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AND FLUVIAL SEDIMENTOLOGY OF THE UPPER CATSKILL FORMATION, CENTRAL : A TEST OF THE DISTRIBUTIVE FLUVIAL SYSTEM MODEL

A Thesis Submitted to the Temple University Graduate Board

In Partial Fulfillment of the Requirements for the Degree MASTER OF SCIENCE

By Christopher Oest July 2015

Advisors

Dr. Dennis O. Terry, Jr., Department of Earth and Environmental Science

Dr. Alexandra E. Krull-Davatzes, Department of Earth and Environmental Science

Dr. David E. Grandstaff, Department of Earth and Environmental Science

ABSTRACT ……

The Upper Devonian Catskill Formation represents marginal marine and alluvial which prograded into the Appalachian Basin during the .

Distributive fluvial systems (DFS) are prevalent in modern actively aggrading basins in all tectonic and climatic regimes and may be common in the rock record. In this study, I reinterpret the Catskill Formation as a prograding distributive fluvial system (DFS) on the basis of up-section variability in , channel textural trends, and alluvial architecture. At least three distinct pedotypes representative of prevailing forming conditions are identified during deposition of the Irish Valley, Sherman Creek, and

Duncannon Members of the Catskill Formation. Increased is inferred from an up-section transition from hydromorphic aqualfs within the Irish Valley Member to non-calcareous, uderts within the Duncannon Member. Qualitative field observations of channel sandstone morphology show an increase in channel size up-section. Channels occur as small isolated bodies at the base of the section, transitioning to relatively larger, amalgamated channels, and finally, large isolated channel bodies up-section. are litharenites and coarsen-upward throughout the Catskill Formation overall. This coarsening upward trend results from increasing paleo-flow competency in larger channels up-section. These results are consistent with deposition of the Catskill

Formation by DFS processes and demonstrate the utility of paleopedological analysis in interpreting alluvial depositional processes. Identifying DFS in the rock record has implications for paleosol-based paleoclimatic studies, as paleosols forming on prograding

DFS have increased paleosol drainage up-section, which could potentially be misinterpreted as a shift from prevailing humid to arid paleoclimatic conditions.

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Recognition of DFS in the rock record also has implications for basin analysis and exploration of fluvial aquifers and hydrocarbon reservoirs, as the stratigraphic architecture of DFS are fundamentally different from tributary systems at the basin scale.

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ACKNOWLEDGEMENTS

The majority of the funding for this project was provided by the Society of

Sedimentary (SEPM) Foundation Student Assistance Grant. Temple University

Department of Earth and Environmental Science provided laboratory supplies and travel support.

DT – You’re responsible for getting me into this business. Thank you for giving me ample freedom to allow me to dive into a project that I was genuinely interested in.

Alix – I can’t recall how many times I’ve come into your office in a mild panic asking “if you had a minute” and you still took the time to talk with me even though you were clearly in the middle of something. Thank you for your guidance. G – I know every time

I bring a manuscript to you, I’ll get it back torn to shreds. But it always turns out for the best. Thank you for constantly keeping me on my toes. Ilya – You’re name might not be on the cover page, but I still consider you a mentor. The ideas that you planted in my head were great contributions to this project. Nick – Whether you realize it or not (or believe me), I genuinely loved both of your classes. Thank you for taking the time to entertain my random questions about all things structural geology and making me a better scientist.

Many, many thanks to my fellow graduate students, especially Haley and Paul.

You were always there to unwind at the end of a long day/week. May there be many dollar taco nights ahead for us.

Steve, Jesse, and Tim – You all have wisdom beyond your years in geology and in life. Thanks for imparting that onto me and keeping me cool on rough days.

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The Valentine Lab - Thanks for the fun times at lunch and helping me remember basic chemistry.

To my undergraduate field assistants (donkeys/wolves) – Phil, Wes, Zach, Chris – the field work for this project literally could not have been completed in time without your help. I am in your debt. Aaron, Chris, Chris, and Jess – Your help with preparing my thin sections was huge.

To all my students – Thank you for not accepting everything I said at face value.

You made me a better geologist. And remember, we like to have fun here at Temple

Geology.

Thanks to my family (Mom, Dad, Patrick, Gran) for you overwhelming support

(both emotional and financial). You got me through this.

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TABLE OF CONTENTS

ABSTRACT …… ...... I

ACKNOWLEDGEMENTS ...... III

LIST OF FIGURES ...... VII

LIST OF TABLES ...... IX

CHAPTER 1. INTRODUCTION ...... 1

1.2 Distributive Fluvial Systems ...... 7

1.3 Stratigraphy, Paleogeography, Tectonic Setting, and Paleoclimate ...... 8

1.4 Methods ...... 17

CHAPTER 2. RESULTS ...... 27

2.1 Paleopedology ...... 27

2.1.1 Pedotype 1 ...... 27

2.1.2 Pedotype 2 ...... 35

2.1.3 Pedotype 3 ...... 38

2.2 Channel Sandstone Petrography ...... 41

2.2.1 Mineralogy ...... 41

2.2.2 Grain Size Analysis ...... 47

2.3 Channel Morphology ...... 52

2.4 Decompaction...... 52

CHAPTER 3. DISCUSSION ...... 54

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3.1 Paleoenvironmental Factors of Soil Formation ...... 54

3.1.1 Paleoclimate ...... 54

3.1.2 Effect of Land Plants ...... 57

3.1.3 Gradient and Base Level ...... 58

3.1.4 ...... 61

3.1.5 Time of ...... 62

3.2 Sandstone Provenance ...... 64

3.3 Up-section Variability in Channel Grain Size and Morphology ...... 65

3.4 Flow Competency Analysis ...... 67

3.5 A Note on Diagenesis ...... 70

CHAPTER 4. CONCLUSIONS ...... 72

REFERENCES CITED ...... 76

APPENDIX A. STUDY SITE LOCATIONS AND DESCRIPTIONS ...... 85

APPENDIX B. GRAIN SIZE DISTRIBUTIONS ...... 86

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LIST OF FIGURES

Figure 1: Hydrofacies variability on DFS...... 4

Figure 2: Study sites relative to the Devonian outcrop belt in Pennsylvania...... 6

Figure 3: Paleogeography and hypothetical cross section of the Appalachian

Basin during the Late Devonian ...... 12

Figure 4: Generalized stratigraphic section of the Catskill Formation in

Central Pennsylvania...... 13

Figure 5: Key to symbols used in Figure 4, Figure 8, and Figure 9...... 14

Figure 6: Outcrop photographs showing various lithologic characteristics of

the Irish Valley and Sherman Creek Members...... 15

Figure 7: Outcrop photographs showing various lithologic characteristics of

the Duncannon and Clarks Ferry Members...... 16

Figure 8: Measured section of the Clarks Ferry and Duncannon Members at

Duncannon, PA showing lithology and location of described

paleosols and sandstone samples...... 23

Figure 9: Measured section of the Duncannon Member at Duncannon, PA,

continued...... 24

Figure 10: Lithologic sections of pedotypes illustrating macroscopic

pedogenic and sedimentary features...... 30

Figure 11: Outcrop photographs of pedotype paleosols...... 31

Figure 12: Photomicrographs of pedogenic features...... 32

Figure 13: X-ray diffractograms from representative samples for each

Pedotype...... 33

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Figure 14: Ternary diagram showing composition of -sized or larger

particles and inferred provenance...... 44

Figure 15: Photomicrographs of channel sandstones from the Catskill

Formation...... 46

Figure 16: Textural trends through the Catskill Formation...... 49

Figure 17: Block diagrams illustrating the relationship between base level

and progradation of DFS...... 60

Figure 18: Shear stress trends through the Catskill Formation...... 69

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LIST OF TABLES

Table 1: Characteristics of Distributive (DFS) and Tributary (TFS) Fluvial

Systems ...... 5

Table 2: Shear Stress Calculation Parameters ...... 25

Table 3: Decompaction Constants ...... 26

Table 4: Pedotype 1 Description ...... 29

Table 5: Pedotype Classifications ...... 34

Table 6: Pedotype 2 Description ...... 37

Table 7: Pedotype 3 Description ...... 40

Table 8: Raw Point Count Data ...... 45

Table 9: Grain Size Data ...... 50

Table 10: Characteristics of Distributive (DFS) versus Tributary (TFS)

Fluvial Systems in the Catskill Formation ...... 75

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CHAPTER 1. INTRODUCTION

1.1 Objectives and Significance

Interpretations of many fluvial deposits, including the Upper Devonian Catskill

Formation, are based on facies models derived from observations of modern tributary fluvial systems (Allen and Friend, 1968; Woodrow et al., 1973; Bridge, 1984; Sevon,

1985; Fielding et al., 2012; Trendell et al., 2013). However, fluvial deposition in modern, actively aggrading basins is dominated by distributive fluvial systems (DFS), and therefore these systems may be prevalent in the geologic record (Weissmann et al., 2010;

Weissmann et al., 2011; Weissmann et al., 2013; Hartley et al., 2010; Hartley et al.,

2013). See Figure 1. The objective of this study is to test the applicability of the DFS model to the geologic record. Table 1 outlines key differences between tributary and distributive fluvial systems.

The Catskill Formation has previously been interpreted as the result of deltaic processes (Barrell, 1913; Willard, 1939; Friedman and Johnson, 1966; Allen and Friend,

1968; Oliver et al., 1967; Glaeser, 1974; Ettensohn 1985; Faill, 1985, Faill, 1997, Faill

,2002; Sevon, 1985; Harper, 2002). In contrast, Walker (1971) and Walker and Harms

(1971) noted the absence of distributary mouth bar facies and repeated coarsening- upward cycles which are indicative of lobe switching on a deltaic complex. The absence of these characteristic deltaic facies implies deposition more consistent with alluvial aggradation rather than delta progradation (Brezinsky et al., 2009). Here, I present evidence from outcrop studies in central Pennsylvania (Figure 2) that deposition of the

Catskill Formation can be attributed to distributive fluvial system (DFS) processes based

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on up-section changes in paleosol development, channel sandstone morphology, grain size, mineralogy, and alluvial stacking patterns (Figure 1 A, B, Table 1).

Difficulty in mapping and correlation of the Catskill Formation within

Pennsylvania may partially reflect a lack of understanding of its lithofacies distributions, which is guided by an understanding of the depositional processes responsible for a particular system. Therefore, identifying depositional processes from their products in the rock record is fundamental in constructing accurate and applicable facies models for the analysis of sedimentary basins. Lithofacies distributions in fluvial environments are dependent on the type of system (e.g. tributary vs. distributary) (Weissmann et al., 2010;

Weissmann et al., 2011) and therefore recognizing DFS in the rock record has implications for predicting reservoir properties in fluvial hydrocarbon plays and understanding heterogeneities in alluvial aquifers (Nichols and Fisher, 2007; Fielding et al., 2012). Identifying prograding DFS in the rock record also has implications for paleosol-based paleoclimatic interpretations (Terry et al., 2012; Trendell et al., 2013;

Weissmann et al., 2013). Drying-upward trends in paleosols within the Chinle

Formation may have been misinterpreted as paleoclimatic change (Cleveland et al., 2008;

Dubiel and Hasiotis, 2011), when in reality this variability is simply the result of increasing paleosol drainage up-section as a function of a prograding DFS

(Figure 1 A, B) (Trendell et al., 2013). Potential examples of DFS in the rock record include the Miocene Luna System (Nichols, 1987; Nichols and Fisher, 2007), the

Eocene-Oligocene White River Group (Terry et al., 2012), the Upper -Early

Cretaceous Monteith Formation (Kukulski et al., 2013), the Upper Triassic Chinle

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Formation (Trendell et al., 2013), the Jurassic Morrison Formation (Weissmann et al.,

2013), and the Lower Beaufort Group (Wilson et al., 2014).

