The Pennsylvania State University

The Graduate School

Department of Geosciences

LATE MIOCENE EROSION AND EVOLUTION OF TOPOGRAPHY ALONG THE

WESTERN SLOPE OF THE COLORADO ROCKIES

A Thesis in

Geosciences

by

Russell H. Rosenberg

 2013 Russell H. Rosenberg

Submitted in Partial Fulfillment of the Requirements for the Degree of

Master of Science

May 2013

The thesis of Russell H. Rosenberg was reviewed and approved* by the following:

Eric Kirby Associate Professor of Geosciences Thesis Advisor

Kevin Furlong Professor of Geosciences

Rudy Slingerland Professor of Geology

Chris Marone Professor of Geosciences Associate Department Head of Graduate Programs

*Signatures are on file in the Graduate School

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ABSTRACT

It is increasingly apparent that dynamic effects associated with changes in mantle flow and buoyancy can influence the evolution of surface topography. In the Rocky Mountain province of the western United States, recent high-resolution seismic imaging of the crust and upper mantle reveals intriguing correlations between mantle velocity anomalies and regions of high topography. To determine whether these regions of low-seismic wavespeed are associated with recent changes in buoyancy structure of the lithosphere, I explore the relationship between the longitudinal profiles of tributaries draining the western slope of the northern Colorado Rockies and the history of Late Cenozoic fluvial incision and exhumation across this region. Major tributaries of the upper Colorado River, including the Gunnison and Dolores Rivers, that drain high topography in central and western Colorado overlie upper mantle with slow seismic wave velocities; these drainages exhibit relatively steep longitudinal profiles (normalized for differences in drainage area and discharge) and are associated with ~1000-1500 m of incision over the past 10 Ma. In contrast, tributaries of the Green River that drain the western slope in northern Colorado (White, Yampa, and Little Snake Rivers) overlie mantle of progressively lower seismic wave velocities to the north. River profiles in northern Colorado are two to three times less steep along reaches within comparable bedrock lithologies. New 40Ar/39Ar ages on basalt flows capping the Tertiary Browns Park Formation in this region range in age from ~11-6 Ma, and provide local datums from which I reconstruct ~500-900 m of incision along tributaries of the

Green River. The correspondence of steep river profiles in regions of greater incision and lower gradient profiles in regions of less incision suggests that the fluvial systems are dynamically adjusting to an external forcing. Moreover, spatial differences in the pattern and magnitude of incision are not readily explained by a putative increase in erosivity associated with late Cenozoic climate change. Rather, fluvial incision appears to reflect relative base level fall along the

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western slope. Given the correspondence of steep channels, deep incision and regions of low seismic velocity mantle, I suggest that differential rock uplift driven, in part, by differences in the buoyancy and/or convective flow of the mantle beneath western Colorado is the likely driver for

Neogene incision.

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TABLE OF CONTENTS

LIST OF FIGURES ...... vii

LIST OF TABLES ...... x

ACKNOWLEDGEMENTS ...... xi

1 Introduction ...... 1

2 Background ...... 4

2.1 Support of High Topography in the Colorado Rockies ...... 5 2.2 Timing, Magnitude and Rates of Incision along the Western Slope ...... 8 2.2.1 Colorado and Gunnison Rivers ...... 10 2.2.2 White, Yampa, and Little Snake Rivers ...... 17 2.3 Analysis of River Longitudinal Profiles ...... 22 2.4 Controls on Channel Profile Form along the Western Slope of the Rocky Mountains ...... 24

3 Analysis of Channel Profiles along the Western Slope ...... 26

3.1 Measuring Normalized Channel Steepness Index ...... 26 3.2 Evaluating the Effect of Substrate Lithology on Channel Steepness ...... 28 3.3 Evaluating Relationships between Discharge and Drainage Area ...... 31

4 Results of Channel Profile Analysis ...... 32

4.1 Relationships Between Discharge and Drainage Area ...... 32 4.2 Colorado, Gunnison, and Dolores Rivers ...... 32 4.3 Yampa, White, and Little Snake Rivers ...... 36

5 New Constraints on Late Miocene Exhumation ...... 41

5.1 Reconstructing Exhumation in the Elkhead Mountains ...... 44 5.1.1 Battle Mountain, Squaw Mountain, and Bible Back Mountain ...... 46 5.1.2 Black Mountain and Mt. Welba ...... 47 5.2 Reconstructing Exhumation in the Flattops ...... 49 5.3 Reconstructing Exhumation in the Yampa River Valley ...... 51 5.3.1 Woodchuck Hill ...... 52 5.3.2 Lone Spring Butte ...... 52 5.4 Summary of Constraints on the Timing and Magnitude of Incision ...... 55

6 Discussion ...... 60

6.1 Extensional Faulting in the Sand Wash Basin ...... 60 6.2 Regional Correspondence between Late Miocene Incision and Channel Steepness ...... 62

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7 Potential Drivers of Late Miocene Incision ...... 65

7.1 Enhanced Fluvial Incision in the Late Miocene ...... 65 7.2 Base-level Fall and Transient Incision Associated with Basin Integration ...... 67 7.3 Differential Rock Uplift and Tilting along the Western Slope ...... 69

8 Conclusions ...... 72

References ...... 74

Appendix: 40Ar/39Ar Analytical Methods and Results ...... 89

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LIST OF FIGURES

Figure 1: Modern topography of the Rocky Mountain physiographic province (left pannel) compared to differential P-wave velocity at 100 km depth (right pannel). Geographic points for reference: GJ--Grand Junction, CO; R--Rifle, CO; SB--Steamboat Springs, CO; NP--North Park, CO; SP--South Park, CO; GM--Grand Mesa; BC--Book Cliffs; FT--Flat Tops. Tomographic data from Schmandt and Humphreys (2010)...... 7

Figure 2: Interpolated normalized channel steepness (ksn) for the entire Colorado River watershed; values calculated with 10 km channel segments and a reference concavity of 0.45 (modified from Karlstrom et al., 2012). The study area for this work, and the extent of Figure 1 (above) is shown by black inset...... 7

Figure 3: Simplified geologic map showing the locations of previously dated landforms which provide constraints on the timing and magnitude of incicision along the Colorado River (modified from Green, 1992; Tweto, 1979). The location of evaporite collapse centers along the Colorado River (from Kunk et al., 2002) are also shwon. Date for prveiously published incision markers along the Colorado River are given in Table 1...... 11

Figure 4: Simplified geologic map showing the extent of the Browns Park Formation (modified from Green, 1992; Tweto, 1979; Green and Drouillard, 1994; Love and Christiansen, 1985; Hintze et al., 2000; Hintze, 1980). The extent of detailed study areas for this work (Figure 11: Elkhead Mountains, Figure: 13 Flat Tops, Figure 14: Yampa River Valley) are shown above by white boxes. Locaities constraining the age of the Browns Park Formation (see text): 1 – Dead Mexican Park; 2 – west bank of Little Snake River; 3 – City Mountain; 4 – Vermillion Creek...... 18

Figure 5: Colorado River long profile with reservoirs removed (blue profile). Reserviors were removed by linear interpolation of the channel profile above and below the reservoir. The segments in red, show areas which were removed. Grandby reservoir is highlighted to contrast with natural steep reaches (e.g. Gore Canyon) which remain in the final profile...... 27

Figure 6: Colorado River long profile with 10 km spaced bins of normalized channel steepness. Normalized channel steepness values were determined by an automated regression of the channel slope every 0.5 km. The average ksn value within each 10 km bin determines the ‘local’ normalized channel steepness value and the standard deviation for each 10 km bin determines the error shown shown for each data point. Colorado River long profile shown with dams removed...... 27

Figure 7: Simplified geologic map showing dominate bedrock lithologies for the study area, quaternary deposits removed (modified from Green, 1992; Tweto, 1979; Green and Drouillard, 1994; Love and Christiansen, 1985; Hintze et al., 2000; Hintze, 1980). Major rivers labeled (north to south): LS--Little Snake River, Y--Yampa River, W--White River, Gr--Green River, C--Colorado River, Gn--Gunnison River, D--Dolorers River...... 30

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Figure 8: Discharge-drainage area relationships of study rivers. Discharge data derived from gridded precipitation model (http://www.prism.oregonstate.edu). Discharge data represent the 30-year average of annual average precipitation models...... 33

Figure 9: Long profiles of rivers from north (top of page) to south (bottom of page) with 10 km spaced bins of normalized channel steepness and color coded bedrock lithology. Error bars show standard deviation of local ksn...... 34/37

Figure 10A: Schematic representation of geologic relationships between basalt flows and the Browns Park Formation (Tbp) in the Elkhead Mountains (modifed from Buffler, 2003). Basalt flows capping the Browns Park Formation provide an estimate of local relief generation. Lithology: JTr – Jurassic/Triassic undifferentiated, Kd – Dakota Sandstone, Km – Mancos Shale, Kmv – Mesa Verde Group, Kls – Lewis Shale, Tbpc – basal Browns Park conglomerate, Tbp – upper Browns Park Formation, Tv – Tertiary volcanics...... 42

Figure 10B: Field relationships between basalt flows, the Browns Park Formation, and the Little Snake River in the Elkhead Mountains (photo: Russell Rosenberg)...... 42

Figure 11: Simplified geologic map of the Elkhead Mountains (modified from Green, 1992; Tweto, 1979). References for ages: 1 this study; 2 Synder, 1980. Major faults which do involve Miocene sediments are dashed...... 45

Figure 12: Schematic cross-section from the White River Plateau and southern part of the Flat Tops on the southwest towards the Yampa River Valley and southern Park Range on the northeast (modifed from Larson et al., 1975). Lithology: Tbpc – basal Browns Park conglomerate, Tbp – upper Browns Park Formation, Tv – Tertiary volcanics (multiple flows)...... 49

Figure 13: Simplified geologic map of the Flat Tops (modified from Green, 1992; Tweto, 1979). References for ages: 1 Kunk et al., 2002; 2 Larson et al., 1975. *Sugar Loaf Mountain ages range from 13.45 +/- 0.16 Ma to 15.57 +/- 0.09 Ma (Kunk et al., 2002). Quaternary deposits are largely coarse debris and landslides...... 50

Figure 14: Simplified geologic map of the Yampa River Valley (modified from Green, 1992; Tweto, 1979). Crowner deposits shown as dashed Tertiary volcanics. References for ages: 1 this study; 2 Izett, 1975...... 53

Figure 15: Ages and magnitude of local incision at selected paleo-land-surface markers with errors in age calculations shown. Values next to each marker represent the time- average incision rate (m/Ma) calculated from each marker to modern river elevation. These plots combine new and previously published data. Grass Mesa (10) is unpublished data from Maureen Berlin’s PhD thesis (Berlin, 2009). Along the Colorado River, only markers associated with fluvial deposits were included in the analysis. Localities and age references: (1)Battle Mountain/Squaw Mountain (this study); (2)Lava Ck B ash at Baggs, WY (Aslan, 2010); (3)Lost Lakes/Sable Pt (Larson et al., 1975); (4)Lava Ck B ash at Meeker, CO (Aslan, 2010); (5)Orno Peak/Flat Top Mountain (Larson et al., 1975); (6)Lone Spring Butte (this study); (7)Woodchuck Hill (this study); (8)Grand Mesa (Kunk

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et al., 2002; Aslan, 2010; Cole, 2010); (9)Spruce Ridge (Kunk et al., 2002), (10)Grass Mesa (Berlin, 2009), (11)Lava Ck B ash at Dotsero, CO (Lidke et al., 2002)...... 56

Figure 16: New and previously published constraints on the magnitude of exhumation (in meters) along the western flank of the Colorado Rocky Mountains within the last 6 - 12 Ma. References for exhumation values (superscript numbers also correspond to information provided in Table 4): 1,2,3,4,5,6,7,8,16 this study; 9,10 this study; Larson et al., 1975; 11,12,13 Kunk et al., 2002; 14,15 Berlin 2008, 2009; 17 Kunk et al., 2002; Aslan et al., 2010; Cole, 2010. Majork knickpoints shown...... 58

Figure 17: By river comparisson of normalized steepness within varying lithologies. Study rivers are shown from north (left) to south (right). Lithologies examined are grouped in Tertiary sandstones and shales; corresponding formations are shown above. Some reaches were excluded from this analysis; within the Flat Tops in the headwaters of the Yampa River and White River, a short reach along the Little Snake River just downstream of a vertical step knickpoint, and the Yampa River through Dinosaur Canyon (see discussion). Dolores River not shown due to absence of similar lithologies...... 63

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LIST OF TABLES

Table 1: Previously published age constraints on Colorado River incision ...... 12

Table 2: Channel profile regressions ...... 38

Table 3: New 40Ar/39Ar ages determined in this study ...... 44

Table 4: Supplemental information for Figure 16; Magnitude of Incision ca. 6-12 Ma ...... 59

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ACKNOWLEDGEMENTS

I would like to thank my advisor, Eric Kirby, for his guidance on all facets of this thesis.

