcite, with microspar filled septarian clay fractures, clay Mora47 587.69 micrite, clay, clotted micrite with microsparry calcite that comprises up to microspar 30% of sample. Veins measure up to 2 mm wide and crosscut entire sample. Cross section of micrite filled tubule measures 2 mm diameter.

161

Table of petrographic observations of thin sections from the Madera and Sangre de Cristo formations of the Rowe-Mora basin, continued.

Sample ID Strat. Primary grains Petrographic observations Height (m) Mora61 715.56 micrite, clay, regions of clotted micrite with common irregularly shaped oxides voids bound by calcitic prisms. Voids may be infilled with iron-oxides (dark, opaque mineral) Mora62 716.3 micrite, microspar, clotted micritic calcite, with microspar filled septarian clay fractures, clay Mora65 717.07 clotted micritic calcite, with microspar filled septarian fractures, clay and diffuse regions of iron oxides Mora70 750.28 dol, micrite, dolomite crystals up to 0.1 mm bound by diffuse regions of oxides iron oxides, with clay-rich regions Mora71 750.65 dol, micrite, dolomite crystals up to 0.1 mm bound by diffuse regions of oxides iron oxides, with clay-rich regions Mora74 913.69 qtz, clay, dol granular accumulations of dolomite or calcite grains, ranging in size from 0.1 mm to 0.05 mm, iron staining on grain boundaries. Crystals (up to 0.5 mm diameter) display excellent rhombohedral cleavage Mora75 916.49 qtz, clay, granular accumulations of dolomite crystals, ranging in size micrite/dol from 0.1 mm to 0.05 mm, iron staining on grain boundaries. Crystals (up to 0.5 mm diameter) display fair to excellent rhombohedral cleavage. Minor amounts of calcite are present lining fractures Mora76 1009.34 micrite, oxides granular micritic calcite grains, ranging in size from 0.1 mm to 0.05 mm, iron staining on grain boundaries Mora77 1061.39 micrite, clay, discrete regions of clotted micrite, bounded by clay minerals oxides and iron staining Mora79 1102.24 calcite, qtz bladed, swallow tail psuedomorph of calcite after gypsum, few silt-size quartz grains Mora80 1132.64 micrite, clay, discrete regions of clotted micrite, bounded by clay minerals oxides and iron staining Mora81 1132.64 silica silicified plant material. Individual, uniformly sized plant cells oriented in regular repeating arrangement, fractures filled by micritic calcite Mora83 1132.64 micrite massive micritic calcite Mora84 1172.84 micrite, microspar, discrete regions of clotted micrite, bounded by clay minerals clay and iron staining, few microsparry calcite fractures Mora86 1188.79 micrite, microspar, discrete regions of clotted micrite, bounded by clay minerals clay and iron staining, microsparry calcite fractures (up to 50% of sample area) Mora88 1189.14 micrite, microspar, discrete regions of clotted micrite, bounded by clay minerals clay and iron staining, microsparry calcite fractures (up to 35% of sample area)

162

Table of petrographic observations of thin sections from the Madera and Sangre de Cristo formations of the Rowe-Mora basin, continued.

Sample ID Strat. Primary grains Petrographic observations Height (m) Mora96 1202.16 micrite, microspar, regions of clotted micrite with common irregularly shaped oxides voids bound by calcitic prisms. Voids may be infilled with iron-oxides (dark, opaque mineral) Mora97 1202.75 micrite, spar, clay, discrete regions of clotted micrite, bound iron staining, sparry oxides calcite fractures (up to 35% of sample area). Several occurrences of clay material organized into tubules up to 1.5 mm in length, 0.5 mm width

Mora99 1226.8 micrite massive clotted micrite Mora102 1236.35 micrite, microspar, massive micritic calcite, clay and iron-oxide lined fractures, clay occluded organic matter, microspar calcite fracture fill, few medium sand-size quartz grains Mora103 1247.5 micrite, microspar, massive micritic calcite, clay and iron-oxide lined fractures, clay occluded organic matter, microspar calcite fracture fill Mora105 1248.25 micrite, clay, clotted micritic calcite, with microspar filled septarian oxides fractures, clay and diffuse regions of iron oxides Mora106 1253.05 micrite, qtz clotted micrite with common (~5%) angular to sub-angular quartz grains Mora107 1253.6 micrite, microspar, massive micritic calcite, clay and iron-oxide lined fractures, clay occluded organic matter, microspar calcite fracture fill Mora108 1253.75 micrite, microspar, Clotted micrite with common irregularly shaped voids bound clay by calcitic prisms Mora110 1288.15 micrite, clay, massive, clotted micrite with few microspar filled fractures, oxides opaque fracture extends 2.0 cm in length, up to 1.0 mm width Mora111 1323.45 micrite, oxides massive micritic calcite Mora112 1330.05 micrite massive clotted micrite Mora117 1446.15 micrite, oxides, micrite granules cemented by calcite, iron staining present at gyp grain boundaries, all grains are <0.1 mm diameter, and very uniform. May possibly be gypsum. Mora118 1447.8 qtz, calcite massive amalgamation of very fine sand- to silt-size grains of quartz cemented by calcite Mora119 1451.57 gyp, calcite, dol gypsum and clay masses crosscut by microsparry dolomite and calcite veins