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Figure 1: Hydrofacies variability on DFS. A) Plan view of a hypothetical DFS. Note change in drainage conditions and channel morphology downstream. B) Cross sections through a hypothetical prograding DFS. A-A’ shows vertical trend in soil drainage and alluvial architecture. B-B’ shows the relationship between depth to the and soil drainage potential down depositional dip on a hypothetical DFS. The spring line is the point where the water table intersects the fan surface.

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Table 1: Characteristics of Distributive (DFS) and Tributary (TFS) Fluvial Systems

Characteristic DFS TFS

Well drained source proximal Soil Drainage* Variable, no predictable trend (Hartley et al., 2013) Soil maturity* Decreases downstream Increases downstream Perennially inundated source-distal Seasonally inundated source Floodplain inundation environments distal environments

Decreases downstream (Weissmann Channel size* Increases downstream et al., 2010)

Decreases downstream (Weissmann Channel Confinement* Increases downstream et al., 2010) Flow competency* Decreases downstream Increases downstream Source proximal: Large, multistory Source proximal: Small, channel complexes isolated channels

Alluvial Architecture* Source distal: Small, isolated Source distal: Large, possibly channels in overbank fines (Nichols amalgamated channel and Fisher, 2007; Hartley et al., complexes 2010; Trendell et al., 2013)

Source proximal: Large, coarse- grained channels Source proximal: Small,

coarse-grained channels Source distal: Primarily sheet Sandstone Geometry* deposition with small isolated Source distal: Large, fine channels grained channels (Nichols and Fisher, 2007; Hartley et al., 2010)

No predictable trend (multiple Sandstone Maturity* Simplistic fractionation sources)

10’s to 100’s of kilometers, fan morphology. Confined to river valleys. Areal Extent Not confined to river valleys. Dependent on size of (Weissmann et al., 2010; Davidson catchment basin. et al., 2013)

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Figure 2: Study sites relative to the Devonian outcrop belt in Pennsylvania. Modified from Miles and Whitfield (2001) and Harper (2002).

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1.2 Distributive Fluvial Systems

DFS are defined as having: 1) an apex from which active and abandoned channels diverge downstream, 2) a detectable topographic relief centered about the apex with slopes decreasing laterally and downstream, 3) rivers showing distributary characteristics

(i.e. channels bifurcating downstream), 4) tributary streams which do not join the DFS downstream, 5) a fan morphology in plan-view, and 6) decreasing channel size and confinement downstream (Figure 1 A) (Hartley et al., 2010; Weissmann et al., 2010;

Weissmann et al., 2011; Weissmann et al.,2013; Davidson, 2013).

Pedogenesis on DFS is largely controlled by hydrologic variability. Limited channel migration, greater channel confinement, and greater depth to the water table in source-proximal DFS environments promotes increased pedogenesis (Figure 1 B).

Conversely, source-distal DFS depositional environments are characterized by shallow water tables and generally unconfined flow (Hartley et al., 2013). These characteristics inhibit soil forming processes such as authigenesis and translocation, as water is not free to percolate from the surface through the underlying (Birkeland, 1999).

Subsurface accumulation of soluble bases (such as Ca, Mg, Na, and K) is also inhibited by these conditions. Therefore, in a prograding DFS paleosols will be increasingly well- drained and potentially well-developed up-section with fewer hydromorphic features, greater evidence of clay translocation, and potentially greater concentrations of subsurface accumulations (Figure 1 B).

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1.3 Stratigraphy, Paleogeography, Tectonic Setting, and Paleoclimate

The Acadian clastic wedge (historically referred to as the Catskill Delta, and more recently, the Catskill clastic wedge) is an approximately 3000 m thick wedge of siliciclastic sediments derived from the Acadian orogenic belt and deposited in alluvial plain, marginal marine, and marine depositional environments within the Appalachian foreland basin (Figure 3 A) (Sevon, 1985; Woodrow, 1985; Slingerland et al., 2009).

The Catskill Formation (Frasnian-) is an approximately 2000 m thick succession of mudstones, siltstones, and sandstones occupying the upper-most portion of the Acadian clastic wedge in Pennsylvania (Figure 3 B, Figure 4). The Catskill

Formation is thickest in central and eastern Pennsylvania and thins and interfingers with marine facies to the west (Figure 3 B) (Berg et al., 1993; Carter, 2007). Diachronous deposition occurred as a result of oblique convergence between promontories extending from the eastern margin of Laurentia and Avalonia, resulting in the Catskill Formation and its lateral equivalents, the Manorkill and Plattekill Formations, becoming younger from north to south in the Appalachian Basin (Faill, 1985; Faill, 1997; Faill,

2002; Ettensohn, 1985; Ettensohn, 2008; Ver Straeten, 2010). Subsidence due to crustal loading from thickening accretionary terranes and the assembling Acadian Orogenic belt, coupled with marine regression during the Late Devonian, created accommodation space in the Appalachian Basin and resulted in progradation of the clastic wedge to the present- day northwest (Ettensohn, 1985, Ettensohn, 2008; Faill, 1985; Haq and Schutter, 2008).

The Catskill Formation is subdivided into three principle lithostratigraphic members in central Pennsylvania (Figure 4). The Irish Valley Member is the lower-most member of the Catskill Formation and conformably overlies the marine Trimmers Rock

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Formation (Frasnian) (Berg et al., 1993). The Irish Valley Member consists of alternating red and green siltstones and predominantly gray, very fine grained sandstones arranged in poorly-defined fining-upward packages (Figure 6 A). This member is sparsely fossiliferous at its basal contact with the Trimmers Rock Formation (Figure 6 B). The

Sherman Creek Member conformably overlies the Irish Valley Member and is primarily composed of red siltstones and gray to reddish-gray very fine-grained sandstones arranged in poorly-defined fining-upward packages (Figure 6 C). Carbonized plant debris and calcareous paleosols are common in this member. Cotter and Driese (1998) document incised-valley fills within the Irish Valley and Sherman Creek Members

(Figure 6 D). The Duncannon Member is the upper-most member of the Catskill

Formation and consists of grayish red and reddish tan medium to coarse-grained crossbedded sandstones, red laminated siltstones, and red hackly mudstones arranged in well-defined fining-upward packages, often capped by thick, well-developed vertic paleosols (Figure 7 A, B). The Duncannon Member is unconformably overlain by the

Mississippian-Devonian Spechty Kopf Formation (Figure 4) (Epstein et al., 1974; Berg et al., 1993; Faill, 2002; Cecil et al., 2004).

The Clarks Ferry Member occurs between the Sherman Creek and Duncannon

Members and is present locally in central Pennsylvania (Figure 4). This unit is largely composed of medium to coarse-grained tan, pebbly, planar and trough crossbedded sandstones with minor conglomerates and red mudstones (Figure 7 C). Macerated carbonized plant debris is common and is often found in association with pyrite nodules.

Detrital charcoal fragments, identified following the criteria of Scott (2010), are common in the middle of this unit and range from millimeters to centimeters in diameter (Figure 7

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D). The Clarks Ferry Member pinches out to the northwest and grades into coarser- grained, source-proximal facies to the east (Glaeser, 1974; Berg et al., 1993).

Paleomagnetic data (Kent, 1985; Miller and Kent, 1988; Van Der Voo, 1988) indicate the

Appalachian Basin was located approximately 20° to 30° S latitude at the time of deposition of the Catskill Formation (Figure 3 A). Woodrow et al. (1973), Woodrow

(1985) and Ettensohn (1985) infer seasonal and arid paleoclimates based on these paleolatitudes. Paleoclimate has also been interpreted as seasonal and semi-arid from calcareous and vertic paleosols throughout much of the Catskill Formation (Mora et al.,

1991, Driese and Mora, 1993; Driese et al., 1997; Harvey and Grandstaff, 1997; Harvey,

1998; Retallack et al., 2009). Cecil et al. (2004) document non-calcic vertic paleosols and paleo- in the upper Hampshire Formation (a lateral equivalent of the Catskill

Formation) and interpret this as a transition from semi-arid to semi-humid during the latest Devonian. Palynological and paleobotanical data from the Catskill

Formation indicate similar paleoclimatic trends (Banks et al., 1985; Streel et al., 2000;

Cressler, 2001; Cressler et al., 2009). Cross-bedded sandstones, granitic dropstones in black marine shales, and a laterally continuous polymictic diamictite present at the base of the overlying Spechty Kopf Formation and time-equivalent strata across the region have been interpreted as glacial in origin, implying the onset of icehouse conditions at the

Devonian-Mississippian boundary (Brezinsky et al., 2008, Brezinsky et al., 2009,

Brezinsky et al., 2010; Ettensohn, 2014). McClung et al., 2013) document fourth-order sedimentary cycles within the Upper Devonian (a lateral equivalent of the Catskill Formation in and ) and attribute the cyclicity to

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glacioeustatic sea-level fluctuations, providing further evidence for the onset of icehouse conditions during the Famennian.

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Figure 3: Paleogeography and hypothetical cross section of the Appalachian Basin during the Late Devonian A) Approximate location of the “Catskill alluvial plane” relative to the Appalachian Basin during the Late Devonian (after Ettensohn, 1985 and Ryder et al., 2012). B) Schematic cross section along A-A’ of the Appalachian Basin during the Late Devonian. The Middle-Upper Devonian boundary is defined as the top of the Tully Limestone and laterally equivalent strata (Oliver et al., 1971; Carter, 2007). The Castkill Formation occupies the upper-most portion of the Acadian Clastic Wedge. Modified from Harper (2002).

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Figure 4: Generalized stratigraphic section of the Catskill Formation in Central Pennsylvania. The Clarks Ferry Member is present locally to the south and pinches out to the northwest. The exact location of the Frasnian-Famennian boundary is uncertain. Lithologic characteristics of the Irish Valley and Sherman Creek Members based on field observations and measured section from Cotter and Driese (1998). See Figure 5 for key.

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Figure 5: Key to symbols used in Figure 4, Figure 8, and Figure 9.

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Figure 6: Outcrop photographs showing various lithologic characteristics of the Irish Valley and Sherman Creek Members. A) Channel sandstone (outlined in yellow) within the Irish Valley Member at Selinsgrove, PA. Large arrow denotes stratigraphic up direction. B) Brachiopod from the Trimmers Rock-Irish Valley contact exposed near Selinsgrove, PA. C) Poorly-defined fining upward package (small arrow) capped by a thin, weakly-developed paleosol characteristic of the Sherman Creek Member at Selinsgrove, PA. Large arrow denotes stratigraphic up direction. S = paleosol surface. D) Incised-valley fill (IVF) within the Sherman Creek Member at Selinsgrove, PA. Large arrow denotes stratigraphic up direction. See Figure 2 for locations.

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Figure 7: Outcrop photographs showing various lithologic characteristics of the Duncannon and Clarks Ferry Members. A) Typical well-defined fining-upward packages (small arrows) capped by vertic paleosols within the Duncannon Member at Duncannon, PA. S = paleosol surface. B) Pedogenic slickenside (SS) (shown by arrow in a paleosol within the Duncannon Member at Duncannon, PA. C) Planar crossbedding in coarse-grained sandstone within the Clarks Ferry Member at Duncannon. D) Charcoal clast (shown by arrow) in hand sample collected from the Clarks Ferry Member at Duncannon. See Figure 2 for locations.

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1.4 Methods

Several hundred meters of the Catskill Formation with minor covered intervals are exposed near Selinsgrove and Duncannon, PA and serve as the primary localities for this study (Figure 2) (Appendix A). To test potential region variability in pedologic and sedimentologic characteristics, I also examine smaller (10’s of meters), isolated sections of the Catskill Formation located approximately along depositional strike of the complete sections near Cedar Run, Hyner (Red Hill locality), and Trout Run, PA (Powys Curve locality) (Figure 2) (Appendix A).