His comments on my work provided were invaluable and, without which, this thesis would not have been possible. Additionally, I would like to thank my committee members, Kevin Furlong and Rudy Slingerland, for offering many thoughtful suggestions on how best to approach specific problems which undoubtedly improved this work. I would also like to thank all of my colleagues and fellow graduate students who lent me their time, support and/or expertise which helped me complete much of the work that went into this thesis. I would like to thank Christine Regalla for many provocative scientific conversations, Kristen Morell for valuable advice on stream profiling, Cassi Knight for a very good sense of color, Andres Aslan for sharing his knowledge of the geology of northern Colorado and his time with me out in the field, Matt Heizler for sharing his lab and assistance determining 40Ar/39Ar ages for this study.

In addition to the thoughtful scientific guidance offered by the individuals mentioned above I would like extend a special debt of gratitude to my fellow interpretive rangers on the

North Rim of the Grand Canyon and to all of the wonderful visitors to the park who asked wonderful and enthusiastic questions of me. You have all reminded me why I chose to become a geologist in the first place; to tell a wonderful story. Thank you for being so eager to learn this story and for reminding me that, “nothing great was ever accomplished without enthusiasm.”

Finally, I would like to thank my family and close friends for their support and patience.

I know that a simple thank you cannot suffice for all that you have given supporting me.

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1 Introduction

One of the outstanding tectonic questions in western North America regards the development and support of high topography in the Rocky Mountains. It has long been recognized that there are correlations among high topography (Gregory and Chase, 1994), low seismic velocity mantle (e.g., Grand, 1994; Schmandt and Humphreys, 2010), high heat flow

(Sass et al., 1971), relatively thin crust (Sheehan et al., 1995), and volcanism. Although these correlations suggest a role for mantle support of high topography (Sheehan et al., 1995), it is less well understood when mantle buoyancy and consequent high topography developed. Widespread exposures of Cretaceous marine rocks across the Colorado Plateau and throughout portions of the

Rocky Mountains attest to approximately two kilometers of elevation increase since that time

(e.g., Pederson, 2002a). However, a more precise history of when the high topography of the western US developed in space and time remains uncertain.

Several competing models for the growth of high elevation beneath the Rocky Mountain-

Colorado Plateau region exist. Although deformation was widespread throughout the region during the Laramide Orogeny (Burchfiel et al., 1992) displacement on thick-skinned structures was generally limited (e.g., Erslev, 1993) and thus probably insufficient to have generated broad regional high elevations. Most alternative explanations involve changes to the buoyancy of the mantle beneath the western US. Hydration of the mantle during Laramide low-angle subduction

(Lipman et al., 1971; Coney and Reynolds, 1977) is thought to have contributed to both the low- density and fertility of the lithospheric mantle (Humphreys, 2003) and may have led to support of regional elevations in the early Tertiary. Thermal re-equilibration following removal and/or fragmentation of the Laramide slab may have driven additional surface uplift of the Colorado

Plateau during middle Tertiary time (Roy et al., 2004; 2009). Finally, changes in the mantle flow field (Moucha et al., 2008; van Wijk et al., 2010) and/or local changes in lithospheric structure

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(e.g., Levander et al., 2011) may have influenced regional elevations during Neogene time

(Karlstrom et al., 2012).

Many of the arguments in favor of a Neogene component of surface uplift in the Rocky

Mountains and Colorado Plateau find support in the timing and patterns of fluvial incision. It has long been recognized that fluvial incision and canyon cutting across the Colorado Plateau and along the western slope of the Rocky Mountains was extensive during late Miocene time (e.g.,

Hunt, 1956; Larson et al., 1975; Izett, 1975; Pederson et al., 2002). Recently, evidence is emerging that portions of the southern Colorado Plateau had a more complicated exhumation history that involved the development of relief in Oligocene – early Miocene time (Flowers et al.,

2008; Cather et al., 2008), and ongoing debate centers on the timing of when the present-day drainage system along the Colorado River was established (c.f., Karlstrom et al., 2012; Flowers et al., 2012).

Within the Rocky Mountains, however, the onset of fluvial incision appears to coincide with the cessation of late Tertiary deposition in intermontane basins (Larson et al., 1975; Buffler,

2003; McMillan et al., 2006). Along the eastern flank of the range, incision post-dates deposition of the ca. 18 - 6 Ma Ogallala Formation (McMillan et al., 2002; 2006). Notably, reconstruction of paleo-transport gradients (McMillan et al., 2002; Duller et al., 2012) in these rocks has been used to argue that long-wavelength tilting exceeds that expected for a simple isostatic response to exhumation (e.g., Leonard, 2002) and that tilting must have been, in part, driven by surface uplift within the Rockies (McMillan et al., 2002; Duller et al., 2012). Along the western slope of the range, fluvial incision appears to have initiated around the same time, ~10-7 Ma (Kunk et al.,

2002; Berlin et al., 2007; Berlin, 2009; Aslan et al., 2008; 2010), but the mechanism driving incision remains uncertain.

In this study, I seek to test whether the incision history of drainages along the western slope of the Colorado Rockies reflects differential rock uplift between the range and the Colorado

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Plateau. Recent analyses of the longitudinal profiles of major rivers draining this region highlight spatial differences in the steepness of these channels (Karlstrom et al., 2012; Pederson and

Tressler, 2012); tributaries of the Green River in northern Colorado (Little Snake, Yampa and

White Rivers) are less steep than the Colorado River and its tributaries to the south. It has been previously suggested that differences in the longitudinal profiles of these fluvial systems may be related to differential rock uplift associated with development of low-velocity mantle beneath central Colorado (the so-called ‘Aspen anomaly’ - Karlstrom et al., 2012). In order to evaluate this hypothesis, I combine a detailed analysis of river longitudinal profiles and their relationship to substrate lithology with geologic constraints on the timing and magnitude of fluvial incision. I develop new constraints for fluvial incision and erosion along the White, Yampa, and Little

Snake rivers through 40Ar/39Ar dating of a series of late Tertiary volcanic flows that cap the upper

Brown’s Park Formation. Collectively, these observations reveal spatial patterns in both channel steepness and in the magnitude of post-10 Ma incision and exhumation that are used to argue for late Cenozoic differential rock uplift along the flank of the range.

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2 Background

The absence of a thick crustal root beneath the high topography of the Colorado Rocky

Mountains implies a role for isostatic and, perhaps, dynamic support from the upper mantle

(Sheehan et al., 1995; Lee and Grand, 1996; Gilbert and Sheehan, 2004; Karlstrom et al., 2012;

Hansen et al., in review). Although spatial correlations between the region of highest topography, low seismic velocity mantle (Schmandt and Humphreys, 2010), and Neogene volcanism and fluvial incision (Karlstrom et al., 2012; Aslan et al., 2010) suggest the possibility of late Cenozoic rock uplift (Karlstrom et al., 2012), the difficulty of deconvolving the complicated history of

Laramide tectonism, possible attainment of high elevations in the early Tertiary (Gregory and

Chase, 1992), Oligocene development of the Rio Grande Rift (Landman and Flowers, 2012), and the possibility of climatically enhanced fluvial incision during the late Cenozoic (Pelletier, 2009;

Wobus et al., 2010) has made this interpretation controversial. This study is motivated by the recognition that, to first order, rivers draining the high topography of central Colorado are steeper than those draining the northern flanks of the range (Karlstrom et al., 2012). This observation does not appear to reflect differences in discharge among the fluvial systems (Pederson and

Tressler, 2012), and thus, must be related to some combination of differences in lithologic resistance, history of drainage integration, and/or differential rock uplift. The history of incision along the Colorado River is relatively well understood (e.g., Larsen et al., 1975; Kunk et al.,

2002, Aslan et al., 2010), so that the primary goal of this study is to combine careful analysis of the controls on fluvial longitudinal profiles with new estimates on the magnitude of fluvial incision along the White, Yampa, and Little Snake Rivers to help elucidate whether climate change, drainage integration, and/or rock uplift best explains the spatial patterns observed in topography, seismic velocity, the record of incision, and channel steepness along the western flank of the Rockies.

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2.1 Support of High Topography in the Colorado Rockies

There is long-standing recognition that crustal thicknesses beneath the Colorado Rockies are insufficient, compared to the adjacent Colorado Plateau and Great Plains, to locally support high topography (e.g., Sheehan et al., 1995; Lee and Grand, 1996; Gilbert and Sheehan, 2004).

This fact implies a role for mantle support in generating and maintaining high topography in the

Colorado Rockies. The modern average elevation of the Colorado Rocky Mountains is approximately 3.2 km (Karlstrom et al., 2012). Simple isostatic models predict that support of an average elevation of ~3200 m would require a crustal thickness of approximately 50 km (e.g.

Coblentz et al., 2011). However, average crustal thicknesses range from ~42-45 km beneath the

Colorado Rocky Mountains (e.g., Coblentz et al., 2011; Karlstrom et al., 2012; Hansen et al., in review). These observations suggest that the crustal thickness underlying the Colorado Rockies is not sufficient to support the overall high elevations in the region. To a simple approximation,

~500-1000 m of elevation within the Colorado Rockies cannot be supported by the present-day crustal thickness.

Despite this recognition, the precise timing and mechanisms responsible for the support of high topography in the Colorado Rockies remain unclear. Recently, it has been proposed that high elevations in the western US are supported by up to ~1000 m of long-wavelength dynamic topography beneath the Colorado Plateau that tapers eastward to a few hundred meters beneath the Rocky Mountain province (Moucha et al., 2009; Liu and Gurnis, 2010; Forte et al., 2010). It has also been suggested that small-scale mantle convection along the western margin of the

Colorado Plateau (e.g., van Wijk et al., 2010; Coblentz et al., 2011) and/or lithospheric delamination beneath the central Colorado Plateau (e.g., Levander et al., 2011) may also play a role in driving surface deformation.

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Within the Colorado Rockies, recent geophysical studies (Aster et al., 2009; Schmandt and Humphreys, 2010) reveal a prominent seismic velocity anomaly (the so-called ‘Aspen anomaly’ - Karlstrom et al., 2012) in the upper mantle beneath the region of highest topography and thinnest crust in the Rocky Mountains (Hansen et al., in review; Karlstrom et al., 2012;

Coblentz et al., 2011) (Figure 1). These observations demonstrate that, in fact, there are negative correlations between crustal thickness and high topography, and further attest to the notion of mantle supported topography (Coblentz et al., 2011; Karlstrom et al., 2012; Hansen et al., in review). However, the present day state of the mantle is not indicative of its time history and a definitive understanding of when or how the ‘Aspen anomaly’ evolved remains poorly understood. A number of processes including, lithospheric delamination, small scale convection, or a broader thermal anomaly associated with the Rio Grande Rift (e.g., Moucha et al., 2008) present possible explanations for the seismic velocities and gravity associated with the Rocky

Mountain Province (e.g., Coblentz and Karlstrom, 2011).

Each of these processes makes specific predictions for the magnitude and wavelength of associated deformation. Mantle convection has been proposed to contribute up to ~500 m of dynamic topography between the central and edge portions of the Colorado Plateau (van Wijk et al., 2010). Delamination of the lithosphere beneath the Colorado Plateau could contribute up to

~400-800 m of surface uplift over wavelengths of several hundred kilometers (Levander et al.,

2011). If either of these mechanisms are contributing dynamic support to overall topography in the Rocky Mountain province, they could be compatible with ~500 m of differential uplift between the crest of the range and the adjacent Colorado Plateau and Great Plains.