163

Table of petrographic observations of thin sections from the Madera and Sangre de Cristo formations of the Rowe-Mora basin, continued.

Sample ID Strat. Primary grains Petrographic observations Height (m) Mora120 1451.74 gyp, calcite, dol dolomitic veins up to 1.0mm width with sparry calcite infilling, gypsiferous matrix, alteration of dolomite to calcite is observed on interior of fractures Mora121 1451.89 gyp, calcite, dol dolomitic veins up to 1.0mm width with sparry calcite infilling, gypsiferous matrix, alteration of dolomite to calcite is observed on interior of fractures Mora122 1454.39 gyp, calcite, qtz gypsum and clay masses crosscut by microsparry calcite veins Mora123 1455.39 gyp, calcite, qtz gypsum and clay masses crosscut by microsparry calcite veins Mora125 1457.5 micrite, lithics micritic calcite matrix with common (~25%) lithic grains ranging from 0.1 to 1.2 mm diameter Mora127 1458.7 clay, calcite, gyp gypsum and clay masses crosscut by microsparry calcite veins Mora128 1459.7 micrite, qtz, lithics calcite cemented very fine quartz sand and lithic grains up to 0.5 mm diameter, clay blebs Mora129 1462.1 micrite, qtz, lithics clotted micrite, with common siliclastics grains and subangular fine quartz Mora131 1462.1 qtz, clay very fine silt and clay with common subangular very fine quartz grains Mora132 1470.15 dol, spar, gyp dolomitic veins up to 1.0mm width with sparry calcite infilling, gypsiferous Mora133 1472.8 qtz, clay very fine silt and mudstone with common (~15%) subangular quartz grains Mora134 1455.76 micrite, qtz, clay clotted micrite with common (~30%) medium to fine sand- size grains and clay coating along fractures

164

APPENDIX DEKESE-1

X-ray diffraction spectra from oriented aggregates of clay size (<2 µm e.s.d.) fraction of samples from the Lakuga Formation of the Dekese core discussed in Chapter 4. X-ray diffraction scans were conducted with at 2° 2θ/minute from 2 to 30° 2θ.

D874 K500 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

165

K500 D911.5 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

D911.7 K500 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

166

K500 D912 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D916 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

167

D961 K500 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1058.8 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

168

K500 D1107.25 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1118.3 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

169

K500 D1180.4 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1184.3 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

170

K500 D1190 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1190_2 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

171

D1202 K500

K Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1227 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

172

K500 D1243 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

D1275.75 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

173

K500 D1280 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

D1283.8 K

Mg Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

174

K500 D1286.7 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1293 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

175

K500 D1294.25 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

D1305.9 K500 K

Mg Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

176

K500 D1310 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1313.6 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

177

K500 D1322 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1329.85 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

178

K500 D1334.25 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1341 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

179

K500 D1343.25 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1348.3 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

180

K500 D1356.5 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1361 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

181

K500 D1365.5 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1365.5 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

182

D1385 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

D1386 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

183

D1394 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

D1432 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

184

D1437 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

D1442 K500 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

185

D1442_2 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

D1458.25 K500 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

186

K500 D1471.5 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

K500 D1501 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ 187

D1531 K500 K Mg

Mg + Gly Relative Intensity

2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 °2Θ

188

APPENDIX D-2

X-Ray diffraction spectra of bulk powder samples from the Lukuga Formation in the Dekese Core. Each analysis was carried out with the following parameters:

D714.8 Relative Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

189

D721.5 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D753.04 Relative Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

190

D782.25 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D801.9 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

191

D820.45 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D874 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

192

D901.9 Relative Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D911.7 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

193

D954 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D1034 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

194

D1048 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D1058.8 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

195

D1129.5 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D1185.5 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

196

D1190 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D1270.3 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

197

D1293 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D1365.5 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

198

D1379.8 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D1405.5 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

199

D1437 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D1458.25 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

200

D1479.8 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D1503 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

201

D1524.5 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D1538.5 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

202

D1539.8 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

D1576 Relative Intensity

2 8 14 20 26 32 38 44 50 56 62 68 °2θ

203

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