A road cut along US Route 11/15 exposes 456 m of the Irish Valley Member and

629 m of the Sherman Creek Member south of Selinsgrove, PA with only minor covered intervals (Figure 2) (Appendix A). Stratigraphic sections measured by Cotter and Driese

(1998) at this locality provide context for detailed paleosol measured sections and sample locations. The Duncannon Member is not exposed at this locality. A railroad cut parallel to US Route 322/22 southeast of Duncannon, PA exposes the upper portion of the

Sherman Creek Member, 57 m of the Clarks Ferry Member, and 235 m of the Duncannon

Member with minor covered intervals (Figure 2, Figure 8, Figure 9) (Appendix A). The upper contact of the Duncannon Member with the Spechty Kopf Formation is exposed at this locality. A stratigraphic section was measured from the Sherman Creek-Clarks Ferry

Member contact to the top of the Duncannon Member at this locality (Figure 8, Figure 9).

I describe three paleosol profiles representative of the most commonly occurring paleosol orders at the top, middle, and base of the Irish Valley, Sherman Creek, and

Duncannon Members. I do not define a pedotype for the Clarks Ferry Member as this interval generally lacks mudstone facies. For each profile, I document horizon

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thicknesses, color (using a Munsell (2000) chart), grain size, ped structures, degree and type of rooting, presence and type of mineral accumulations, and reaction with dilute HCl following methods of Retallack (1988). Samples from each horizon where taken for petrographic analysis following methods of Brewer (1964) and

Fitzpatrick (1993). Clay mineralogy was determined from representative, unoriented samples from each pedotype by X-ray diffraction at Bruker laboratories (analysis by

Henderson, 2015). XRD spectra were obtained using a Bruker D2 diffractometer at

30 kV and 10 mA equipped with a copper cathode. Based on these observations, I taxonomically classify these paleosols based on criteria of Mack et al. (1993) and USDA

Soil Survey Staff (2014) and develop pedotypes (sensu Retallack, 1994) for the Irish

Valley, Sherman Creek, and Duncannon Members.

I interpret prevailing soil forming conditions in each of the distinct depositional environments represented by individual lithostratigraphic members based on these observations and taxonomic classifications. Jenny (1941) states that the environmental factors which influence pedogenesis can be expressed as:

Equation 1: 풔 = 풇(풄풍, 풐, 풓, 풑, 풕) where s, the soil being evaluated, is a function of , (cl), flora and fauna, or the biotic factor (o), the topographic position on the landscape where pedogenesis is occurring (r), the parent material within which the soil is forming (p), and time available for pedogenic modification (t). To determine the relative contribution of each of these factors on the formation of a particular soil, one factor is allowed to vary while the others are hypothetically held constant. The relative effect of each factor of soil formation is then evaluated for each pedotype.

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I describe channel sandstone morphology in the field and collected oriented samples from sandstones at the top, middle, and base of each member of the Catskill

Formation for petrographic analysis. Twenty thin sections were prepared and stained following methods in Hutchinson (1974) for rapid identification of feldspars (Figure 8,

Figure 9). Point counting (n = 400) was performed to quantify variability in sandstone composition and sediment provenance up-section through the Catskill Formation.

Following the method of Van Der Plas and Tobi (1965), I determine point count modal percentage uncertainty of 2% to 5% at the 95% confidence level. To quantify grain size and evaluate up-section textural trends, I utilize Nikon NIS Elements image acquisition software to fit 5-point ellipses to the boundaries of at least 200 grains. This provides semi-major and semi-minor axis dimensions and maximum grain diameter is obtained by doubling the semi-major axis dimension.

th The shear stress (τC) necessary to entrain the 90 percentile grain diameter (D90) is calculated as a proxy for paleo-stream competency (Andrews, 1983; Komar, 1987,

Komar, 1989; Thompson and Croke, 2007; Oest et al., 2015). This assumes that the total area of distal channels on a DFS is greater than the area of the feeder channel and that discharge is constant.

Shear stress is calculated from the equation of Shields (1936):

Equation 2: 흉푪 = 흉푪∗(흆풔 − 흆풇)품푫풊 where 휏퐶 is shear stress, 휏퐶∗ is critical shear stress, 휌푠 is the density of sediment, 휌푓is the density of the fluid, 푔 is the acceleration due to gravity at Earth’s surface (-9.81 m s-2),

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and 퐷푖 is the grain diameter of interest. Shields (1936) determined that 휏퐶∗ is a function of the particle Reynolds number, which is given by

풖 푫 Equation 3: 푹 = ∗ 풊⁄흊 where R is the particle Reynolds number, 푢∗ is the shear velocity, Di is the grain diameter, and 휐 is kinematic viscosity. For R greater than 100, 휏퐶∗ approaches a constant of 0.06.

For streams with bed material 2 mm in diameter (coarse sand to fine pebbles) or greater,

R is typically greater than 500 (Andrews, 1983). I use at least coarse sand size bed material based on grain size data from channel sandstone samples and therefore use 휏퐶∗

-3 of 0.06. The sediment density, 휌푠, of 2650 kg m is used, as point count data show high

-3 percentages of quartz in sandstones sampled, and 휌푓 of 1000 kg m because water is the fluid. For these calculations, I use D90 for Di in order to obtain a maximum shear stress value and therefore gain insight into maximum flow competency. D90 is calculated from grain size populations for each thin section (see Appendix B for grain size distributions).

Critical shear stress varies with bed surface roughness (Andrews, 1983;

Thompson and Croke, 2007). Andrews (1983) demonstrates a scaling relationship between 휏퐶∗ and bed roughness, D50 by

∗ 푫풊 −풃 Equation 4: 흉푪풊 = 흉푪ퟓퟎ( ⁄ ) 푫ퟓퟎ

∗ where 휏퐶푖 is the critical shear stress for a given bed roughness, 휏퐶50 is an empirically derived constant based on D50 of the bed surface, Di is the grain diameter of interest, and b is curve fitting constant. (Andrews, 1983; Komar, 1987; Komar, 1989) apply this relationship to the original equation of Shields (1936) (Equation 2) to obtain

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풃 ퟏ−풃 Equation 5: 흉푪 = 흉푪ퟓퟎ∗(흆풔 − 흆풇)품푫ퟓퟎ푫ퟗퟎ where 휏퐶50∗ is the critical shear stress adjusted to the D50 of the bed surface. In this study, a 휏퐶50∗ of 0.0834 and b of 0.872 are used for calculation of shear stress given a bed surface of coarse sand to pebbles. Both values are taken from Andrews (1983). Similarly,

휏퐶50∗ of 0.045 and b of 0.60 are used to calculate shear stress assuming a pebble bed surface. These values are taken from Komar (1987). Bed roughness variables, D90 and

D50, are calculated from grain size populations for each thin section. Density parameters and g are not changed from the calculation using Shields (1936) equation. Parameters used in shear stress calculations are summarized in Table 2.

Profile thicknesses are not representative of the original thickness due to burial compaction. To evaluate the effect of compaction due to burial, I utilize the relationship between and compaction derived by Sheldon and Retallack (2001):

−푺풊 Equation 6: 푪 = 푭 [( ퟎ⁄ )−ퟏ] 풆푫풌 where C is compaction as a percent of original profile thickness, Si is the initial solidity,

F0 is the initial porosity, D is burial depth in kilometers (input as negative), and k is a curve-fitting constant specific to a particular soil order. I use initial solidity (Si) and initial porosity (F0) values for each pedotype from Sheldon and Retallack (2001) (Table 3).

Fluid inclusion from veins within Middle Devonian shales in the central

Valley and Ridge province indicate burial depths ranging from 5 to 11 km (Evans et al.,

2014). Given a total thickness of approximately 500 m of Middle Devonian rocks in central Pennsylvania (Oliver et al., 1971), I use a maximum burial depth for the base of

21

the Catskill Formation of 10.5 km. Table 3 summarizes constants used to calculate values of compaction for each pedotype.

22

Figure 8: Measured section of the Clarks Ferry and Duncannon Members at Duncannon, PA showing lithology and location of described paleosols and sandstone samples. Continued in Figure 9. See Figure 5 for explanation of symbols. DC- CHSST-CO-# = sandstone sample, DC-PS3-14700 = Paleosol profile location. C = Clay, St = , S = Sand, G = Gravel.

23

Figure 9: Measured section of the Duncannon Member at Duncannon, PA, continued. See Figure 5 for explanation of symbols. DC-CHSST-CO-# = sandstone sample, DC-PS1-29700 = Paleosol profile location.C = Clay, St = Silt, S = Sand, G = Gravel.

24

Table 2: Shear Stress Calculation Parameters

Source Equation Bed Surface Parameters

-1 Shields, 1936 휏퐶 = 휏퐶∗(휌푠 − 휌푓)푔퐷90 Glass spheres 휏푐∗ = 0.6, g = 9.81 m s

Coarse sand to Andrews, 1983 휏 = 휏 (휌 − 휌 )푔퐷푏 퐷1−푏 휏 = 0.0834, b = 0.872, g = 9.81 m s-1 퐶 퐶50∗ 푠 푓 50 90 pebbles 퐶50∗

Komar, 1987; 휏 = 휏 (휌 − 휌 )푔퐷푏 퐷1−푏 Pebbles 휏 = 0.045, b = 0.60, g = 9.81 m s-1 Komar, 1989 퐶 퐶50∗ 푠 푓 50 90 퐶50∗

25

Table 3: Decompaction Constants

Soil Order Initial Solidity Initial Porosity k (Curve fitting Pedotype (USDA) (Si) (F0) constant)

1 0.65 0.35 0.15

2 0.51 0.49 0.27

3 0.69 0.31 0.12

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CHAPTER 2. RESULTS

2.1 Paleopedology

2.1.1 Pedotype 1

Pedotype 1 is developed based on descriptions of three paleosols within the Irish

Valley Member. The type paleosol for Pedotype 1 is located 239 m above the base of the

Irish Valley Member at Selinsgrove (Table_4) (Figure_10) (see Figure 4 of Cotter and

Driese (1998) for measured section of the Irish Valley Member.). The paleosol is composed of predominately grayish red (5R 4/2) mudstone with grayish yellow green

(5GY 7/2) to pale olive (10Y 6/2) reduction mottles occurring throughout the profile

(Figure 10, Figure_11A, B). Horizons are weakly developed and have gradual, undulatory contacts (Figure 11 A). Drab-haloed root traces oriented perpendicular to bedding are rare and occur within the upper 25 to 30 centimeters of the profile.

Petrographic analysis shows clay coating ped surfaces and sub-millimeter scale clay infilled roots occur throughout the profile (Figure_12 A, B). Fine to medium (1 to 5 mm diameter) granular ped structures dominate the upper portion of the profile and gradually transition to medium to coarse (10 to 50 mm diameter) angular blocky, clay-coated ped structures with depth. Parent material is grayish red siltstone to very fine, massive sandstone. None of the horizons from any of the paleosols of this pedotype react with dilute hydrochloric acid. Aligned clay are present in voids and argillans on ped structures. (Figure 12 A, B). Mosepic, omnisepic, skelsepic and vosepic soil microfabrics are present in this paleosol. The upper portion of the profile is dominated by mosepic fabric transitioning to omnisepic fabric with depth. The base of the profile is dominated by skelsepic to vosepic fabric. Irregularly shaped, opaque grains interpreted as detrital

27

organic matter are common at the top of the profile. The X-ray diffractogram for

Pedotype 1 shows the presence of chlorite, illite, muscovite, feldspars, and quartz

(Figure_13). Pedogenic slickensides on ped surfaces are rare, occurring in only one of the profiles examined. Pedotype 1 is classified as an aqualf (an occasionally waterlogged alfisol) using the criterial of Staff (2014) and an argillic using the criteria of Mack et al. (1993) (Table_5). These classifications are based on the presence of illuviated clays (t), and hydromorphic features (g) throughout the profile (Figure_10).