Quantifying the timescales and pattern of landscape response is critical to identifying potential drivers of differential rock uplift. Along the eastern flank of the Rockies, reconstruction of paleo-transport gradients from Ogallala Formation gravels (McMillan et al., 2002) suggests up to 600 m of tilting in excess of isostatic adjustment over a wavelength of a few hundred

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kilometers (Leonard, 2002), and the restricted temporal window over which this tilting has occurred (post 6 Ma) has been argued to represent relatively localized differential uplift between the Rockies and the Great Plains (Duller et al., 2012). Although the possibility of base-level changes associated with drainage integration complicate the history of incision along the western slope, further study of the history of incision along the western slope may aid in ascertaining whether differential rock uplift along the Rocky Mountain-Colorado Plateau transition can be attributed to mantle processes.

2.2 Timing, Magnitude and Rates of Fluvial Incision along the Western Slope

Much of the geologic evidence for a young component of surface uplift in the Rocky

Mountains comes from a record of overall basin filling switching to basin exhumation and relief production in the Rockies during the late Miocene (e.g., Larson et al., 1975; Izett, 1975; Buffler,

1967, 2003; Kunk et al., 2002; McMillan et al., 2002, 2006). These basins received terrestrial sediments following the Laramide orogeny. Their elevations at the time of deposition, however, are not well known (e.g., Larson et al., 1975; Izett, 1975; Buffler, 1967, 2003; Kunk et al., 2002;

McMillan et al., 2002, 2006). Many of the early workers who documented the late Miocene switch from deposition to exhumation proposed that a pulse of young surface uplift was responsible. However, they had no proposed mechanism for driving changes in surface elevation

(e.g., Izett, 1975; Larson et al., 1975). Recent research has shown broad tilting associated with

Neogene dynamic topography to be possible (e.g., McMillan et al., 2002; Leonard, 2002;

Riihimaki et al., 2007). However, the driver of exhumation remains debated as fluvial incision documents both increases in rock uplift and changes in erosional efficacy (e.g., Molnar and

England, 1990).

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Other workers attribute the most recent pulse of fluvial incision and relief generation to reflect an increase in the transport efficiency of rivers draining the Rocky Mountains. For example, Pelletier (2009) suggests that overall cooling during the late Miocene increased snow- melt discharge for high-elevation channels and induced enhanced fluvial transport. Similarly,

Wobus et al. (2010) develop a model of fluvial incision along the eastern flank of the Rockies and conclude that enhanced fluvial incision due to climate change is a plausible explanation of the patterns and rates of Late Cenozoic incision in this region.

Along the western slope of the Rocky Mountains, a third potential mechanism exists that may drive fluvial incision – base level fall associated with drainage integration across the western

Colorado Plateau (Hunt, Lucchitta, etc.). Base-level change associated with basin integration has long been thought to be a primary driver of incision along the western slope of the Rockies (e.g.,

Pederson et al., 2002b; Kunk et al., 2002), linked to cutting of the Grand Canyon across the western Colorado Plateau (e.g., Pederson et al., 2002b; Kunk et al., 2002; Pederson et al., 2013).

The final integration of the Colorado River through the Grand Canyon and across the Grand

Wash Cliffs is known to have occurred between ca. 5-6 Ma (e.g., Lucchitta, 1990; Dorsey et al.,

2007; Ingersoll et al., 2013). Thus, one prediction of this history is that incision driven by this integration event upstream of the Grand Canyon should only have occurred in the last 6 Ma

(Pederson et al., 2013).

In contrast to this expectation, a growing body of evidence points to a more complicated history of integration for the upper Colorado River. Definitive evidence for a rapid pulse of transient incision related to Grand Canyon integration is not observed upstream of the Glen

Canyon region (e.g., Wolkowinsky and Granger, 2004; Karlstrom et al., 2008; Cook and

Whipple, 2009; Darling et al., 2012). Moreover, a large body of work in western Colorado suggests that relatively rapid incision has been sustained along the Colorado River since ~10 Ma

(e.g., Kunk et al., 2002; Aslan et al., 2010; Karlstrom et al., 2012). In addition, emerging

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evidence from the southern Colorado Plateau suggests a period of significant relief development during late Oligocene – middle Miocene time (Flowers et al., 2007; Cather et al., 2008).

Although, the influence of early relief is not seen upstream, these observations appear to require a more complicated history of incision along the upper Colorado than previously recognized.

2.2.1 Colorado and Gunnison Rivers

Much of the evidence for late Cenozoic tectonism in the Rocky Mountains relies on the history of incision along major drainages (McMillian et al., 2006; Riihimaki et al., 2007). An extensive body of work over the past two decades indicates that the Colorado River has incised

~1100-1500 m across the western slope of the Rockies during the past 10 Ma (e.g., Larsen et al.,

1975; Kunk et al., 2002, Aslan et al., 2010). Below, I summarize the key constraints that underlie this conclusion. In the vicinity of Carbondale, Colorado (Figure 3), regionally-extensive normal faulting is related to collapse during dissolution of underlying evaporite deposits (Scott et al.,

1999; Kunk et al., 2002); localities within this region have subsided by up to 800 m relative to regions outside the collapse center (Kunk et al., 2002). I exclude these sites from discussion here, and refer the reader to the comprehensive work of Kunk et al. (2002) for a synopsis of evaporate collapse along the Colorado River.

Outside the region of evaporite collapse, however, most of the key markers used to reconstruct fluvial incision rely on an association of basalt flows and fluvial gravel deposits

(Figure 3, Table 1). The westernmost of these is located at Grand Mesa, just upstream from

Grand Junction, CO (Figure 3). Here, a sequence of at least 10 individual basalt flows ranging in age from 9.49 +/- 0.16 Ma – 10.76 +/- 0.24 Ma (Kunk et al., 2002) overlie fluvial gravels with

Colorado River provenance (Aslan et al., 2010; Cole, 2010). These river gravels contain clasts that could only have source regions in granitic ranges upstream (Gore, Park, or Sawatch Ranges -

Aslan et al., 2010) and attest to the presence of an established fluvial system draining west from

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the crest of the Rockies at this time. The age of the basal flow at Grand Mesa is 10.76 +/- 0.24

Ma (Kunk et al., 2002), and the underlying river gravels rest ~1500 m above the present-day river. These deposits provide the critical evidence that an ancestral Colorado River was established across the western slope of the Rockies by ~10 Ma, capable of carrying sizeable sediment loads from the crest of the range onto the Colorado Plateau (Aslan et al., 2010).

A second constraint is found ~40 km upstream of Grand Junction, between Battlement

Mesa, south of the Colorado River (Figure 3), and Mt. Callahan, north of the river. Debris-flow deposits on the southern shoulder Mt. Callahan were documented by Berlin and Anderson (2007) and Berlin (2008, 2009) to contain boulders of basalt > 1 m in diameter that overlie ancestral

Colorado River gravels. The boulders are similar in age (~9.17 Ma; Berlin, 2008, 2009) to flows on Battlement Mesa (~9.3 Ma; Berlin 2008, 2009) and are interpreted to represent debris-flow deposits derived from these units and shed northward into the ancient Colorado River valley

(Berlin, 2008, 2009). Deposits on Mt. Callahan overlie fluvial gravels that sit ~1100m above the modern river and thus place a minimum constraint on the timing and rate of incision (Berlin,

2008, 2009). However, the average slope of younger, well-preserved debris-flow surfaces inset along the northern flank of Battlement Mesa (~0.07) is similar to present-day difference in elevation between Battlement Mesa and Mt. Callahan (slope of ~0.03). Moreover, the northern extent of basalt flows on Battlement Mesa has been removed by incision and widening of the canyon along the Colorado River, reconstruction of the paleotopography is somewhat uncertain; flows may have extended ~8-10 km farther north (Stover, 1984). Using the average modern transport slopes of ~0.07 (Berlin, 2009) and the distance from Mt. Callahan to the present day position of the Colorado River (~4-5 km), one can estimate that there may have been up to ~280-

350 m of additional relief. Thus, although ~1100 m is an absolute minimum value of incision

(Berlin, 2009), it seems likely that incision in the vicinity of Mt. Callahan and Battlement Mesa is most likely in the range of ~1380-1450 m.

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The easternmost locality where dated basalts are in direct association with Colorado

River gravels is located at Spruce Ridge (Figure 3). Here, a 7.8 +/- 0.04 Ma basalt flow (Kunk et al., 2002) overlies river gravels with a significant constituent percentage of granitic clasts which are of probable Colorado River origin (Kirkham et al., 2001; Brown et al, 2007). The gravels are

~750 m above the Colorado River (Kunk et al., 2002).

Several other localities between Battlement Mesa and Spruce Ridge constrain Colorado

River incision, but these localities lack a clear association with gravel deposits of Colorado River origin. The exact relationship of these markers to the history of incision is provided in Table 1 and the references therein. At Little Baldy Mountain (Figure 3), the upper of two basalt flows is dated to 10.38 +/- 0.12 Ma and is ~1190 m above the modern Colorado River (Kunk et al., 2002).

Aslan (2012, pers. comm.) describes river gravels made up of dominantly volcanic clasts preserved in the sediments which underlie the basalt flow. Thus, although the deposits at Little

Baldy Mountain may not represent the magnitude of incision along the main-stem of the

Colorado River, they probably help constrain the amount of incision into a locally expansive high elevation erosion surface (e.g., Larson et al., 1975; Kirkham and Scott, 2002). On the south side of Basalt Mountain (Figure 3), a series of five stacked basalt flows lies directly atop the Mancos shale (Kunk et al., 2002). The basal flow has an age of 10.49 +/- 0.07 Ma (Kunk et al, 2002) and is ~1020 m above the modern Colorado River. The presence of numerous stacked flows (Kunk et al., 2002), and the presence of locally derived river gravels overlying the northernmost flow

(Aslan et al., 2010) would seem to suggest that Basalt Mountain was at or near local base level.

Collectively, the observations summarized above suggest that, downstream of Glenwood

Canyon (Figure 3), the Colorado River has incised > 1000 m and probably as much as ~1400-

1500 m during the past 10 Ma. The pace of incision through time, however, remains somewhat poorly understood, in large part because of uncertainty in the timing of when incision began.

The oldest basalts associated directly with gravels of Colorado River provenance occur at Grand

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Mesa where the lowermost basalt flow has an age of 10.76 +/- 0.24 Ma (Kunk et al., 2002). The river gravels and the basalt flow would have both occupied the local base level at the time of their emplacement. Thus, the age of the basalt, and elevation of the river gravels constrain the elevation of local base level for the ancestral Colorado River prior to ~10.7 Ma. They do not, however, directly constrain at what time an incipient canyon had begun to form. In fact, the presence of extensive, ~10 Ma basalt flows, at broadly similar elevations along the Colorado

River suggests the presence of a broad, low relief erosional and/or transport surface prior to ~10

Ma (e.g., Larson et al., 1975; Kunk et al., 2002).

Kunk et al., (2002) suggest that the presence of a young, 3.03 +/- 0.02 Ma, high-elevation basalt at Gobbler’s Knob (Figure 3), ~730 m above the modern Colorado River, suggests an increase in the rate of incision around ~ 3 Ma. Kunk et al., (2002) conclude that the incision rate along the Colorado was slow, ~24 m/Ma, from ~7.8 Ma (Spruce Ridge) to ~3.3 Ma (Gobbler’s

Knob) but then increased to ~240 m/Ma from ~3.3 Ma to present. However, the base of the basalt flow at Gobbler’s Knob is unexposed and is not known to be associated with river gravels

(Kunk et al., 2002). Thus, the flow may have been emplaced significantly above the ancestral

Colorado River ca. 3 Ma and most likely does not directly constrain incision (Aslan et al., 2010).