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Table 4: Pedotype 1 Description

Depth Interpreted Macromorphology Micromorphology (cm) Horizon 5R 4/2 (grayish red) mudstone, small, crumby Insepic to mosepic peds, 3 – 5 cm diameter 5G microfabric, 6/1 (greenish gray) reduction abundant irregularly spots, top 1 cm of profile 0- 15 shaped opaque Ag entirely 5G 6/1, no reaction grains interpreted as with dilute hydrochloric acid, detrital organic sharp undulatory contact with matter underlying horizon

5R 4/2 (grayish red) siltstone, small to medium angular blocky peds, argillans, rare drab-halo root traces (~3 cm in length), 3 – 5 cm diameter Omnisepic 5G 6/1 (greenish gray) microfabric, clay 15-30 Btg reduction spots, gradual, no infilled roots, reaction with dilute reduced zones hydrochloric acid, undulatory contact with underlying horizon

10R 4/2 (weak red) very fine grained sandstone, occasional Skelsepic to vosepic large blocky peds, 3 – 5 cm microfabric, clay diameter 5G 6/1 (greenish infilled roots and 30-123 gray) reduction spots, no Cg  Btg drab haloed root reaction with dilute traces, reduction hydrochloric acid, base of spots, relict bedding profile

29

Figure 10: Lithologic sections of pedotypes illustrating macroscopic pedogenic and sedimentary features.

30

Figure 11: Outcrop photographs of pedotype paleosols. Large arrows indicate stratigraphic up directions. A) Pedotype 1 profile. B) Mottling (M) and granular peds (GP) in lower portion of Ag horizon from A. C) Pedotype 2 profile. Note gradual transitions between horizons. D) Carbonate nodules (CN) near the base of Pedotype 2. E) Pedotype 3 profile. Note wedge shape ped structures (W) and pedogenic slickensides (SS). F) Close up of pedogenic slickensides (SS) and angular blocky ped structures (AB) within Bss horizon of Pedotype 3.

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Figure 12: Photomicrographs of pedogenic features. A) Clay (C) lining ped structure (P) boundaries in Btg horizon of Pedotype 1. of Pedotype 1. C) Original pedogenic carbonate (PC) surrounded by recrystallized

Large arrow indicates stratigraphic up direction. E) Omnisepic and clinobimasepic soil microfabric, Bss horizon, Pedotype 3. Note rare skeletal quartz grains (Q). F) Omnisepic soil microfabric, Ass horizon, Pedotype 3. All photomicrographs taken in cross-polarized light and are unoriented unless otherwise noted.

32

Figure 13: X-ray diffractograms from representative samples for each Pedotype. Note the abundance of illite (I) and chlorite (Cl). M = muscovite, Q = quartz, F = feldspar

33

Table 5: Pedotype Classifications

Pedotype USDA (Soil Survey Staff, 2014) Mack et al. (1993)

Pedotype 1 aqualf argillic gleysol

Pedotype 2 ustept argillic

Pedotype 3 udert argillic vertisol

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2.1.2 Pedotype 2

The type paleosol for Pedotype 2 is located 249 m above the base of the Sherman

Creek Member at Selinsgrove (Table 6) (Figure 10) (See Figure 6 in Cotter and Driese

(1998) for measured section of the Sherman Creek Member.). This paleosol is composed of grayish red (10R 4/2) to very dusky red (10R 2/2) mudstone and laminated, occasionally rippled siltstone at its base (Figure 10, Figure 11 C). A thin, very fine sandstone interval occurs within the upper laminated siltstone portion of the profile.

Horizons are poorly to moderately developed and are separated by gradual contacts

(Figure 11 C). Fine to medium angular blocky ped structures (1-20 mm diameter) occur in the upper 24 cm of the profile. Thin platy peds (1-5 mm) are present from 24 cm depth to the base of the profile. Pedogenic carbonate nodules ranging from 4 to 20 mm in diameter occur at ~52 cm depth and are increasingly abundant with depth (Figure 11 D).

All horizons react strongly to dilute hydrochloric acid. Very fine clay-infilled roots are present in the same interval as the carbonate nodules. Carbonate nodules are not evident in outcrop for one paleosol examined within the Sherman Creek Member, but are present in thin section. Nodules occur as channel lags in sandstones overlying truncated paleosols. Argillans are not evident at the macroscopic scale, but petrographic analysis shows abundant illuviated clay material in burrows and root voids (Figure 12 D).

Similarly, pedogenic calcite is only evident at the micro-scale in horizons which lack visible carbonate nodules. Much of the calcite has a sparry morphology, although micritic textures are still occasionally present (Figure 12 C). Parent material is mudstone, grayish red laminated siltstones, and very fine grained, grayish red sandstones. Soil microfabric transitions from bimasepic to silasepic with increasing depth in the profile. Vosepic

35

fabric is occasionally present in zones dominated by silasepic fabric. This trend in microfabric occurs twice within the profile, suggesting that this paleosol is a compound soil. The X-ray diffractogram for Pedotype 2 is similar to Pedotype 1 indicating the presence of chlorite, illite, muscovite, feldspar, and quartz (Figure 13). The presence of these minerals is confirmed by petrographic analysis. Vertic features are largely absent from paleosols within this stratigraphic interval and are poorly expressed when they are present. Pedotype 2 is classified as an ustept (an inceptisol with carbonate accumulations) using the criteria of Soil Survey Staff (2014) and an argillic calcisol using the criteria of

Mack et al. (1993) (Table 5). These classifications are based on the presence of pedogenic carbonate (k) and argillic (t) and cambic (w) horizons (Figure 10).

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Table 6: Pedotype 2 Description

Depth Interpreted Macromorphology Micromorphology (cm) Horizon 10R 4/2 (grayish red) mudstone Silasepic to masepic to siltstone, small to medium microfabric, occasional angular blocky peds, strong unaltered feldspar 0- 24 reaction with dilute grains and abundant Bwk hydrochloric acid. Gradual, unaltered muscovite undulatory contact with grains underlying horizon

Silasepic to vosepic microfabric, clay 10R 4/2 (grayish red) infilled roots, laminated, occasionally rippled recrystallized 24-52 siltstone, strong reaction with carbonate, oriented Ck  Btk dilute hydrochloric acid. Sharp clay infilling burrow, contact with underlying horizon primary depositional fabric still identifiable

10R 4/2 (grayish red) to 10R 2/2 (very dusky red) mudstone, platy ped stuctures with occasional argillans, carbonate Vosepic to mosepic nodules 4 to 20 mm in diameter microfabric, recrystallized Ck  Btk 52-116 which become more abundant with depth, very fine clay carbonate, clay infilled infilled roots. Gradual contact roots, isolated patches with underlying laminated of oriented clay films siltstone bed

37

2.1.3 Pedotype 3

The type paleosol for Pedotype 3 is located 297 meters above the Sherman Creek-

Clarks Ferry Member contact within the Duncannon Member at Duncannon (Figure 9)

(Table 7) This paleosol is a predominately grayish red (5 to 10R 4/2) mudstone at the top to very fine grained sandstone at the base (Figure 10, Figure 11 E). Horizons are strongly developed with sharp well-defined, undulatory contacts (Figure 11 E). Pedogenic slickensides, curvilinear fracture planes, and wedge shaped ped structures are the most common marcromorphological feature in these paleosols and dominate the upper portion of the profiles (Figure 11 E, F). Wedge shaped and medium to large (10 to 50 mm diameter) angular blocky ped structures are common throughout the profile

(Figure 11 E). Crumby ped structures occur in the top 10 to 20 cm of the profile and platy ped structures gradually transition to unaltered sandstone at the base of the profile. Drab haloed root traces approximately 5 to 10 cm in length oriented perpendicular to bedding are visible at the macro-scale and occur through much of the profile. Sub-millimeter scale clay-infilled roots are evident at the micro-scale. Pedotype 3 generally occurs on top of fine to medium-grained crossbedded sandstone. Soil microfabric is omnisepic through much of the profile transitioning to argillasepic and silasepic at the base of the profile

(Figure 12 E). Clinobimasepic fabric is present in horizons with visible pedogenic slickensides in outcrop (Figure 12 E). The X-ray diffractogram of Pedotype 3 indicates the presence of chlorite, illite, muscovite, feldspars, and quartz (Figure 13). Variability within Pedotype 3 includes the abundance of pedogenic slickensides, profile thickness, and the presence, type, and abundance of rooting. Pedotype 3 is classified as an udert

(a vertisol lacking carbonate accumulations) using the criteria of Soil Survey Staff (2014)

38

and an argillic vertisol using the criteria of Mack et al. (1993) (Table 5). These classifications are based on the presence of pedogenic slickensides (ss), translocated clays

(t), wedge shaped peds, and a lack of pedogenic carbonate (Figure 10).

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Table 7: Pedotype 3 Description

Depth Interpreted Macromorphology Micromorphology (cm) Horizon 5R 4/2 (grayish red) Omnisepic and mudstone, crumby to medium occasionally angular blocky ped structures, argillasepic minor pedogenic slickensides, microfabirc, drab 0- 64 drab haloed root traces, no haloed root traces, Ass reaction with dilute argillans hydrochloric acid. Sharp occasionally contact with underlying coating skeletal horizon grains

5R 4/2 (grayish red) siltstone, medium to large angular Clinobimasepic blocky and wedge shaped ped and omnisepic structures, abundant microfabric, 64-146 pedogenic slickensides, drab argillans Bss haloed root traces, no reaction occasionally with dilute hydrochloric acid. coating skeletal Sharp contact with underlying grains horizon

Omnisepic and 10R 4/2 (grayish red) very occasionally fine to fine grained sandstone, argillasepic angular blocky ped structures, microfabric, BtC 146-171 drab haloed root traces, no infilled roots, reaction with dilute oriented clay films, hydrochloric acid delaminating muscovite grains

5R 4/2 (grayish red) very fine grained sandstone, angular Skelsepic to blocky ped structures argillasepic and gradually transitioning to silasepic 171-205 platy ped structures at base, microfabric, clay CBt no reaction with dilute infilled roots, hydrochloric acid. Gradual isolated oriented contact with underlying very clay films fine grained sandstone

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2.2 Channel Sandstone Petrography

2.2.1 Mineralogy

Channel sandstones within the Irish Valley Member range in composition from sublitharenites to litharenites, with an average composition of litharenite based on the

Folk classification (1980) (Figure 14). The average relative percentages of total quartz, feldspar, and lithic grains is 68±18%, 6±3%, and 26±17%, respectively. Quartz grains are almost exclusively monocrystalline, with only 6 polycrystalline grains counted from

3 samples taken from the Irish Valley Member (Table 8). Feldspars are largely plagioclase, with only 4 potassium feldspar grains counted (Table 8). Lithic fragments are primarily composed of schist grains, although quartzites are occasionally present

(Figure 15 A, B). Chert and siltstone clasts are the principle sedimentary lithic clasts.

Sandstones are grain supported with concavo-convex and occasionally sutured grain boundaries. Although syntaxial quartz overgrowth and phyllosilicates infilling primary porosity are the most prevalent cements in these sandstones, poikilotopic calcite cement is also present. Accessory minerals include muscovite, chlorite, and rare detrital zircons

(Figure 15 A).

Sandstones within the Sherman Creek Member are classified as litharenites with an average composition of 58±5% total quartz, 6±3% feldspar, and 37±6% lithic grains

(Figure 14). Quartz grains are largely monocrystalline, although minor amounts of polycrystalline quartz are present (Table 8). Feldspars are mainly plagioclase with minor amounts of potassium feldspar. One sample from the top of the Sherman Creek Member contains an anomalous amount of potassium feldspar (Table 8). Lithic grains consist of schist, quartzite, chert, and shale clasts (Figure 15 C). These sandstones are grain

41

supported with concavo-convex grain boundaries and, to a lesser extent, sutured grain contacts (Figure 15 C). Poikilotopic calcite cement is common with lesser amounts of syntaxial quartz overgrowths, phyllosilicate, and iron oxide cements. Muscovite is common as an accessory constituent. Feldspar grains are partially altered to clays.

Channel sandstones within the Clarks Ferry Member range in composition from sublitharenites to litharenites, with an average relative percentage of total quartz, feldspar, and total lithics of 71±5%, 2±1%, and 28±6%, respectively (Figure 14). Most quartz grains are monocrystalline, although minor amounts of polycrystalline quartz are present (Table 8). Sandstones within the Clarks Ferry Member generally lack plagioclase and potassium feldspar (Table 8). Metamorphic lithic grains consist largely of quartzites and gneisses (Figure 15 D). Sedimentary lithics include chert and shale clasts, and less frequently, sandstone clasts. Channel sandstones of the Clarks Ferry Member are grain supported with concavo-convex and sutured grain boundaries. Cements include iron oxides (likely hematite) and phyllosilicates (likely illite and sericite) infilling pore spaces, as well as syntaxial quartz overgrowths (Figure 14 D, E). Chlorite, muscovite, and to a lesser extent biotite occur as accessory minerals. Phyllosilicates show deformation due to compaction when in contact with more competent grains.

Channel sandstones of the Duncannon Member are classified as litharenites with an average composition of 65±4% total quartz, 3±1% feldspar, and 32±4% lithic grains

(Figure 14). Monocrystalline quartz is the principle variety of quartz within these sandstones (Table 8). Plagioclase and potassium feldspar are both present in minor amounts, with plagioclase more common than potassium feldspar. Lithic grains largely consist of schists and quartzites with occasional gneiss clasts. Chert clasts comprise the

42

majority of sedimentary lithic grains but siltstone and sandstone clasts are also present

(Figure 15 F). These sandstones are grain supported with concavo-convex and sutured grain boundaries. Poikilotopic calcite and syntaxial quartz overgrowth cements are common (Figure 15 F). Iron oxides (likely hematite) and phyllosilicates (likely illite and sericite) infilling pore spaces also act as cements (Figure 15 F). Accessory minerals include muscovite, chlorite, and minor amounts of biotite.

43

Figure 14: Ternary diagram showing composition of sand-sized or larger particles and inferred provenance. Modal percentages of total quartz, feldspar, and total lithic grains indicate a recycled orogenic source terrane for the Catskill Formation. Most sandstones of the Catskill Formation are litharenites in composition. Two sandstones from the Clarks Ferry Member and one sandstone from the Irish Valley Member are sublitharenites. One sample from the Irish Valley Member is a feldspathic litharenite.

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Table 8: Raw Point Count Data Counts Percentages Sample ID Member Qt Qm Qp F P K L Lm Ls Lv Qt F L SG-CHSST-CO-23900 Dciv 170 170 0 21 21 0 151 64 87 0 50±5 6±4 44±5 SG-CHSST-CO-34500 Dciv 318 317 1 12 11 1 43 5 38 0 85±4 3±2 12±3 SG-CHSST-CO-44600 Dciv 287 282 5 34 21 3 88 30 58 0 70±5 8±3 22±4 Irish Valley Average 68±18 6±3 26±17 SG-CHSST-IVF1-C0-5800 Dcsc 203 185 18 20 17 3 169 71 98 0 52±5 5±3 43±5 SG-CHSST-CO-14500 Dcsc 225 197 28 19 19 0 110 70 40 0 64±5 5.±3 31±5 SG-CHSST-CO-24700 Dcsc 215 215 0 6 6 0 154 71 83 0 57±5 2±1 41±5 SG-CHSST-CO-61800 Dcsc 239 224 15 33 32 1 114 64 50 0 62±5 9±3 30±5 DC-CHSST-CO-20 Dcsc 207 185 22 29 7 22 155 48 107 0 53±5 7±3 40±5 Sherman Creek Average 58±5 6±3 37±6 DC-CHSST-CO-680 Dccf 271 271 0 6 0 6 110 66 44 0 70±5 2±1 28±5 DC-CHSST-CO-2800 Dccf 323 294 29 3 1 2 79 25 54 0 80±4 1±1 20±4 DC-CHSST-CO-3191 Dccf 246 246 0 7 1 6 108 62 46 0 68±5 2±1 30±5 DC-CHSST-CO-3200 Dccf 243 243 0 1 0 1 120 61 59 0 67±5 0 33±5 DC-CHSST-CO-4100 Dccf 223 207 16 6 5 1 66 34 32 0 76±5 2±2 22±5 DC-CHSST-CO-6800 Dccf 216 206 10 8 8 0 111 44 67 0 64±5 2±2 33±5 Clarks Ferry Average 71±5 2±1 28±6 DC-CHSST-CO-7000 Dcd 220 216 4 11 11 0 86 34 52 0 69±5 3±2 27±5 DC-CHSST-CO-10003 Dcd 214 207 7 9 5 4 81 26 55 0 70±5 3±2 27±5 C-CHSST-CO-14929 Dcd 229 228 1 6 4 2 132 59 73 0 62±5 2±1 36±5 DC-CHSST-CO-18830 Dcd 212 207 5 16 10 6 126 61 65 0 60±5 5±3 36±5 DC-CHSST-CO-19900 Dcd 252 235 17 11 8 3 116 45 71 0 66±5 3±2 31±5 DC-CHSST-CO-29100 Dcd 183 169 14 6 5 1 101 58 43 0 63±5 2±2 35±5 Duncannon Average 65±4 3±1 32±4 Note: Dciv = Irish Valley Member, Dcsc = Sherman Creek Member, Dccf = Clarks Ferry Member, Dcd = Duncannon Member

45

Figure 15: Photomicrographs of channel sandstones from the Catskill Formation. A) Monocrystalline quartz (MQ), sedimentary lithic clasts (LS), and detrital zircon (ZR) as accessory mineral within the Irish Valley Member.. B) Monocrystalline quartz (MQ) and metamorphic lithic (schist) clasts (LM) within the Irish Valley Member. C) Recrystallized metamorphic quartz (LM) and monocrystalline quartz (MQ) within the Sherman Creek Member. Note deformed muscovite grain (DM) due to compaction. D) Sutured grain boundaries (S) between monocrystalline quartz (MQ) and recrystallized metamorphic quartz (LM) within the Clarks Ferry Member. Cement in this sample is illite and sericite (I/S) infilling primary porosity. E) Syntaxial quartz overgrowth (QO) cement between chert (CT), monocrystalline quartz (MQ), and metamorphic lithic (LM) grains within the Clarks Ferry Member. F) Poikilotopic carbonate (PC), quartz overgrowth (QO), and phyllosilicate (PS) cement between monocrystalline quartz (MQ) grains and sedimentary lithic (LS) grains within the. Duncannon Member. All photomicrographs taken in cross polarized light and oriented perpendicular to bedding.

46

2.2.2 Grain Size Analysis

Channel sandstones within the Irish Valley Member have mean grain sizes which range from 89 to 103 µm (very fine sand) (Figure 16) (Table 9). Maximum grain size increases up-section from 221 to 267 µm (fine to medium sand) (Figure 16) (Table 9).

Similarly, minimum grain size increases from 32 to 33 µm (coarse silt) up-section

(Figure 16) (Table 9). These sandstones are moderately to well-sorted with standard deviations ranging from 32 to 40 µm and sub-rounded to sub-angular (Figure 16)

(Table 9). Grain size distributions from all sandstones within the Irish Valley are unimodal and positively skewed (Appendix B).

Mean grain size of channel sandstones within the Sherman Creek Member range from 135 to 212 µm (fine sand) and generally fine upward within this stratigraphic interval (Figure 16) (Table 9). Maximum grain size also fines up-section from 646 to 364

µm (coarse to medium sand) with the exception of one sample taken from the top of the

Sherman Creek Member at Duncannon, PA which has a maximum grain size of 939 µm

(coarse sand) (Figure 16) (Table 9). Minimum grain size shows a similar trend of fining up-section from 84 to 39 µm (very fine sand to coarse silt) with the exception of the sample from the top of Sherman Creek Member, which has a minimum grain size of

48 µm (coarse silt) (Figure 16) (Table 9). Sorting is moderate throughout this member although sandstones become increasingly well-sorted up-section as standard deviations decrease from 82 to 54 µm up-section (Table 9). All samples have unimodal and positively skewed grain size distributions, although the degree of skew varies slightly between samples (Appendix B). Rounding ranges from angular (plagioclase grains) to well-rounded (quartz and sedimentary lithic grains).

47

Mean grain size of channel sandstones fines up-section through the Clarks Ferry

Member from 502 to 118 µm (coarse to very fine sand) (Figure 16) (Table 9). Maximum and minimum grain size fine upward in general from 1650 to 265 µm (very coarse to medium sand) and 156 to 52 µm (fine sand to coarse silt), respectively (Figure 16)

(Table 9). Sorting varies from well to poorly sorted throughout the Clarks Ferry Member with standard deviations ranging from 34 to 232 µm with increased sorting up-section

(Figure 16) (Table 9). Grain size distributions are unimodal and positively skewed with the exception of one sample which shows a nearly normal distribution (Appendix B).

Grains are sub-rounded to sub-angular.

Channel sandstone mean gain size within the Duncannon Member generally coarsens up-section from 115 to 336 µm (very fine to medium sand) (Figure 16)

(Table 9). Maximum grain size increases from 229 to 1613 µm (fine to very coarse sand) from the base to the middle of the section before decreasing from 1088 to 815 µm (very coarse to coarse sand) from the middle to the top of the section (Figure 16) (Table 9).

Standard deviation increases up-section from 34 to 120 µm, although much of the section is moderately sorted (Table 9). Sandstones within the Duncannon Member are sub-rounded to sub-angular. All samples show unimodal and positively skewed grain size distributions (Appendix B).

48

Figure 16: Textural trends through the Catskill Formation. A) Irish Valley – Sherman Creek contact is approximately 500 m above the base of the section. Note fining-upward packages denoted by arrows. B) Clarks Ferry – Duncannon contact is approximately 70 m above the Sherman Creek – Clarks Ferry contact. Note change in horizontal scale, fining-upward packages denoted by arrows, and overall coarsening-upward trend.