Comparison of incision rates along this reach of the Colorado River over different timescales, in fact, suggests relatively constant incision rates since ~10 Ma. Correspondence of incision rates have been determined from basalt flows at Basalt Mountain (~100 m/Ma from ~10

Ma) and Spruce Ridge (~100 m/Ma from ~7.8 Ma), and from fluvial terrace gravels at Dotsero

(Figure 3) that contain Lava Creek B tephra (Lidke et al., 2002) (~130 m/Ma from ~640 Ka). All three of these localities are at the same approximate position along the river and collectively, these data suggest that the reach of the river near Glenwood Canyon has experienced relatively constant incision rates of ~100 m/Ma that began sometime between ~10 Ma and ~7.8 Ma (Aslan et al., 2010).

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Likewise, near Rifle, CO (Figure 3), apatite fission-track data from the MWX well along the Colorado River show a 10 Ma partial annealing zone which corresponds to paleotemperatures of ~110 °C, ~3 km paleodepth, at 10 Ma (Kelly and Blackwell, 1990). This partial annealing zone has been exhumed ~ 1500 m in the last 6-10 Ma (Karlstrom et al., 2012). These data also seem to suggest that the onset of rapid Colorado River incision beginning ca. 6-10 Ma (Karlstrom et al., 2012).

In this same region, bedrock strath terraces preserved beneath fluvial gravel and debris- flow/alluvial fan deposits provide additional constraints on the pace of incision through time.

Terrace deposits are preserved beneath tributary fan-terrace complexes that flank the southern valley wall beneath Battlement Mesa (Figure 3) (Stover, 1993; Berlin, 2009; Aslan et al., 2010).

Dating of two of these fan-terrace surfaces via 26Al/10Be burial dating of quartz provide constraints on incision rate over the past 1-2 Ma. The Grass Mesa terrace (Figure 3) sits ~225 m above the Colorado River and is 1.77 + 0.71/-0.51 Ma (Berlin, 2008; Berlin et al., 2009); this date suggests a minimum incision rate of 115 +27/-23 m/Ma (Berlin, 2008, 2009). A second terrace at

Morrisania Mesa sits ~94 m above the modern river and is dated to 0.44 Ma +/- 0.3Ma (Darling et al., 2012). These data yield a poorly constrained incision rate of ~214 m/Ma (127-671 m/Ma)

(Darling, 2010). Overall, these observations suggest that in the vicinity of Grand Mesa, the

Colorado River began relatively rapid and sustained incision of ~100-150 m/Ma beginning around10 Ma; in general agreement with data from upstream near Dotsero.

Although the Colorado River appears to have a history of relatively steady incision, the

Gunnison River displays a major and pronounced knickpoint in the Black Canyon region (e.g.,

Sandoval, 2007; Aslan et al., 2008; Darling et al., 2009; Donahue et al., in review). Abundant strath terraces that contain the ~0.64 Ma Lava Creek B tephra (Lanphere et al., 2002) reveal striking spatial differences in incision rate across this knickzone. Downstream, near Grand

Junction, incision rates are ~150 m/Ma (Sandoval, 2007; Aslan et al., 2008; Darling et al., 2009).

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These rates increase within the Black Canyon to ~400-550 m/Ma (Sandoval, 2007; Aslan et al.,

2008), but decrease again upstream to ~90 m/Ma above the knickzone (Sandoval, 2007; Aslan et al., 2008). Thus, the Black Canyon knickzone is clearly a prominent expression of transient incision along this system that appears to be a consequence of the abandonment of Unaweep

Canyon by capture of the Gunnison River (e.g., Hansen, 1987; Donahue et al, in review; Aslan et al., in review); an event that appears to have taken place in the last 1-2 Ma (Donahue et al, in review; Aslan et al., in review).

2.2.2 White, Yampa, and Little Snake Rivers

In contrast to the reasonably well-understood history of incision along the Colorado

River, relatively little is known regarding the timing and magnitude of incision along the western slope of the Rockies in northern Colorado. Here, the White, Yampa and Little Snake rivers are not entrenched in narrow canyons, and deposits of ancient fluvial gravels are exceedingly rare.

However, the region was the locus of sediment accumulation during Oligocene through Miocene time (Kucera, 1962; Buffler, 1967; 2003; Izett, 1975; Larson, 1975; McMillan et al., 2006), and these deposits, collectively referred to as the Browns Park Formation (originally described by

Powell, 1876 and summarized by Kucera, 1962), have been incised, dissected and eroded. Thus, the degree of preservation of basin sediments allows for a minimum estimate of both the timing and magnitude of mass removed by fluvial activity (e.g., McMillan et al., 2006).

Regionally, the Browns Park Formation extends from the Elkhead Mountains in the northeast, the Flattops in the south, and the Browns Park graben sensu stricto in the west (Figure

4). There are two primary members of the Browns Park Formation; a lower basal conglomerate and an upper sandstone member as described by Buffler, (2003). The basal conglomerate rests everywhere unconformably on older strata including Precambrian rocks exposed in the Park

Range, varies in thickness up to ~90 m, and is composed primarily of calcareous-cemented,

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poorly sorted gravel and coarse sand dominated by varied Precambrian lithic clasts, i.e. ‘red granite’, granite gneiss, biotite and amphibole gneiss, and schist (Buffler, 2003). The conglomerate generally thickens and becomes coarser and more angular as it approaches the margins of the basin, and is interpreted to represent alluvial fans being shed from the Park and

Sierra Madre Ranges to the east and southeast of the Sand Wash basin (Buffer, 2003).

The upper sandstone member of the Browns Park Formation is up to ~670 m thick and consists of two prominent, recognizable facies; a very fine to medium grained, ‘white sandstone’ facies and a very fine grained sandstone to siltstone, ‘brown sandstone’ facies (Buffler, 2003).

The white sandstone facies lacks finer or coarse grained material, is sub-angular to sub-rounded, contains abundant frosted and pitted grains, and is characterized by large scale (9 -15 m high) trough cross bedding (Buffler, 2003). This facies is interpreted to represent primarily eolian deposition (Buffler, 2003). The brown sandstone facies is moderately well sorted, angular to sub- rounded, and massive, but crudely stratified and while not as definitive as the ‘white sandstone’ facies is also interpreted to represent primarily eolian deposition (Buffler, 2003). Although dominated by eolian deposits, sub-facies within the upper member show locally coarser sandstone deposits with small-scale cross-stratification and ripple marks associated with massive to laminated siltstone suggestive of local fluvial channel and floodplain deposits (Buffler, 2003).

Sedimentary structures within the Browns Park Formation suggest major changes in the dominant mode and direction of sediment transport through time. Paleo current directions of eolian dunes and localized fluvial channels within the Browns Park Formation show an overall

NNE sediment transport direction (Buffler, 1967; 2003). The modern sediment transport direction within areas of previous Browns Park deposition is predominately westward via the major drainages of the White, Little Snake, and Yampa Rivers. Additionally, the upper member is often inter-bedded with tuffaceous deposits (Izett, 1975; Naeser et al., 1980; Snyder, 1980;

Luft; 1985; Honey and Izett; 1988; field observations from this study). Where thick sections of

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the upper Browns Park are exposed they often contain multiple ash beds spanning a large age range (Snyder, 1980) suggesting that deposition was relatively continuous. The switch from a time of approximately continuous, ~NNE, eolian, aggradation to westward fluvial transport and basin dissection may reflect a change in the local base-level or uplift that is approximately coeval with the cessation of Browns Park deposition along the western flank of the Rockies.

The age range of the Browns Park Formation is approximately 24.8 Ma to 7.2 Ma (Izett,

1975; Luft, 1985; summarized by Buffler, 2003). The age of the basal conglomerate member is constrained by several deposits of volcanic ash which underlie the formation (e.g., Izett, 1975;

Luft, 1985). The oldest ash deposit lies directly on Precambrian rocks on the flank of the Sierra -

Madre range at Dead Mexican Park (Locality 1, Figure 4) and has an age of 24.8 +/- 2.4 Ma

(Snyder, 1980). This age agrees generally with that of an ash layer located along the west bank of the Little Snake River just north of the Yampa confluence (Locality 2, Figure 4), ~ 30 m above the base of the Browns Park, dated to 24.8 +/- 0.8 Ma (Izett et al., 1970). The upper member of the Browns Park is dated by basalt flows which overlie the formation and numerous shallow intrusions and ash layers within the deposits themselves (e.g., Buffler, 1967, 2003; Izett, 1975;

Honey and Izett, 1988; Naeser et al., 1980; Segerstrom and Young, 1972; Larson et al., 1975;

Luft, 1985). Basalt flows that cap thick sections of the Browns Park have K-Ar ages which average around ~ 10 Ma (Buffler, 1967). At City Mountain (Locality 3, Figure 4), a latite porphyry intruding the formation is 7.6 +/- 0.4 Ma (Buffler, 1967). Additionally, an ash reported near the top of the Browns Parks located along Vermillion Creek (Locality 4, Figure 4) has an age of 7.2 +/- 0.6 Ma (Naeser et al., 1980). However, this age may be a minimum as it is a fission track age of hydrated glass shards (Naeser et al., 1980). A zircon fission track age from the same sample is 9.8 +/- 0.4 Ma (Naeser et al., 1980). Collectively, these data suggest that sediment accumulation in the Sand Wash basin continued from ~ 24 – 9 Ma, and possibly as young as 7

Ma, across the region.

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Of particular relevance to this study are basalt flows that cap mesas throughout the region and often overlie thick sections of Browns Park Formation. Near the headwaters of the Little

Snake River in the Elkhead Mountains, basalt flows have previously published K-Ar ages ranging from 9.5 +/- 0.5 Ma to 10.7 +/- 0.5 Ma and overlie ~400-600 m of Browns Park Formation

(Buffler, 1967, 2003). In the headwaters of the White and Yampa River, numerous basalt flows in the Flat Tops range inter-finger with a slightly thinner (~300 m) section of Browns Park

(Larson et al., 1975; Kunk et al., 2002). In both of these regions, the age of the uppermost

Browns Park Formation is similar to the flows themselves. As these flows overlie the Browns

Park Formation, they broadly confirm a minimum age for the formation of ~ 10 Ma (Buffler,

2003). Field relationships suggest, however, that local relief generation during fluvial incision most likely post-dates basalt emplacement, and thus incision appears to have begun sometime after 10 Ma in this region.

Although fluvial terraces inset beneath the top of the basin fill are sparse, the Lava Creek

B tephra has been discovered within terrace tread alluvium at several localities (Izett, 1975; Izett and Wilcox, 1982; Dethier, 2001). Rates of incision at these localities range from ~75 m/Ma (Elk

River) to ~160 m/Ma (White and Yampa rivers) (Dethier, 2001; Aslan et al., 2008, 2010).

Although rates are in general agreement with those derived from the ~10 Ma surface defined by basalt flows and the basin fill, they are a slightly greater. Whether this apparent increase represents a climatically enhanced increase in fluvial incision rates during the mid-Pleistocene

(e.g., Pelletier, 2009) or the influence of base level fall associated with incision downstream (e.g.,

Pederson and Tressler, 2012) remains to be determined.

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2.3 Analysis of River Longitudinal Profiles

Analysis and interpretation of longitudinal profiles of bedrock channels, those that are actively incising into mountainous landscapes (e.g., Whipple, 2004), has become a relatively common tool to guide the interpretation of landscape evolution in erosional settings. Although applications of this type of analysis are typically applied in convergent mountain ranges where differential rock uplift is associated with permanent deformation of the crust (e.g., Seeber and

Gornitz, 1983; Kirby and Whipple, 2001; Kirby et al., 2003; Wobus et al., 2006; Harkins et al.,

2007; Ouimet et al., 2009; Merritts and Vincent, 1989; Snyder et al., 2000; Duvall et al., 2004;

Safran et al., 2005), a number of recent studies try to export these techniques to regions of long- wavelength, epiorogenic uplift (e.g., Miller et al., 2010; Crow and Karlstrom, 2011; Miller et al.,

2011; Karlstrom et al., 2012). Here, I use these techniques to examine possible drivers of

Miocene exhumation related to possible epiorogenic uplift along the western flank of the Rocky

Mountains. Outlined below are some of the theoretical principles behind channel profile analysis; although, the reader is directed to several reviews of the subject for a more comprehensive examination of this technique (e.g., Whipple, 2004; Whipple, 2011; Kirby and Whipple, 2012).