49

Table 9: Grain Size Data Mean Grain Maximum Grain Minimum Grain Median Grain Standard Sample ID Member n Size (µm) Size (µm) Size (µm) Size (µm) Deviation SG-CHSST-CO-23900 Dciv 338 89 221 32 84 32 SG-CHSST-CO-34500 Dciv 200 78 262 22 73 32 SG-CHSST-CO-44600 Dciv 245 103 267 33 99 40 Irish Valley Average 90 250 29 86 35 SG-CHSST-IVF1-C0-5800 Dcsc 206 212 646 84 194 82 SG-CHSST-CO-14500 Dcsc 216 169 458 38 160 69 SG-CHSST-CO-24700 Dcsc 301 146 431 48 139 58 SG-CHSST-CO-61800 Dcsc 352 135 364 39 124 54 DC-CHSST-CO-20 Dscs 201 207 984 48 185 120 Sherman Creek Average 174 577 51 160 76 DC-CHSST-CO-680 Dccf 185 502 1650 156 462 232 DC-CHSST-CO-2800 Dccf 224 146 430 43 109 87 DC-CHSST-CO-3191 Dccf 220 297 887 100 273 136 DC-CHSST-CO-3200 Dccf 253 160 390 58 145 67 DC-CHSST-CO-4100 Dccf 212 287 575 67 287 109 DC-CHSST-CO-6800 Dccf 201 118 265 52 113 34 Clarks Ferry Average 252 699 79 232 111 DC-CHSST-CO-7000 Dcd 200 116 229 44 112 34 DC-CHSST-CO-10003 Dcd 241 168 496 60 153 76 DC-CHSST-CO-14929 Dcd 237 207 1613 48 191 128 DC-CHSST-CO-18830 Dcd 201 309 1088 81 294 133 DC-CHSST-CO-19900 Dcd 202 257 588 42 242 115 DC-CHSST-CO-29100 Dcd 222 336 815 71 320 120

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Table 9: Grain Size Data (Cont'd) Mean Grain Maximum Grain Minimum Grain Median Grain Standard Sample ID Member n Size (µm) Size (µm) Size (µm) Size (µm) Deviation Duncannon Average 232 805 58 219 101 Note: Dciv = Irish Valley Member, Dcsc = Sherman Creek Member, Dccf = Clarks Ferry Member, Dcd = Duncannon Member

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2.3 Channel Morphology

The base of the Irish Valley Member is characterized by rare channel sandstones occurring as isolated bodies in overbank fines and common thin to medium bedded sheet sandstones. Channel sandstones become more common up-section starting at 247 m above the base of the Irish Valley Member. Channels are asymmetric, single-story bodies occasionally showing asymmetric fill patterns, suggesting lateral migration and accretion of isolated channels (Gibling, 2006). Within the Sherman Creek Member, channel sandstones become larger, relatively coarser-grained, and more common than in the underlying Irish Valley Member (Figure 16) (Table 9). Channels occur as amalgamated, succession-dominated multistory bodies within poorly-defined fining-upward packages commonly capped by thin ustepts of Pedotype 2. A distinct change from relatively fine- grained, amalgamated channels associated with the Sherman Creek Member to coarser- grained (occasionally conglomeratic), succession-dominated, multistory channel bodies with thick, well-defined lateral accretion sets occurs seven meters above the Sherman

Creek-Clarks Ferry Member contact at Duncannon, PA (Figure 8). Channel sandstones within the Duncannon Member are thick, laterally extensive, asymmetric single-story channel bodies which occur at the base of well-defined fining-upward packages often capped by uderts of Pedotype 3.

2.4 Decompaction

Given a maximum burial depth of 10.5 km, Pedotypes 1, 2, and 3 are 66%, 52%, and 75% of their original thicknesses, respectively. Variability in compaction is attributed

52

to different initial porosities and solidities for each soil order. Original profile thicknesses for Pedotypes 1, 2, and 3 are 177 cm, 172 cm, and 256 cm, respectively.

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CHAPTER 3. DISCUSSION

3.1 Paleoenvironmental Factors of Soil Formation

3.1.1 Paleoclimate

Appalachian basin Paleoclimate is characterized by a shift from semiarid greenhouse conditions to sub humid icehouse conditions during the Late Devonian

(Buggisch, 1991; Cecil et al., 2004; Averbuch et al., 2005; Brezinski et al., 2008;

Brezinski et al., 2009; Brezinski et al., 2010; Haq and Schutter, 2008). This transition coincides with the onset of glaciation on Gondwana and Euramerica (Caputo and

Crowell, 1985; Streel et al., 2000; Brezinski et al., 2008, 2009, 2010). In particular, mean annual precipitation (MAP) within the Appalachian Basin increased during the latest

Devonian before decreasing again at the Devonian-Mississippian boundary (Cecil et al.,

2004; Brezinski et al., 2008, 2009, 2010; Retallack et al., 2009; Retallack, 2011). The relative effect of paleoprecipitation on the formation of each pedotype can be inferred from rooting depth and morphology, the presence and abundance of translocated clay minerals, accumulations of soluble bases, and the up-section variability in interpreted paleosol orders.

Rooting in Pedotype 1 typically extends 25 to 30 cm into the soil profile and consists largely of drab haloed root traces. However, fine clay-infilled roots are also present at the sub-millimeter scale (Figure 12). Sub-millimeter scale clay-infilled roots occur up to 70 cm depth in one Pedotype 1 profile. This profile also lacks hydromorphic features common to other Pedotype 1 profiles, which suggests that this soil formed on a local topographic high which resulted in well-drained conditions. Rooting is not evident at the outcrop scale in all but one of the Pedotype 2 profiles, however sub-millimeter

54

scale roots are identifiable in thin section at depths up to 1 m. Millimeter-scale roots are present at depths up to 2 m in one Pedotype 2 profile. Drab haloed root traces up to 10 cm in length and oriented perpendicular to bedding are common and occur at depths up to

2 m in all Pedotype 3 profiles. Clay-infilled roots occur over the same depth interval for all Pedotype 3 profiles but are only evident at the micro-scale. The overall increasing depth and abundance of rooting from Pedotype 1 to Pedotype 3 suggests relatively decreasing paleoprecipitation throughout the section. However, this trend is not consistent with other accounts of Late Devonian paleoclimate (Cecil et al., 2004;

Brezinski et al., 2008; Brezinski et al., 2009; Brezinski et al., 2010; Retallack et al., 2009;

Retallack, 2011), which imply rooting depth is not controlled by paleoclimate. Depth to the water table, and therefore soil drainage potential, also controls rooting depth.

Hydromorphic features and shallow rooting within Pedotype 1 suggests a shallow water table. Pedogenic carbonate nodules, clay infilled roots penetrating deeper into Pedotype 2 profiles, and a lack of hydromorphic features suggest better drainage than Pedotype 1.

Similarly, deep rooting in Pedotype 3 reflects greater water table depths rather than decreased paleoprecipitation given the lack of pedogenic carbonate in these paleosols, although Retallack et al. (2009) document calcareous paleosols similar to Pedotype 3 within the Duncannon Member at the Red Hill locality, suggesting regional variability in soil drainage potential. These overall trends in rooting depth up-section through the

Catskill Formation suggest increasing water table depth consistent with the DFS deposition model (Figure 1) (Hartley et al., 2013; Weissmann et al., 2013).

Argillans on ped surfaces (Figure 12 A, B, D) and oriented clay infilling voids

(either roots or burrows) are observed in all pedotypes and suggest sufficient

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paleoprecipitation to mobilize clay minerals downward through the soil profile. Illuviated clays, well-developed soil microfabric, and hydromorphic mottling in all Pedotype 1 profiles suggests overprinting of two distinct soil forming conditions. The co-occurrence of illuviated clays and well-developed microfabric with hydromorphic features suggests seasonality during pedogenesis in Pedotype 1 paleosols. During drier periods, water table depth increases locally and allows for downward percolation of water through the profile.

Prolonged periods of precipitation results in decreasing depth to the water table, which inhibits translocation of clays. Alternatively, a longer term, overall increase in paleoprecipitation could also produce this overprinting. Pedogenic carbonate accumulation coupled with translocated clays in Pedotype 2 suggests increasing aridity.

During times of increased paleoprecipitation, carbonate and clay minerals would have been mobilized through the profile. Increased aridity then allowed for the precipitation of carbonate nodules before burial and preservation of the soil. The lack of carbonate accumulation and the presence of illuviated clays and vertic features in Pedotype 3 suggests a relative increase in paleoprecipitation compared to Pedotype 2. The co-occurrence of translocated clay with hydromorphic features and pedogenic carbonate in Pedotypes 1 and 2 suggests an influence of a seasonal paleoclimate overprinting the effect of variation in geomorphic position.

Classification of paleosols based on USDA Soil Taxonomy (Soil Survey Staff,

2014) allows for interpretation of paleoclimatic conditions based on known climate conditions necessary to form different soil orders. Pedotypes 1, 2 and 3 are classified as an aqualfs, ustepts, and uderts, respectively. form in semiarid (250-500 mm yr -1

MAP) to humid (1000-2000 mm yr -1 MAP) climates (Bull, 1991; Retallack, 2001; Cecil

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and Dulong, 2003; Soil Survey Staff, 2015). form over a wide range of climatic conditions and therefore the classification of inceptisol alone is a poor indicator of climate during pedogenesis (Soil Survey Staff, 2015). form in dry subhumid to moist humid climates (500 – 2000 mm yr -1 MAP) (Bull, 1991; Retallack, 2001; Cecil and Dulong, 2003; Soil Survey Staff, 2015). While the up-section transition from predominantly alfisols to vertisols could potentially reflect increasing seasonality, it is more likely that this shift is the result of a change in parent material, as vertisols require expanding clays to form.

3.1.2 Effect of Land Plants

The expansion of vascular plants during the Early and Middle Devonian resulted in fundamental changes in landscape stability and climate (Banks et al., 1985: Driese et al., 1997, Driese et al., 2000; Retallack, 1997; Algeo and Scheckler, 1998; Streel et al.,

2000; Beerling and Berner, 2005; Berner, 2006; Le Hir et al., 2011). Increasing height of arborescent plants during the Devonian resulted in deeper, more extensive rooting systems, which in turn enhanced mechanical and chemical (Driese et al.,

1997, 2000; Algeo and Scheckler, 1998; Beerling and Berner, 2005). Increased biologic weathering of parent material coupled with increased landscape stability allowed for greater amounts of time for pedogenesis to occur (Retallack, 1997; Algeo and Scheckler,

1998; Mintz et al., 2010).

Pedotype 1 contains shallow, millimeter-scale clay infilled roots and relatively deeper penetrating drab haloed root traces. Petrographic analysis shows that clay infilled roots occur throughout the entirety of all three Pedotype 1 profiles. Large arborescent plants are documented in older, correlative strata in New York (Driese et al., 1997; Mintz

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et al., 2010; Retallack, 2011) and therefore the absence of large roots in Pedotype 1 reflects the depositional environment rather than a lack of deep-rooting plants at this time. Hydromorphic features within Pedotype 1 and the abundance of gray and green overbank fines lacking pedogenic modification within the Irish Valley Member indicate waterlogged overbank environments which may have only been suitable for colonization by specialized plants. Macro-scale roots are rare in Pedotype 2, however, petrographic analysis reveals clay infilled roots throughout these profiles. Amalgamated, multistory channel complexes characteristic of the Sherman Creek Member indicate a relatively higher energy, dynamic depositional environment than the underlying Irish Valley

Member. Pedotype 2 paleosols contain Stage 1 to 2 carbonate accumulation, indicating these formed over a period of up to 5,000 years (Gile et al., 1966). Plants stabilizing overbank environments would promote such extended periods of pedogenesis in an otherwise high-energy, dynamic landscape. Larger, deeper penetrating drab haloed root traces are common in well-developed Pedotype 3 paleosols. Pedotype 3 paleosols cap well-defined fining-upward packages interpreted to be the result of lateral accretion of point bars. Colonization of flood plain environments by deep rooting plants promotes greater landscape stability allowing for the formation of well-developed vertisols of

Pedotype 3.

3.1.3 Landscape Gradient and Base Level

Increasing landscape gradient is coupled with generally greater depth to the water table. Since the Catskill Formation has been interpreted to represent a prograding, low- gradient alluvial plane, depth to the water table will increase up-section as depositional environments become increasingly source-proximal (Sevon, 1985; Woodrow, 1985;

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Slingerland et al., 2009). Accumulations of soluble bases, deeper rooting, and a lack of hydromorphic features up-section (from Pedotype 1 aqualfs to Pedotype 3 uderts) are interpreted to be the manifestation of increasing depth the water table. This trend is consistent with increased gradients in source-proximal DFS depositional environments.

Clay translocation is promoted by increased soil drainage, and therefore greater water table depths. Since all pedotypes show evidence of clay translocation, seasonal water table fluctuations or high-frequency paleoclimate and base level change may be overprinting the effect of topography on soil drainage. This overprinting is especially evident in Pedotype 1 aqualfs. Fourth and fifth order sedimentary cycles are documented by Cotter and Driese (1998) in the Sherman Creek and Irish Valley Members, respectively, and result in locally-well developed paleosols adjacent to incised-valley fills. Overall regression during the Late Devonian created accommodation space, allowing for progradation of the Catskill Formation into the Appalachian Basin

(Figure 17). This resulted in the vertical succession from poorly- to well-drained soil forming environments consistent with a prograding DFS (Hartley et al., 2013;

Weissmann et al., 2013; Trendell et al., 2013).

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Figure 17: Block diagrams illustrating the relationship between base level and progradation of DFS. A) Base level is at highstand and gradually falling. B) An increased rate of base level fall results in deposition of a tongue of the Clarks Ferry Member at Duncannon (DC). C) Base level continues to fall, resulting in the vertical succession preserved at Selinsgrove (SG) and Duncannon (DC).

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3.1.4 Parent Material

All pedotypes form on overbank and . However, grain size, sorting, and composition vary with stratigraphic position. Petrographic analysis of sandstones and paleosols indicates that parent material is largely composed of quartz and lithic fragments, with minor amounts of feldspars and micas. Feldspars show evidence of alteration to clay minerals when present. Micas are often fresh and unaltered, but are occasionally partially delaminated. The abundance of vertisols, and therefore presence of smectitic parent material, increases from Pedotype 1 to Pedotype 3. One of three paleosols described from the Irish Valley Member contained vertic features. When present in paleosols within the Irish Valley Member, vertic features are weakly expressed and suggest minimal input of smectitic material. Vertic features are absent in all paleosols described from the Sherman Creek Member. Vertic features are ubiquitous in all three paleosols described from the Duncannon Member, however, non-vertic paleosols are also present within this member (Figure 8, Figure 9). Vertic paleosols are designated as the pedotype for the Duncannon Member due to the abundance of vertic paleosols relative to non-vertic paleosols within this member. Intervals of vertic paleosols punctuated by non-vertic paleosols within the Duncannon Member can be attributed to more frequent episodic input of volcanic material that was hydrolyzed to smectite clays (Odom, 1984;

Terry et al., 2015; Ver Straeten et al., 2015). Smectitic parent material derived from the weathering of the Acadian Orogenic belt would result in exclusively vertic paleosols occurring within a particular stratigraphic interval, rather than vertisols occurring in punctuated intervals.

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3.1.5 Time of Pedogenesis

Time available for pedogenesis in alluvial depositional environments is a function of channel confinement and mobility (Kraus, 2002; Hartley et al., 2013). Pedotype 1 aqualfs and Pedotype 2 usterts have poorly to moderately expressed horizons in outcrop and discontinuous patches of well-developed plasmic microfabric, indicating these paleosols experienced a short period of pedogenesis. Thick, well defined horizons in outcrop and well-developed plasmic microfabric are present in Pedotype 3 uderts, indicating these soils formed over longer periods of time relative to Pedotypes 1 and 2.

Channel sandstones associated with Pedotype 1 are relatively small and occur as isolated bodies within overbank fines and as abundant sheet sandstones. The abundance of overbank material relative to channel material implies poor channel confinement

(Weissmann et al., 2013). Two of three paleosols described for Pedotype 1 are thin and show moderately to poorly developed horizons in outcrop. Unaltered to partially altered feldspars are present in these paleosols, which is consistent with short periods of pedogenic alteration. It is unlikely that increased exposure of crystalline source rocks through the denudation of the Acadian Orogenic belt would produce differing feldspar content in Pedotype 1 compared to Pedotypes 2 or 3, as parent material generally lacks feldspars throughout the section (Figure 14). Multistory, amalgamated channel sandstones are associated with Pedotype 2, suggesting a highly mobile channel belt which would have promoted erosion and alluvial aggradation in overbank environments and inhibited pedogenesis. Pedotype 2 ustepts are interpreted to have formed over a maximum of 5,000 years based on the presence of Stage 1 and 2 carbonate accumulation in all three paleosols described. Pedotype 3 uderts often cap well-defined fining-upward

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packages that are interpreted as the product of sinuous, meandering river deposits. Thick overbank facies are found in association with large, laterally extensive channel sandstones. Well-defined horizons in thick pedogenically modified overbank fines and less common sheet sandstones suggests greater channel confinement in the Duncannon

Member, and thus greater time for pedogenic modification in Pedotype 3 paleosols relative to Pedotypes 1 and 2.

The degree to which characteristics of subsurface horizons are expressed can be utilized to estimate time of pedogenesis (Birkeland, 1999). Bt and Bk horizons may reach steady state in 100 to 100,000 years. Variability in the time of formation results from other factors of soil formation. Pedotype 2 ustepts contains Btk horizons in their incipient stages of development, indicating a minimal time of pedogenesis. Since Pedotype 2 is a compound soil it is difficult to determine its age due to multiple periods of pedogenesis.

However, the presence of cambic (Bw) horizons within this paleosol indicate ages of 10 to 10,000 years. Pedotype 2 soils are likely younger than 10,000 years, as the cambic horizons are not fully expressed. Pedogenic carbonate morphology in this pedotype constrains a time of formation to approximately 5,000 to 10,000 years. Aqualfs of

Pedotype 1 contains well-expressed Bt horizons, which may indicate an age of 100 to

100,000 years for these paleosols. The abundances of overbank fines relative to channel sandstones within the Irish Valley Member indicates poor channel confinement and therefore high rates on floodplains. These conditions inhibit pedogenesis and therefore Pedotype 1 paleosols are likely 10,000 years old or younger. The base of all

Pedotype 3 uderts show the incipient stages of Bt horizon formation indicating a minimum time of formation of 1,000 years.

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3.2 Sandstone Provenance

Sandstone samples from all members of the Catskill Formation contain high percentages of quartz (average 57.51% to 70.79%) and lithic grains (average 25.73% to

36.89%) and generally lack feldspars (average 1.49% to 5.89%) (Table 9). Plotting relative percentages of quartz, feldspars, and lithic grains indicates a recycled orogenic terrane as the sediment source of the Catskill Formation based on the relationship between sandstone composition and tectonic setting developed by Dickinson and Suczek

(1979) (Figure 14). The scatter in point count data may be attributed to an evolving source terrane, as sedimentary cover is eroded exposing the crystalline core of the

Acadian Orogen. Although not evident in this dataset, Simonides (2014) documents this trend in correlative rocks in New York. Feldspar content decreases up-section which may suggest sandstone composition is the result of changing weathering conditions rather than changing source terrane lithologies. Further analysis is required to distinguish whether the range in composition presented here is the result of a changing source rock composition or paleoclimate, and therefore, weathering conditions.

Sandstones generally occur in fining-upward packages nested within the overall coarsening-upward trend throughout the Catskill Formation (Figure 16). Average sandstone grain size coarsens upward overall from 90.14 µm within the Irish Valley

Member, 173.86 µm within the Sherman Creek Member, 251.79 µm within the Clarks

Ferry Member and 232.25 µm within the Duncannon Member. Coarser grain size is coupled with larger channels up-section. The overall coarsening-upward nature of the

Catskill Formation indicates a general prograding trend through time. However, maximum grain size increases up-section within the Irish Valley and Duncannon

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Members while maximum grain size decreases up-section within the Sherman Creek and

Clarks Ferry Members. Coarsening upward trends in the Irish Valley and Duncannon

Members may reflect periods of increased sediment supply or progradation over shorter time scales. Conversely, fining upward trends within the Sherman Creek and Clarks Ferry

Members may reflect decreased sediment supply or aggradation on shorter time scales.

Sorting decreases up-section coupled with overall larger grain size and implies increasingly source-proximal depositional environments up-section (Figure 16). The

Clarks Ferry Member is anomalously coarse relative to the Irish Valley, Sherman Creek, and Duncannon Members and is interpreted to be a tongue of the coarser-grained source- proximal facies present in northeast Pennsylvania (Berg et al., 1993).

3.3 Up-section Variability in Channel Grain Size and Morphology

Channel sandstones within the Irish Valley Member occur as single-story bodies isolated in overbank fines (Figure 6 A). The abundance of overbank fines and sheet sandstones relative to channel sandstones reflects poorly channelized flow throughout much of this stratigraphic interval, characteristic of source-distal DFS deposits (Gibling,

2006; Nichols and Fisher, 2007; Weissmann et al., 2010; Hartley et al., 2010; Hartley et al., 2013; Weissmann et al, 2011; Weissmann et al, 2013). Grain size increases to fine to medium sand within the Sherman Creek Member and channels become larger, amalgamated, succession-dominated multistory bodies (Figure 6 C). This geometry indicates predominantly vertical rather than lateral accretion (Gibling, 2006) and is consistent with interpreted aggradation based an overall fining upward trend within this interval. The Sherman Creek Member is interpreted as a fine-grained braid plain similar to medial environments on the modern Okavango DFS (Hartley et al., 2013) based on

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alluvial architecture, grain size, and poorly defined fining upward packages often capped by ustepts of Pedotype 2. Exposures of channel sandstones within the Sherman Creek

Member at Cedar Run and Powys curve show a similar morphology and suggests this fluvial style is characteristic of the Sherman Creek Member and not locally restricted to central Pennsylvania. Sandstones within the Clarks Ferry Member occur as multistory, amalgamated channel bodies with prominent lateral accretion sets and are relatively larger than the sandstones within the underlying Sherman Creek Member. Grain size increases abruptly at the base of the Clarks Ferry Member to medium to coarse sand before fining upward to very fine to fine sand. These characteristics are interpreted as source-proximal meandering feeder channels on DFS (Gibling, 2006; Nichols and Fisher,

2007). Sandstones within the Duncannon Member are texturally similar to sandstones within the Clarks Ferry Member, however, sandstones within the Duncannon Member tend to occur as large-scale (100’s of meters wide), massive bodies at the base of well-defined fining upward packages. Exposures of channel sandstones within the

Duncannon Member at Red Hill show a similar size and morphology to channel sandstones exposed at Duncannon, suggesting large, laterally extensive channels are not restricted to central Pennsylvania. These channels are also interpreted as meandering feeder channels in a source-proximal DFS environment. Since the Duncannon Member occurs stratigraphically above the Clarks Ferry Member (and its lateral equivalents), the transition from the coarser-grained Clarks Ferry Member sandstones to the relatively finer-grained Duncannon Member sandstones is attributed to increased landscape gradient as a result of tectonic activity. Higher-order base level fluctuations result in a

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basinward shift of depositional environments resulting in deposition of a tounge of the

Clarks Ferry Member at Duncannon (Figure 17).

3.4 Flow Competency Analysis

The shear stress necessary to entrain the D90 grain size is used as a proxy for paleoflow competency. At Selinsgrove, shear stress varies between 0.05 and 0.15 Pa within the Irish Valley Member (Figure 18 A). Low shear stress values indicate lesser flow competency and supports the interpretation of poorly channelized flow occurring in relatively smaller channels within this stratigraphic interval. Approximately 50 m above the base of the Sherman Creek Member, shear stress increases to a maximum of 0.28 Pa and then gradually decreasing to between 0.11 and 0.19 Pa (Figure 18 A). Shear stress within the Sherman Creek Member increases relative to the Irish Valley Member and gradually decreases up-section. Greater overall shear stress indicates greater flow competency, which is consistent with relatively larger channel sizes within the Sherman

Creek Member. The overall decrease in shear stress through the Sherman Creek Member is interpreted to be the result of increased gradient due to pulses of tectonic activity followed by quiescence (Figure 18 A). The trends in shear stress correspond to finer grain size and increased sorting up-section through the Sherman Creek Member, which can also be attributed increased gradient resulting from tectonism followed by quiescence.