Channel profile analysis exploits the empirical scaling relation between the local channel gradient (S) and the contributing drainage area upstream (A). In graded channel profiles (Mackin,

1948) from mountain ranges around the world, the empirical relationship follows the form,

where ks is a measure of the relative channel steepness, or so called ‘channel steepness index’, and θ is the ‘concavity index’, or a measure of how rapidly slope varies with upstream drainage area (e.g., Flint, 1974; Snyder et al., 2000; Kirby et al., 2003; Wobus et al., 2006). In practice, the steepness index (ks) and concavity index (θ) can be determined by linear regression of slope

(S) against drainage area (A) in log-log space. However, small uncertainties in the slope of this

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regression (θ) yield large variations in the regression intercept (ks) (Wobus et al., 2006). Thus, several methods for measuring a normalized gradient index, which address this effect, have been proposed (e.g., Sklar and Dietrich, 1998; Royden et al., (2000); Sorby and England, 2004; Wobus et al., 2006). However, determining normalized channel steepness (ksn) (c.f. Wobus et al., 2006) utilizing a reference concavity index (θref), addresses this effect and allows for the comparison of channels with different contributing drainage areas (Wobus et al., 2006)

There are a number of empirical studies that examine how the steepness of a channel, gradient normalized for area (ksn) (e.g., Wobus et al., 2006) varies with erosion rate. Early in the development of the field, studies were limited to steady-state landscapes where uplift rates were known from independent geomorphic markers (e.g., Snyder et al., 2000; Kirby and Whipple,

2001; Duvall et al., 2004); these results support theoretical predictions (e.g., Whipple and Tucker,

1999) that under conditions of uniform substrate (Moglen and Bras, 1995), spatially invariant uplift (Kirby and Whipple, 2001), and uniform precipitation (Roe et al., 2003) i) the normalized channel steepness (ksn) scales monotonically with uplift rate and ii) the concavity index () is relatively insensitive to uplift rate. These results bolstered the use of channel profile analysis as a tool to determine spatial patterns of rock uplift (Wobus et al., 2006). In recent years, the application of cosmogenic isotopic inventories in modern sediment to measure basin averaged erosion rates (e.g., Bierman and Steig, 1996; Granger et al., 1996) has enabled comparisons of channel steepness (ksn) and local, catchment scale, erosion rates (e.g., Safran et al., 2005; Harkins et al., 2007; Ouimet et al., 2009; DiBiase et al., 2010; Cyr et al,. 2010). These studies all suggest that erosion rates correlate well with channel steepness and allow for the application of stream profile analysis to a wide variety of landscapes (Kirby and Whipple, 2012).

The expected correlation between channel steepness and erosion rate allows this analysis to reflect the combined effects of substrate lithology, climatic influences, and rock uplift, making it potentially useful for assessing the rates and patterns of differential rock uplift if the effects of 23

lithology and climate can be controlled for (Kirby and Whipple, 2012). However, as transient river profiles have been recognized in tectonically active landscapes around the world (e.g.,

Wobus et al., 2006; Harkins et al., 2007; Kirby et al., 2007; Berlin and Anderson, 2007; Cook et al., 2009; Morell et al., 2012), interpretation of these landscapes requires an understanding of the response of river profiles to forcings that disrupt the equilibrium profile (Kirby and Whipple,

2012). Perturbations such as uplift or base level lowering can generate transient knickpoints (e.g.,

Howard, 1994; Whipple and Tucker, 1999, 2002; Whittaker et al., 2007). These can be thought of as a migrating boundary between downstream reaches that are adjusting to the new forcing and upstream reaches that are adjusted to the previous background state (Howard et al., 1994; Crosby and Whipple, 2006). As different channel segments will typically have different values of channel steepness index and/or concavity index (Wobus et al., 2006), channel steepness analysis can still be used as a method of understanding and interpreting transients in the fluvial system.

2.4 Controls on Channel Profile Form along the Western Slope of the Rocky Mountains

Preliminary analysis of modern channel profiles draining the western slope of the

Colorado Rockies provides motivation for the present study. Karlstrom et al., (2012) show that channels that drain high topography above the ‘Aspen anomaly’ have higher normalized steepness indices (c.f. Wobus et al, 2006) than those that lie along the northern flank of the anomaly (Figure 2). This observation suggests that the southern channels, which overlie the anomaly, may be responding to recent surface deformation associated with buoyant mantle.

However, the influences of potential variations in discharge and/or substrate lithology were not evaluated in this analysis. A recent re-evaluation by Pederson and Tressler (2012) attempted to account for variations in discharge, and importantly, this study found that the difference in

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channel steepness between the Colorado and northern rivers is not an artifact of variations in discharge.

The role of variations in substrate lithology, however, remains an open question.

Pederson and Tressler (2012) argue that substrate lithology is the dominant control on the position of knickpoints along the Green-Colorado River system. They argue that knickpoints and knickzones are anchored to resistant substrate and act to ‘decouple’ topography from proposed loci of uplift (e.g., van Wijk et al., 2010). In contrast, I examine the present steepness of channels and the timing and magnitude of incision for rivers draining the western slope with the hope that this method may identify areas of possible uplift along the western flank of the Rockies.

In this study, I begin with the premise that correlations exist between erosion rate and channel steepness (ksn) and that how these vary in space may provide insight into the potential controls on fluvial incision. Because these rivers may not be in steady state (e.g., Berlin and

Anderson, 2007), it is important to identify transients in the system that may be associated with knickpoints which separate regions of different channel gradient (Kirby and Whipple, 2012) and distinguish these from steep channel segments that are anchored to locally resistant substrate

(e.g., Pederson and Tressler, 2012). Finally, I test the hypothesis that there is young deformation associated with the ‘Aspen anomaly’, by examining whether spatial correlations exist between channel steepness, the record of incision, and regions of low-velocity mantle which underlie rivers along the western slope.

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3 Analysis of Channel Profiles along the Western Slope

3.1 Measuring Normalized Channel Steepness Index

I extract channel long-profiles and determine normalized channel steepness values (ksn) for six of the major rivers draining the western flank of the Rockies: the Colorado, Gunnison, and

Dolores Rivers, and the White, Yampa, and Little Snake Rivers upstream of their respective confluences with the Green River. Calculation of channel steepness values and extraction of channel profiles follow the methods, and utilize the ArcGIS add-ons and MATLAB scripts, of

Wobus et al., (2006); available at www.geomorphtools.org. Topographic data and upstream drainage area were extracted from USGS 30m DEMs (The National Map – http://nationalmap.gov; Data available from U.S. Geological Survey, Earth Resources

Observation and Science (EROS) Center, Sioux Falls, SD). Channel profiles were smoothed using a moving-average window of 1 km and vertical sampling interval of 12.192 m.

The initial long profiles extracted for the Colorado, Gunnison, and Dolores rivers contain sharp, localized, breaks in slope that correspond spatially to major reservoirs. The locations of these reservoirs were verified against a USGS shapefile (http://nationalmap.gov) and against

USGS topographic maps (http://nationalmap.gov). Only the Colorado, Gunnison, and Dolores rivers displayed reservoirs not significantly filtered by the 1 km smoothing window. Reservoirs that remained after smoothing were manually removed by linear interpolation of the channel elevation just upstream and downstream of each reservoir (Figure 5). I acknowledge that this alters the local gradient and local ksn values. It should not, however, affect regional interpretations of channel steepness patterns.

Once reservoirs were removed, log(S)-log(A) plots were extracted for all six study area rivers and regressed to determine values of ksn along each channel (c.f. Wobus et al., 2006). A reference concavity (θref) of -0.45 was used for this, and all further, ksn analyses in this study. 26

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Additionally, log(S)-log(A) plots were used to identify and classify major knickpoints/knickzones along each river. Prominent knickpoints were classified into either vertical-step or slope-break knickpoints based on their characteristics in log(S)-log(A) space (e.g., Haviv et al., 2010; Whipple et al., 2011; Kirby and Whipple, 2012). Vertical-step knickpoints are characterized by a local, discrete increase in channel steepness, whereas, slope-break knickpoints are characterized as separating two distinct reaches of a channel, each with a different steepness (Whipple et al., 2011;

Kirby and Whipple, 2012). However, neither the log(S)-log(A) regressions nor the knickpoint classifications are used as the primary data set for this study as these channels flow over similar, but highly spatially variable, lithologies along the western slope upstream of their respective confluences with the Green River. Therefore, these preliminary estimates of channel steepness may be at least partially controlled by the underlying bedrock lithology (e.g., Whipple, 2004;

Kirby and Whipple, 2012).

3.2 Evaluating the Effect of Substrate Lithology on Channel Steepness

In order to evaluate the effect of substrate lithology on channel steepness, ksn values were calculated at an interval of 0.5 km along each river utilizing an automated regression (c.f.,

Whipple et al., 2007). Following the automated regression, the 0.5 km spaced measurements of ksn were placed into 10 km bins along each channel. The average ksn value for each bin then provides a measure of ‘local normalized channel steepness’, or ‘local ksn’, at a spacing of 10 km and the standard deviation provides an estimate of the error for each bin. Figure 6 shows an example of local ksn values plotted along the long profile of the Colorado River. Displaying ksn values in the form shown in Figure 6 allows for visual examination of ksn values along the long profile in distance-elevation (normal) space.

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The bedrock geology for the study area was then extracted from the digital geologic maps of Colorado (Green, 1992, after Tweto, 1979), Utah (Hintze et al., 2000; Hintze, 1980), and

Wyoming (Green and Drouillard, 1994, after Love and Christiansen, 1985) and divided into map units that are expected to have broadly similar resistances to erosion (Figure 7). In general, I expect that older, more consolidated and compacted, sediments will show more resistance to erosion and will, therefore, have potentially higher values of local ksn. These simplified lithologies are then plotted along channel profiles and compared to the 10 km spaced values of local ksn (Figure 9). This allows for the comparison of local normalized channel steepness and lithology.

However, in order to examine local steepness variations, within differing lithologies, across the six study area rivers, ksn measurements were binned by formation and, where appropriate, classified into either ‘shales’ or ‘Tertiary sandstones’. The Browns Park Formation and equivalent, Uinta Formation and equivalent, and the Wasatch formation and equivalent are grouped into ‘Tertiary sandstones’ while the Mancos Shale, Lewis Shale, and Green River

Formation and equivalent are grouped into ‘shales’. These two lithologic groups were selected because they represent broadly different lithologies, shale vs. sandstone, but encompass formations that are readily identifiable from one another and are present along many of the study rivers. For each river, all of the 0.5 km ksn values within reaches of self-similar lithology (e.g., the Wasatch Formation and equivalent) were binned; with bin value and error calculated as before. Only lithologic-reaches longer than 10 km were included in this analysis. The Dolores

River was excluded in its entirety because it flows only over Paleozoic/Mesozoic rocks (Figure 7 and Figure 9) and is, therefore, not readily compared to the other study rivers. Additionally, some reaches were excluded because they displayed a signal interpreted to be of local origin; a discussion of this is presented later.

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3.3 Evaluating Relationships between Discharge and Drainage Area

Typically, gradient indices such as those used here (ksn) utilize upstream drainage area as a proxy for discharge (c.f. Wobus et al., 2006). This means that comparison of ksn values among different channels implicitly assumes that the relationship between discharge and upstream drainage area does not vary between drainage basins. However, differences in the amount of evapotranspiration among catchments may yield varying values of discharge at equivalent drainage areas. If systematic variations in the relationship between discharge and upstream drainage area exist between channels, then differences in normalized channel steepness may be misleading (e.g., Pederson and Tressler, 2012).

In order to examine whether it is reasonable to compare ksn values between the study rivers, I compare the discharge-drainage area relationships of each river. Gridded discharge values were derived from the PRISM precipitation model (http://www.prism.oregonstate.edu) by calculating the resultant flow-accumulation for each catchment from 30 m DEMs using ArcGIS.