Within the Clarks Ferry Member, shear stress varies from 0.09 to 0.74 Pa and generally decreases up-section (Figure 18 B). Shear stress is not a reliable indicator of flow competency within the Clarks Ferry Member, as this calculation was restricted to sand- size clasts. The Clarks Ferry Member is occasionally conglomeratic, especially at the base of the member, which indicates greater flow competency relative to the underlying

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Sherman Creek Member. Although shear stress does not vary considerably from the underlying Sherman Creek Member, channels are larger and coarser grained within the

Clarks Ferry Member. The overall decrease in shear stress up-section is similar to the trend within the Sherman Creek Member and is also attributed to increased gradient due to tectonism, and is consistent with fining grain size and increased sorting up-section through the Clarks Ferry Member. The Duncannon Member shows an overall increase in shear stress, from 0.12 – 0.25 Pa at the base of the section to 0.28 – 0.48 Pa at the top of the section (Figure 18 B). Increasing shear stress throughout the Duncannon Member corresponds with channels becoming larger and coarser grained as depositional environments become increasingly source proximal.

Grain size of the bed surface will change with increasing distance from the source area. Therefore, the applicable equation for calculating shear stress will change in a section which represents large lateral distances. Empirically derived constants of

Andrews (1983) were developed from streams with a bed surface of coarse sand to pebbles. These constants yield minimum shear stress values throughout the section and represent averaged surface roughness from source proximal to distal environments.

Constants given by Komar, 1987, and Komar, 1989 are derived from stream beds with a pebble grain size. These constants yield maximum shear stresses throughout the section and are an upper limit of realistic bed roughness and are likely more applicable to the

Clarks Ferry Member, as this member is anomalously coarse grained in comparison to the

Irish Valley, Sherman Creek, and Duncannon Members.

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Figure 18: Shear stress trends through the Catskill Formation. A) Calculated shear stress from the Sherman Creek and Irish Valley Members. Note an overall increase in shear stress up-section. B) Calculated shear stress from the top of the Sherman Creek Member, Clarks Ferry Member, and Duncannon Member. Note an overall increase in shear stress up-section and the spike at the contact between the Sherman Creek and Clarks Ferry Members.

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3.5 A Note on Diagenesis

Chlorite is abundant in most sandstone and paleosol thin sections and also appears the in x-ray diffractograms of Pedotypes 1 and 2. Chlorite is present in thin sections from

Pedotype 3 and the absence of chlorite peaks in the XRD spectra from Pedotype 3 is attributed to a lesser abundance of chlorite present in these paleosols. X-ray diffractograms (Figure 13) show that illite is the primary clay mineral present in all pedotypes. However, the presence of vertic features, which are the result of seasonal expansion and contraction of swelling clays, suggests illitization of smectite clays rather than the presence of illite as original parent material. Given that original clays have been diagenetically altered, clay mineralogy cannot be utilized to infer paleoclimate during pedogenesis. Additionally, climofunctions (e.g. Sheldon et al., 2002; Sheldon and Tabor,

2009; Nordt and Driese, 2010) based on paleosol bulk geochemistry cannot be applied reliably to geochemical data derived from these paleosols as due to the addition of Mg and K from diagenetic chlorite and illite, respectively.

Adjusting rooting depths for burial compactions yields original rooting depths of

43 cm in Pedotype 1, 148 cm in Pedotype 2, and 250 cm in Pedotype 3. Accounting for compaction yields deeper rooting depths, however, the trend of increasing rooting depth up-section does not change in comparison to observed rooting depths for this study. Since compaction is dependent on initial soil properties specific to soil orders, compaction should be accounted for when utilizing rooting depth as a proxy for depth to the water table. Paleoclimatic proxies which rely on depth to calcic horizons (e.g. Retallack, 2005) must account for compaction of original profile thickness. Retallack (2009) shows minimal variability in depth to calcic horizon based paleoprecipitation estimates from

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paleosols at Red Hill. However, burial depths used in this decompaction equation must be site specific, as the Catskill Formation in the central Valley and Ridge Province

(Selinsgrove and Duncannon, PA) has experienced greater burial depths than the Catskill

Formation in the Appalachian Plateau (Red Hill).

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CHAPTER 4. CONCLUSIONS

I interpret deposition of the Catskill Formation to be the product of distributive fluvial system processes (DFS) based on vertical trends in paleosol development and drainage, channel sandstone morphology and grain size, and alluvial architecture

(Table 10). Paleosols become progressively well-drained and well-developed up-section through the Catskill Formation. At least three distinct pedotypes are identified within the

Irish Valley, Sherman Creek, and Duncannon Members, each representing a distinct depositional environment. Pedotype 1 is associated with the Irish Valley Member and consists of aqualfs with moderately developed horizons, hydromorphic features, and shallow rooting. Illuviated clay material indicates increased soil drainage potential due to periodic increased water table depth as a result of base level fluctuations. Pedotype 1 is associated with relatively small, asymmetric, single-story channel sandstones which are isolated in gray to greenish-gray overbank fines. The abundance of overbank fines and sheet sandstones relative to channel sandstones suggests poorly-channelized flow and perennially inundated floodplain environments within the Irish Valley Member.

Pedotype 2 is associated with the Sherman Creek Member and consists of ustepts with weak to moderately developed horizons, pedogenic carbonate accumulations, and illuviated clays. The transition from hydromorphic paleosols with shallow roots to calcareous paleosols with relatively deeper penetrating roots is contradictory to other paleoclimate proxies, which indicate increasing paleoprecipitation during the Late

Devonian. Varibility in paleosol drainage is therefore attributed to soil forming environments becoming increasingly well-drained up-section rather than recording an increase in paleoprecipitation. Amalgamated, succession-dominated multistory channel

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sandstones in poorly-defined fining-upward packages are characteristic of the Sherman

Creek Member and are interpreted as fine-grained braided fluvial deposits characteristic of medial DFS environments. Channel avulsion and vertical aggradation on floodplains inhibits soil formation, resulting in Pedotype 2 paleosols being generally thin and weakly-developed due to insufficient time for pedogenesis. Pedotype 3 is associated with the Duncannon Member and consists of uderts with well-developed horizons, abundant pedogenic slick planes, and deeper rooting relative to both Pedotypes 1 and 2. Pedotype 3 often caps well-defined fining-upward packages interpreted as meandering fluvial deposits of feeder channels on a DFS. Thick, laterally extensive, asymmetric single-story channel sandstones indicate channelized flow, which promotes extended periods of pedogenesis in overbank environments. A pedotype is not defined for the Clarks Ferry

Member due to an overall lack in pedogenically modified mudstone facies within this member. Increased paleosol development, lack of hydromorphic features, and increased rooting depth from Pedotype 1 aqualfs to Pedotype 3 uderts is consistent with soil development trends on modern DFS (Hartley et al., 2013; Weissmann et al., 2013).

Up-section variability in channels sandstone morphology is also consistent with deposition by DFS processes. Channel sandstones within the Irish Valley Member consist of very fine to fine grained litharenites. These sandstones occur as small, isolated single story bodies within overbank fines. Within the Sherman Creek Member, sandstones are poor to moderately-sorted fine-grained litharenites which occur as succession dominated, amalgamated multistory channel complexes. When present, channel sandstones within the Clarks Ferry Member are characterized by fine to medium grained, occasionally conglomeratic sandstones with prominent laterally accretion foresets and are larger

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relative to channels within the underlying Sherman Creek Member. These sandstones are interpreted as coarse grained, source-proximal meandering feeder channels on a DFS.

The presence of the Clarks Ferry Member below the relatively finer-grained Duncannon

Member is attributed to eustatic base level fluctuations. Channel sandstones within the

Duncannon Member are fine to medium grained laterally extensive bodies and are interpreted as meandering feeder channels on a DFS. Shear stress trends are consistent with qualitative channel size and grain size observations. Mineralogical composition of channel sandstones varies up-section and is attributed to evolving source terrane lithologies during the weathering of the Acadian Orogenic belt rather than depositional fractionation.

Although these results are consistent with the DFS depositional model, the limited lateral extent of this study cannot unequivocally allow for the interpretation of basin-scale depositional processes. However, this study documents trends in paleosols and channel sandstones which are consistent with the DFS model and demonstrates the utility of pedofacies and alluvial architecture analysis in identifying DFS in the rock record.

Identifying DFS in the rock record has implications for paleosol based paleoclimatic interpretations, as paleosols associated with prograding DFS show a drying-upward trend as a result of hydrologic variability rather than increasing aridity. Recognizing DFS in the rock record also has implications for basin analysis and characterizing fluvial hydrocarbon reservoirs, as the geometry and scale of lithofacies distributions associated with DFS are fundamentally different from tributary systems.

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Table 10: Characteristics of Distributive (DFS) versus Tributary (TFS) Fluvial Systems in the Catskill Formation

Characteristic DFS TFS

Soil drainage  Soil maturity  Channel size  Flow competency  Alluvial architecture  Sandstone geometry  Sandstone maturity Inconclusive

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APPENDIX A. STUDY SITE LOCATIONS AND DESCRIPTIONS

Locality Latitude Longitude Notes Road cut only accessable Selinsgrove 40.76340 -76.86536 from southbound side of US 11/15 Private property. Must obtain permission from Duncannon 40.38857 -77.01884 CSX Corporation prior to access. Road cut along PA Rt. Cedar Run 41.53434 -77.42699 414 parallel to Pine Creek Road cut only accessable Powys Curve 41.34236 -77.08953 from southbound side of US 11/15. Road cut along PA Rt. 120 parallel to West Red Hill 41.34478 -77.68178 Branch Susquehanna River

85

APPENDIX B. GRAIN SIZE DISTRIBUTIONS

DC_CHSST_CO_20

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 200.0 400.0 600.0 800.0 1000.0 Grain Size (um)

86

DC_CHSST_CO_680

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 400.0 800.0 1200.0 1600.0 2000.0 Grain Size (um)

87

DC_CHSST_CO_2800

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 90.0 180.0 270.0 360.0 450.0 Grain Size (um)

88

DC_CHSST_CO_3191

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 200.0 400.0 600.0 800.0 1000.0 Grain Size (um)

89

DC_CHSST_CO_3200

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 50.0 120.0 190.0 260.0 330.0 400.0 Grain Size (um)

90

DC_CHSST_CO_4100

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 120.0 240.0 360.0 480.0 600.0 Grain Size (um)

91

DC_CHSST_CO_6800

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 50.0 100.0 150.0 200.0 250.0 300.0 Grain Size (um)

92

DC_CHSST_CO_7000

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 50.0 100.0 150.0 200.0 250.0 Grain Size (um)

93

DC_CHSST_CO_10003

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 100.0 200.0 300.0 400.0 500.0 Grain Size (um)

94

DC_CHSST_CO_14929

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 400.0 800.0 1200.0 1600.0 2000.0 Grain Size (um)

95

DC_CHSST_CO_18830

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 240.0 480.0 720.0 960.0 1200.0 Grain Size (um)

96

DC_CHSST_CO_19900

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 120.0 240.0 360.0 480.0 600.0 Grain Size (um)

97

DC_CHSST_CO_29100

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 200.0 400.0 600.0 800.0 1000.0 Grain Size (um)

98

SG_CHSST_CO_14500

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 100.0 200.0 300.0 400.0 500.0 Grain Size (um)

99

SG-CHSST-CO-23900

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 50.0 100.0 150.0 200.0 250.0 Grain Size (um)

100

SG_CHSST_CO_24700

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 90.0 180.0 270.0 360.0 450.0 Grain Size (um)

101

SG-CHSST-CO-34500

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 60.0 120.0 180.0 240.0 300.0 Grain Size (um)

102

SG_CHSST_CO_44600

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 60.0 120.0 180.0 240.0 300.0 Grain Size (um)

103

SG_CHSST_CO_61800

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 80.0 160.0 240.0 320.0 400.0 Grain Size (um)

104

SG_CHSST_IVF1_C0_5800

100.0

80.0

60.0

40.0

Frequency (%)

20.0

0.0 0.0 140.0 280.0 420.0 560.0 700.0 Grain Size (um)

105