The gridded precipitation data are a 30-year average of 30 annually-averaged precipitation models from 1971-2000 (http://www.prism.oregonstate.edu) and represent values of 30-year average discharge. Finally, discharge values were extracted from the spatial grid along the channel paths of the study rivers and plotted against contributing drainage area (Figure 8).

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4 Results of Channel Profile Analysis

4.1 Relationships between Discharge and Drainage Area

Previous studies have demonstrated that differences in discharge along the western slope is not a sufficient explanation for the observed differences in channel steepness (Pederson and

Tressler, 2012). Discharge-drainage area relationships determined for this study (Figure 8) reveal similar scaling relationships among all six study area rivers. Although the Gunnison and Dolores

Rivers carry somewhat lower discharge at equivalent drainage areas (Figure 8), the Colorado,

White, Little Snake, and Yampa Rivers have nearly identical discharge-drainage area relationships. Values of gridded discharge generated from the PRISM precipitation model compare favorably with the record of discharge from USGS gauging stations

(www.waterdata.usgs.gov) averaged in the same manner (30-year average of annual discharge records from 1971-2000) and provide a reasonable way to estimate mean annual discharge along reaches of rivers in the study area with limited numbers of gauging stations. Thus, at least from the Colorado north to the Little Snake, it appears that differences in channel steepness do not reflect differences in the mean annual discharge and how it varies downstream. These results broadly confirm the conclusions of Pederson and Tressler (2012).

4.2 Colorado, Gunnison, and Dolores Rivers

The profile of the Colorado River shows broad variations in channel steepness between the headwater reaches and the confluence with the Green River (Figure 9A). Generally, the lowest values of local ksn (~20 - 40) are observed immediately upstream of the confluence of the

Green River, whereas the highest values of local ksn (~90 - 100) occur in the middle of the profile.

The uppermost ~200 km of the profile are again gentler (local ksn ~60 - 70). Superimposed on

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this general trend, local high values of ksn correlate with the position of distinct knickpoints or knickzones along the Colorado River. These correspond directly to major canyons along the river including Westwater Canyon, Glenwood Canyon, and Gore Canyon (Figure 9A and Table 2).

The general association of these knickpoints with reaches developed in crystalline basement rocks and the relatively similar steepness values for reaches both upstream and downstream of the knickzones suggest that these steep reaches are likely anchored to the underlying bedrock lithology (Kirby and Whipple, 2012). A similar interpretation has been put forward by Pederson and Tressler (2012) for many of these features along the entire extent of the river profile.

Notably, local ksn values gradually increase upstream from the Green River confluence to approximately the position of Glenwood Canyon (Figure 9A). Regression of slope-area data

0.9 through this reach yields a fit of ksn = 82.8 m , θ = 1.1 +/- 0.24. The high concavity index of the lower reach in this case reflects the streamwise change in channel steepness. Although there is a

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slight decrease in local ksn upstream of Glenwood Canyon (Figure 9A), there does not appear to be a sharply defined break in slope-area scaling between the reach directly above and below

Glenwood Canyon (Figure 9A). The general pattern of progressively increasing normalized steepness that does not differ systematically across locally steep reaches could reflect either 1) a river profile that is still adjusting to a transient perturbation and has not yet reached an equilibrium shape, or 2) uplift rates that vary along the profile (e.g., Kirby and Whipple, 2001).

These possibilities are discussed in greater detail in section 7.

Regression of slope-area data along the Colorado River above and below Gore Canyon, however, is more likely consistent with a slope-break knickpoint. The reach below the knickpoint

0.9 exhibits relatively high normalized steepness and high concavity index (ksn = 89.6 m , θ = 0.8

+/- 0.19) whereas the reach above the knickpoint exhibits both lower steepness and concavity (ksn

= 55.2 m0.9, θ = 0.49 +/- 0.034) from ~570-700 km (Table 2).

In contrast to the Colorado, the channel profile of the Gunnison River is characterized by a prominent knickzone within the Black Canyon of the Gunnison, ~400-500 km upstream from the Colorado – Green confluence (Figure 9B). Regression of slope-area data above the knickzone

0.9 yields a fit of ksn = 43.6 m , θ = 0.55 +/- 0.19, whereas the reach below the knickzone is similar

0.9 (ksn = 55.4 m , θ = 0.61 +/- 0.41). Locally, ksn values within the knickzone are much greater,

(Figure 9B). Similar to knickzones along the Colorado River, the steep reach within Black

Canyon of the Gunnison is developed within Precambrian crystalline rocks. Although the high normalized steepness values observed within the knickzone could be singularly controlled by lithology, constraints on short term incision rates across the knickzone clearly illustrate that this feature is related to an upstream migrating wave of incision (Sandoval, 2007; Darling, 2009;

Donahue et al, in review; Aslan et al., in review). Due to the influence of this transient knickzone, the Gunnison is excluded from further discussion.

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Directly upstream from its confluence with the Colorado River, the Dolores River displays variable, but relatively high values of local ksn (Figure 9C). These high values of local ksn near the confluence may suggest adjustment of the Dolores River to base level lowering along the Colorado River or the influence of variable substrate (the Dolores flows through the Permian

Cutler sandstone and the Morrison Formation through this section). Regression in log(S)-log(A)

0.9 space, however, yields a fit of ksn = 44.7 m , θ = 0.2 +/- 0.13. A prominent knickpoint occurs

~450 km above the confluence with the Green River. Regression above the knickpoint yields a

0.9 fit of ksn = 29.5 m , θ = 0.69 +/- 0.16. However, as poor control on the magnitude of incision or the effect of lithology exists along the Dolores it is excluded from further discussion.

4.3 Yampa, White and Little Snake Rivers

The White, Yampa and Little Snake rivers are all tributaries of the Green River that drain the western slope of the northern Colorado Rockies (Figure 1). Because the Little Snake

River is itself a tributary of the Yampa, we discuss this profile first. The lower reach of the

Yampa River, between the confluence of the Little Snake and the Green Rivers, coincides with

Dinosaur Canyon (Figure 9E) which is cut through the eastern tip of the Uinta block (Hansen,

1986). Within this reach, values of local ksn are generally high and exhibit rather substantial scatter. Within the Uinta block, rocks of the Uinta Mountain Group are typically quite resistant and may contribute to the steepening of the river profile, either directly (Pederson and Tressler,

2012) or through the input of coarse debris (Grams and Schmidt, 1999). In addition, this region is the locus of late Cenozoic faulting (Hansen, 1986), and the profile may be steepened by young or ongoing deformation. Alternatively, the Dinosaur Canyon knickzone could represent a transient feature associated with integration of the Yampa River to the Green River (Darling et al., 2012).

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Upstream of Dinosaur Canyon, the Yampa River displays relatively uniform values of local ksn along its profile. A slope-area regression of the entire channel upstream of Dinosaur

0.9 Canyon, however, yields a reasonable fit of (ksn = 67.9 m , θ = 0.6 +/- 0.044). There is, however, a noticeable increase in the headwater reaches of the river (Figure 9E). There are two possible explanations for this. First, the river heads in the basalt fields that comprise the Flat

Tops, and it seems reasonable to infer that profile steepening may be associated with coarse debris shed from this range. As shown in Figure 13, the majority of Quaternary deposits mapped within the Flat Tops province are landslides and coarse debris (Larson et al., 1975). However, the headwater region of the Yampa also overlie the western flank of the Aspen Anomaly (Figure 1), and so it is also possible that these steepened reaches reflect long-wavelength tilting associated with this feature.

The profile of the Little Snake River exhibits relatively uniform values of normalized channel steepness (Figure 9D). One exception to this is a locally steepened reach that corresponds with a prominent knickpoint at ~310 km above the confluence with the Green River.

0.9 Regressions of slope-area data above and below the knickpoint yield fits of ksn = 36.3 m , θ =

0.9 0.55 +/- 0.11 and ksn = 40.4 m , θ = 0.58 +/- 0.11, respectively (Table 2). Because the knickpoint along the Little Snake separates two channel reaches of approximately equal steepness it likely represents a vertical-step knickpoint (e.g., Whipple et al., 2011; Kirby and Whipple,

2012). As with knickzones along the Colorado River, the association with underlying crystalline rocks suggests that it is probably a static feature associated with the local variations in substrate erodibility.

The White River exhibits a smooth profile with relatively uniform values of local ksn along its entire length and no obvious knickpoints. An increase in local ksn values is observed in the uppermost headwaters of the White River (Figure 9F). Again, this increase may be due to coarse debris being shed off of the Flat Tops, or may be a signal of deformation associated with

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0.9 the Aspen Anomaly. A slope-area regression of the entire channel yields a fit of (ksn = 82.7 m ,

θ = 0.5 +/- 0.022).

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5 New Constraints on Late Miocene Exhumation

Although a broad body of work demonstrates that the Colorado River has incised up to

1500 m in the last 10 Ma along the western slope of the Rockies (e.g., Larsen et al., 1975; Kunk et al., 2002, Aslan et al., 2010), our knowledge of the incision histories of the White, Yampa, and

Little Snake rivers to the north is relatively sparse. Preliminary constraints from the preserved thickness of the Tertiary Browns Park Formation suggest substantially less incision along the

Little Snake River (Buffler, 2003). However, the timing and magnitude of incision along these northern rivers remains poorly constrained overall. New 40Ar/39Ar ages from this study on basalt flows capping the Browns Park Formation provide robust constraints on the incision histories of the White, Yampa, and Little Snake Rivers.

Following the cessation of deposition in the Sand Wash Basin, incision along the White,

Yampa, and Little Snake river systems has removed a significant volume of both Browns Park strata and the underlying bedrock. Where the Browns Park Formation is capped by basalt flows or shallow intrusive volcanic features, the topography has been inverted and the flows stand high above the surrounding landscape (Figure 10). I use these relationships to placed bounds on the magnitude of late Miocene exhumation by measuring the amount of relief generated locally adjacent to markers that represent the relative position of the ancestral land surface. Generally, these markers are basalt flows that cap mesas, but in a few cases, I use the exposed thickness of the Browns Park formation where the upper most strata are well dated by interbedded tuffs, or other dateable deposits that place constraints on the position of the ancestral land surface.

Additionally, shallow intrusions place constraints on the magnitude or timing of exhumation in places.

As basalt flows in the Sand Wash basin are not known to be associated with deposits of fluvial gravels, as along the Colorado River, there is some uncertainty in using these to estimate

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relief generated by fluvial incision. Without the association of river gravels to constrain the position of the ancestral base-level, there is an imprecise knowledge of the relationship between the position of an individual flow and the fluvial network at the onset of significant fluvial activity. In this study area, the presence of late Miocene faulting, previously described to be on the order of a few hundred meters (e.g., Kucera, 1962; Buffler, 1967) could have generated significant relief prior to the onset of exhumation. Thus, I suggest that the local preserved thickness of Miocene basin fill sediments (the Browns Park Formation) records the minimum possible relief generated by fluvial incision and the height of flows above the modern river records the potential maximum. In both instances, I assume that minimal deposition occurred following emplacement of lava flows and that the elevation of the base of each flow is the best marker available for constraining the position of the ancestral land surface locally.

Minimum estimates of local relief generated during incision were determined by taking the greatest preserved thickness of Browns Park Formation as mapped on USGS 1 x 2 degree quads (Tweto, 1976) and the geologic map of Wyoming (Love and Christiansen, 1985). The positions of Browns Park contacts mapped by the USGS are taken to be accurate to +/- 10 m. In a few localities, basalt flows are present on both sides of the modern river channel, and these provide robust constraints on the amount of fluvial incision. Correspondence between estimates of incision in these locations and the exposed thickness of Browns Park Formation suggests that the amount of relief generated locally is a reasonable estimate of fluvial incision following the emplacement of basalt flows.

In order to more precisely estimate when exhumation began, I supplement existing chronology with new 39Ar/40Ar ages from basalt flows. Table 3 shows a summary of new age data from this study (see Appendix for detailed 39Ar/40Ar methods and release spectra information). To address local relationships between the timing of deposition and emplacement between volcanic units, I carefully evaluated the geologic relationships at each locality where I

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reconstruct the magnitude of exhumation. For the purposes of this discussion, I organize localities into three areas of approximately self-similar geologic relationships: the Elkhead

Mountains, the Flattops, and the Yampa River Valley.

5.1 Reconstructing Late Miocene Exhumation in the Elkhead Mountains

The Elkhead Mountains comprise a region of high topography near the

Colorado/Wyoming border (Figure 4 and Figure 11) drained primarily by the Little Snake River.

Here, late Cenozoic volcanics intrude and overlie the Browns Park Formation (Buffler, 2003) and hold up the areas of highest relief. Of importance to this study, significant late Cenozoic extensional faulting is documented in the region and may be on the order of ~300 – 600 m

(Buffler, 1967).

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5.1.1 Battle Mountain, Squaw Mountain, and Bible Back Mountain

At Battle Mountain (Figure 11), basalt flows overlying the upper member of the Browns

Park Formation are exposed atop the mesa. In a recent landside, the basal contact of an ~15 m thick flow with the underlying Browns Park and two thin, ~0.5 m, tuffaceous layers present in the

Tertiary strata beneath are exposed. The elevation of the flow base is ~2680 m and is ~650 m above the elevation of the Little Snake River. The local thickness of Browns Park Formation at

Battle Mountain is ~620 m, and thus I estimate local relief generated by fluvial incision to be

620-650 m. We determined a new 40Ar/39Ar age of 11.46 +/- 0.04 Ma of the basalt flow. Taking the flow base as a datum, this implies average incision rates of ~54-57 m/Ma.

Squaw Mountain sits directly across the Little Snake River southeast of Battle Mountain

(Figure 11). Here, basalts also cap the mesa, but their base is not exposed, complicating the interpretation of whether these outcrops represent extrusive flows. Outcrops are non-vesiculated, free of significant phenocrysts and evidence for an intrusive or extrusive origin is equivocal.

However, exposed just below the base of the outcrop are deposits of a volcanic breccia that is typically associated with extrusive flows elsewhere in the region (Buffler, 1967). These volcanic breccia deposits strongly suggest a local surface vent and that the volcanic feature atop Squaw

Mountain is a flow, similar to the interpretation of Buffler (1967).

The exposed thickness of the probable flow atop Squaw Mountain is ~20 m. The new

40Ar/39Ar age on the most basal exposure found is 11.45 +-/- 0.04 Ma. The most basal exposure is at an elevation of ~ 2550 m and is ~520 m above the modern elevation of the Little Snake

River. The Squaw Mountain flow caps ~510 m of Browns Park Formation; making the value of local relief generation 510-520 m. This yields an estimate for incision rate of 45 +/- 1 m/Ma.

Just southwest of Battle Mountain and across the Little Snake River, at Bible Back

Mountain (Figure 11), the base of ~10 m thick, columnar jointed, flow is exposed on the south

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side. Here, it appears that there may be a second flow of similar thickness above this outcrop, but the nature of the exposure made this upper outcrop inaccessible. The new 40Ar/39Ar age of the basal flow outcrop is 11.46 +/- 0.04 Ma. Map relations in Figure 11 show volcanic units at Bible back trending downslope to the northwest; however, these deposits are discontinuous remnants as mapped by Buffler (1967) and may represent debris downslope of the unit. The elevation of the flow base, where it was sampled for this study, is ~2550 m and is ~550 m above the modern Little

Snake River. The local exposed thickness of Browns Park Formation is approximately 450 m, suggesting 450-550 m of relief generated by incision.

The basalt flows at Battle Mountain and Squaw Mountain lie directly across the Little

Snake River from one another (Figure 11), are of essentially identical age, and are at broadly similar elevations. The relationship of these two basalt flows to the Little Snake River thus provides an opportunity for estimating the magnitude of fluvial incision along the Little Snake directly. Here, I assume that the ca. 11.5 Ma land-surface extended between Battle Mountain and

Squaw Mountain. Taking the average elevation of the two flow bases, ~ 2600 m, above the modern elevation of the Little Snake, ~2030 m, yields an estimate of fluvial incision of ~580 m since ca. 11.5 Ma. This direct reconstruction of fluvial incision is similar to the local exposed thickness of Browns Park and suggests that the preserved thickness of Browns Park Formation is a reasonable, conservative, estimate of incision along the Little Snake and Yampa Rivers.

Additionally, the constraints on exhumation at Bible Back Mountain are broadly similar to those at Battle and Squaw mountains. Thus, these three localities seem to represent a fairly consistent and robust estimate of local relief generation since basalt emplacement at ca. 11.5 Ma.

5.1.2 Black Mountain and Mt. Welba

In contrast, geologic relationships at Black Mountain and Mount Welba in the western

Elkhead Mountains (Figure 11) show a markedly different relationship between the local

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thickness of Browns Park Formation and their elevation above the modern river. At Black

Mountain, a fairly extensive flow, or very likely a vertical sequence of flows, are poorly exposed beneath thick vegetation. Extensive deposits of vesiculated, basaltic debris cover the area adjacent to and directly below the mesa-shaped topographic high, but exposures are rare and the base of the flow is buried. The new 40Ar/39Ar age determined from an outcrop discovered on the northeast end of the main ridge is 10.92 +/- 0.16 Ma (Table 3). The most basal expose of the flow, where the above sample was taken, is at an elevation of ~3160 m. The local exposed thickness of Browns Park Formation to the southwest of Black Mountain is ~350 m.

Nearby at Mt. Welba, exposures were also poor and difficult to access. There are three topographic highs in the vicinity of Mt. Welba (Figure 11). The southernmost point, Mt.

Oliphant, appears to be intermediate in composition and does not display noticeable flow textures. The easternmost point, Buck Point, was not easily accessed and was not described for this study. Mt. Welba, the northernmost point, has outcrops of weathered, vesiculated basalt.

Samples from an exposure on the southern edge of Mt. Welba yielded a new 40Ar/39Ar age of

12.60 +/- 0.06 Ma (Table 3). The basal exposure is at an elevation of ~3150 m and the thickness of exposed Browns Park Formation is ~400 m to the northeast of Mt. Welba.

The flows at Black Mountain and Mount Welba are ~500 m higher in elevation than Battle

Mountain yet sit atop a slightly thinner section of Browns Park Formation. This marked difference in elevation highlights the difficultly of using the Browns Park Formation to estimate the magnitude of exhumation (e.g., Buffler, 2003). Using these constraints, Black Mountain and

Mt. Welba record 350-1180 m and 400-1170 m of relief generated by incision respectively.

These values may reflect additional relief generated by late Cenozoic faulting (e.g. Buffler, 1967,

2003) and do not seem to represent the most reasonable estimates of relief generation in the region. Instead, here, I project the basal elevations of the flows at Black Mountain and Mt.

Welba to more local tributaries within the same fault block, Slater Creek/Elkhead Creek (Figure

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11); both with upper elevations of ~2500 m. This estimates incision to be ~350-660 m at Black

Mountain and ~400-650 m at Mt. Welba.

5.2 Reconstructing Exhumation in the Flat Tops

At the headwaters of both the Yampa and White Rivers (Figure 4), a laterally expansive sequence of at least 27 stacked basalt flows (Larson et al., 1975) make up the large, high elevation, mesas for which the Flat Tops Range (Figure 13) is named. Here, basalt flows have been previously constrained to range in age from approximately 23 to 9.6 Ma (Larson et al.,

1975; Kunk et al., 2002). Individual basalt flows range in thickness from 3m to ~60 m where locally ponded (Larson et al., 1975). Typically, however, flows are on the order of ~3-15 m thick

(from ~24-20 Ma) or ~15-23 m thick (from ~14-9 Ma) (Larson et al., 1975). The flows exhibit a dominantly flow on flow relationship to the southwest with increased inter-bedding of the

Browns Park Formation to the northeast with an overall thickness of these deposits of approximately 470 m (Larson et al., 1975) (Figure 12).

Basalts in the Flat Tops region appear broadly titled south approximately 3-4° (Larson et al., 1975). However, the sequence of stacked basalt flows broadly describes a region in which

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basalt flows ranging in age from approximately 23-10 Ma are deposited relatively conformably and lie within a few hundred meters elevation from one another. This relationship suggests that a low relief surface existed in the Flat Tops region from ~23 to 10 Ma. This relationship also allows for estimation of post ca. 10 Ma fluvial incision by assuming that the highest exposed basalts everywhere in the Flat Tops represent the approximate land surface at ~10 Ma.

I calculate the magnitude of fluvial incision on the uppermost headwaters of the Yampa and White rivers by averaging the highest elevation of the basalt surface on both sides of the modern valley and subtracting the elevation of the modern river channel. Here, I assume that prior to ~10 Ma these river valleys were filled with material up the elevation of the low-relief surface defined by the basalts that has been subsequently removed by fluvial incision. To date this surface, I use the youngest radiometric age determined in the Flat Tops (9.6 +/- 0.5 Ma –

Larson et al., 1975). Although this date is an old K-Ar age, the K-Ar ages from Larson et al.,

(1975) for the oldest basalt flows in the Flat Tops are consistent with newer 40Ar/39Ar ages from

Kunk et al., (2002).

For the headwaters of the White River (A-A’, Figure 13) from Lost Lakes Peak to Sable

Point, this reconstruction yields ~ 900 m of fluvial incision in the last 9.6 +/- 0.5 Ma. For the headwaters of the Yampa River (B-B’, Figure 13) from Orno Peak to Flat Top Mountain, this reconstruction yields ~ 700 m of fluvial incision within the last 9.6 +/- 0.5Ma. These estimates yield incision rates of 94 +/- 5 m/Ma and 73 +/- 4 m/Ma, respectively.

5.3 Reconstructing Exhumation in the Yampa River Valley

As the Yampa River flows out of the Flat Tops, it passes its namesake at the town of

Yampa, CO and then flows north in a fault bounded valley before making a series of sharp b- ends; east towards Woodchuck Hill, north parallel to the flank of the Park Range and eventually

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west at Steamboat Springs, CO (Figure 4 and Figure 14). Before the river turns north and parallels the Park Range, it flows through the ‘Yampa River Valley’. Here, the Yampa River flows over primarily the Mancos Shale and the Browns Park Formation, which has been intruded by many young dikes and volcanic necks (e.g., Kucera, 1962).

5.3.1 Woodchuck Hill

Near where the Yampa River makes a sharp turn to the east is a second basalt flow atop

Woodchuck Hill (Figure 14). The flow, or potential vertical sequence of flows, here is poorly exposed, but appears to be quite thick. At the top of Woodchuck Hill, the topography is expansive and approximately flat, suggesting the top of a flow surface. Here, a sample from a very poor outcrop yields a new 40Ar/39Ar age of 6.04 +/- 0.04 Ma (Table 3). Approximately 65 m lower in elevation, at ~2620 m, on the southwest side of Woodchuck Hill an outcrop of dark basalt with minor olivine and mildly developed flow banding is exposed. The new 40Ar/39Ar age from this sample is 5.97 +/- 0.06 Ma. These observations suggest that a ~60 m thick flow, or sequence of flows cap Woodchuck Hill. I use the age and elevation of the lower sample to constrain local relief generation to be ~460 m at a rate of ~77 m/Ma. The entirety of this relief appears to have been generated following basalt emplacement as Browns Park Formation is exposed continuously down to the Yampa River.

5.3.2 Lone Spring Butte

Near Yampa, CO at Long Spring Butte (Figure 14), a ~10 m thick porphyritic basalt flow with moderately well-developed flow banding is exposed. The basalt has phenocrysts of olivine, plagioclase, and mafic accessory minerals, which average ~0.5 mm in diameter and are ~2 mm in maximum diameter. The base of the flow, sampled by Andres Aslan, is at an elevation of ~3090 m (Aslan, pers. comm.); ~630 m above the modern Yampa River. The new 40Ar/39Ar age for the

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flow is 6.15 +/- 0.03 Ma (Table 3). The basalt flow atop Lone Spring Butte unconformably overlies Browns Park Formation which is dipping gently toward the southwest, ~20°.

Additionally, Browns Park Formation underlying the basalt flow is exposed continuously down to the river suggesting that the entire ~630 m of relief was carved subsequent to basalt eruption at

~6.15 Ma; at a rate of 102 m/Ma.

Volcanic deposits local to Lone Spring Butte also provide important constraints on the timing of exhumation along northern rivers in the study area. A tuff deposited near the top of

Lone Spring Butte (Figure 14) at an elevation of ~3010 m, ~80 m below the basalt flow, has a zircon fission track age of 23.5 +/- 2.5 Ma as described by Izett, (1975) and Luft, (1985). The relatively thin exposure of Browns Park Formation (~80m) preserved between the tuff (~25 Ma) and the basalt flow (~6 Ma) suggests the possibility that a significant amount of sediment was removed by erosion prior to the emplacement of the basalt flow atop Lone Spring Butte.

Deposits of volcanic breccia, previously described by Kucera (1962) and Buffler (1967), are also exposed in the Yampa River Valley. These deposits, referred to as the Crowner

Formation (Kucera,1962), and henceforth simply referred to as ‘Crowner deposits’, occur in a topographic depression on the southwestern slope of Lone Spring Butte ~300-400 m below the base of the basalt flow. The Crowner deposits at Lone Spring Butte contain very poorly sorted, subangular to angular, cobbles of volcanics and Browns Park Formation. The bedding of the

Crowner deposits is moderately well developed and is on the order of a meter thick. However, the thickness of these beds tends to very, and the internal structure of the beds is poorly developed. The orientations of beds dip concentrically inward in a ring-like geometry. These observations suggest that the Crowner deposits likely represent maar deposits (Buffler, 1967).

Thus, at Lone Spring Butte, it is likely that these units were deposited close to the position of the ancestral-land-surface.

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New 39Ar/40Ar dates constrain the age of the Crowner deposits at Lone Spring Butte. A new 39Ar/40Ar age for a basaltic clast, contained within bedded Crowner deposits is 7.0 +/- 0.4 Ma and a new 39Ar/40Ar age for a mafic dike that cross-cuts the Crowner deposits is 4.62 +/- 0.05

Ma. The age of the dike (4.62 +/- 0.05 Ma) cross-cutting the Crowner deposits provides an absolute minimum age for the Crowner deposits. Notably, the age of the basaltic clast (7.0 +/-

0.4 Ma) is similar to the age of the basalt flow atop Lone Spring Butte (6.15 +/- 0.03 Ma) and may suggest that the Crowner deposits are approximately coeval with volcanism at ca. 7 – 6 Ma.

If the Crowner deposits are similar in age to the basalt flow capping Lone Spring Butte, then the significant, ~300-400 m, of relief between the flow and the Crowner deposits suggests the presence of significant relief generated by ca. 6 Ma. An alternate interpretation is that relief generation post-dates 6 Ma; with the Crowner deposits being closer to ~4.6 Ma in age.

Collectively, however, the relationships observed between the basalt flow atop Lone Spring

Butte, the underlying ash, and the Crowner deposits seem to suggest the onset of relief generation beginning prior to, and certainly not much later, than ~ 6 Ma in the Yampa River Valley.

5.4 Summary of Constraints on the Timing and Magnitude of Incision

Local relationships between volcanic features with new 39Ar/40Ar ages from this study

(Table 3) and the Browns Park Formation provide new constraints on the timing and magnitude of incision along northern rivers (White, Yampa, Little Snake). Examination of new incision rates along the Yampa and White rivers show relative agreement between the ca. 10 Ma time- averaged incision rates (~70-100 m/Ma) and ca. 6 Ma time-averaged incision rates (~77-102 m/Ma) (Figure 15B) suggesting that incision began after 10 Ma and prior to 6 Ma. This is broadly confirmed by field relationships documented for this study observed at Lone Spring

Butte. Although, the timing of incision along the Little Snake River (Figure 15A) is more

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difficult to constrain given the absence of incision markers between ~10-1 Ma, the regional minimum age of the Browns Park Formation and the absence of significant thicknesses of the

Formation observed that post-date ~8 Ma (e.g., Buffer, 2003) also suggest that incision here began between ~8-10 Ma. Collectively, these observations suggest that incision along northern rivers, probably began sometime between ~8-10 Ma, but certainly not later than 6 Ma; in general agreement with the timing of incision along the Colorado River (e.g., Aslan et al., 2010;

Karlstrom et al., 2012). New constraints from this study show that the magnitude of incision along northern rivers has been on the order of ~500-900 m in the last ca. 12 Ma (Figure 16 and

Table 4). In contrast, previously published estimates along the Colorado River show that incision has been on the order of ~1000-1500 m in the last ca. 10 Ma (e.g., Larsen et al., 1975; Kunk et al., 2002, Aslan et al., 2010).

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6 Discussion

6.1 Extensional Faulting in the Sand Wash Basin

Deposits of the Browns Park Formation in the Sand Wash basin have been displaced along north to northwest-trending normal faults (Figure 11). Of importance to this work, two explanations have been proposed to explain the present distribution of the Browns Park

Formation. Buffler (2003) suggested that the Browns Park Formation was deposited as a relatively uniform blanket of sediment such that variations in the present distribution of the

Browns Park (Figure 4) are the result of post Browns Park faulting and erosion. However, other workers have suggested that the Browns Park Formation was deposited in pre-existing areas of low elevation, either in structural grabens or erosional topography (e.g., Izett 1975; Hansen,

1986).

In the evaluation of relief generated by fluvial incision presented in section 4, I have tried to be conservative by projecting datums represented by basalt flows to the nearest tributary stream. In nearly all sites throughout the northern part of the study area, this analysis was confined to a single fault block. However, in the Elkhead Mountains, basalts of similar age

(~11.5 Ma) sit at elevations that differ by ~700 m. Basalt flows on Battle Mountain, Squaw

Mountain, and Bible Back Mountain sit at elevations around ~2600 m, whereas basalt flows on

Black and Welba Mountains sit at ~3150m elevation. Notably, these flows cap peaks in the footwall block of one of the two major graben-bounding faults (Figure 11). This relationship suggests the possibility that late Cenozoic faulting has influenced the evolution of topographic relief in the Elkhead range.

The timing and magnitude of displacement along late Cenozoic faults in the Elkhead

Mountains that cut the Browns Park Formation, however, is not well constrained. Motion along these faults has been previously interpreted to post-date a significant portion of Browns Park 60

deposition (Buffler, 1967). Deformation of structure contours along the basal conglomerate appears be mimicked by thickness variations in the upper member of the Browns Park Formation, and displacement along graben-bounding faults is thought to be on the order of ~300-600 m

(Buffler, 1967).

If fault-generated relief existed between the high and low elevation basalts prior to the onset of fluvial incision, then projecting the elevations of Black Mountain and Mt. Welba across documented normal faults (Figure 11) to the Little Snake River is not a reasonable estimate of incision. A more reasonable estimate would be to project these elevations to the head of Slater

Creek (Little Snake watershed) or Elkhead Creek (Yampa watershed); both of these flow presently at ~2500 m within the same fault block as the basalts (Figure 11). Utilizing this approach suggests that local relief generated by fluvial incision ranges from ~350-660 m at Black

Mountain and ~400-650 m at Mt. Welba. Notably, this estimate is similar to the maximum preserved thickness of the upper member of the Browns Park Formation mapped in this region.

Thus, although late Cenozoic normal faulting appears to play a role in displacing the Browns Park

Formation and associated volcanic deposits relative to one another, the estimates of fluvial incision presented here are not significantly influenced by this effect.

Late Cenozoic faulting has also been documented within the Flat Tops region and the

Yampa River Valley (Figure 13 and Figure 14 respectively) in the southern part of the study area

(Kucera, 1962). In the Flat Tops region, Kucera (1962) documented that late Cenozoic displacement along faults was minimal, apparently less than about 100 m. In the Yampa River

Valley, however, a number of the faults have Laramide ancestry, but late Cenozoic displacement appears to exceed ~600 m (Kucera, 1962). At Lone Spring Butte (Figure 14), the basal conglomerate of the Browns Park Formation is exposed beneath the 23.5 +/- 2.5 Ma tuff at dated by Izett et al. (1975). As the Browns Park conglomerate here is ~600 m higher elevation than in the Yampa Valley, the age of the tuff, places a maximum age constraint on this deformation.

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However, tilting of the upper Browns Park sandstone beneath the basalt flow at Lone Spring

Butte and throughout the region suggests a component of younger deformation (e.g., Buffler,

1967; Aslan et al., 2010).

6.2 Regional Correspondence between Late Miocene Incision and Channel Steepness

One of the notable results of this study is that the Colorado River exhibits a steeper profile than rivers in the north. Along the western slope, the lower reach of the Colorado is

0.9 relatively steep (ksn = ~90 – 100 m ), whereas the Little Snake is significantly gentler (ksn = ~40

0.9 0.9 0.9 m ). The White (ksn = ~80 m ) and Yampa (ksn = ~70 m ) Rivers are intermediate in both geographic distribution and normalized steepness (Table 2). To account for the potentially complicating effects of differences in lithology in these average values, I evaluate the mean channel steepness (ksn) of reaches that are underlain by similar substrate (Figure 17). I focus on two primary rock types – Tertiary sandstones, which include the Wasatch and Uinta Formations, as well as the Brown’s Park Formation, and Cretaceous shales (Lewis and Mancos Formations).

In a recent study of rock strength, Tressler (2011) found that variations in compressive strength among the former group are minimal. Compressive strength of Cretaceous shales was unable to be determined, due to the overall mechanical weakness of these units (Tressler, 2011), but we assume that variations across the study area are minimal. Therefore, comparison of channel steepness indices along these reaches should reflect differences in stream profile gradient that are independent of substrate erodibility.

The Colorado River exhibits relatively high values of ksn in Browns Park and equivalent

0.9 0.9 sediments (ksn = 81.6 +/- 38.5 m ), within the Wasatch Formation (ksn = 107.3 +/- 39.0 m ),

0.9 and, notably, within the Mancos Shale (ksn = 82.6 +/-14.8 m ). In contrast, the profile of the

Little Snake River, is approximately half as steep within the Browns Park and equivalent

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0.9 sediments (ksn = 36.8 +/- 0.1 m ), and nearly three times less steep within the Wasatch

0.9 Formation (ksn = 36.5 +/- 6.6 m ). This analysis provides compelling evidence that substrate lithology is not the dominant control on channel steepness across the study area.

One exception to the absence of a correlation between steepness and lithology is within reaches of crystalline Precambrian rocks. Along the Gunnison, Colorado and Little Snake Rivers, reaches underlain by crystalline, Precambrian bedrock often coincide with vertical step knickpoints that are associated with locally elevated ksn values. I interpret these correlations as indicative of locally resistant substrate (e.g., Tressler, 2011; Pederson and Tressler, 2012) and thus do not include these reaches in my regional comparison.

The Yampa River as it flows through the Dinosaur Canyon knickzone (Figure 5-1B), was also excluded from regional comparison. Here, locally resistant substrate (e.g., Darling, 2009;

Pederson and Tressler, 2012), input of coarse debris (Grams and Schmidt, 1999) or ongoing late

Cenozoic faulting (Hansen, 1986) all have the potential to locally influence channel steepness.

Furthermore, even if the Dinosaur knickzone does represent a transient feature (e.g., Darling et al., 2012) it would not represent a signal of regional base-level lowering. Finally, the uppermost

~50 km of both the Yampa River and White River were excluded, as coarse debris being shed off of the Flat Tops (Larson et al., 1975) could be locally steepening the headwaters.

Collectively, however, these observations suggest that lithology is not the primary control on the steepness of the study rivers; nor is differences in discharge between rivers along the western slope (c.f., Pederson and Tressler). Furthermore, the steepest channels in the study area correlate with the largest magnitude of incision over the last ~10 Ma (Figure 17 and Figure 16).

0.9 The lower reach of the Colorado displays a ksn of ~90 – 100 m and has incised ~1500 m in the last ~10 Ma, while the Little Snake River has a normalized steepness of ~40 m0.9 and has only incised ~450-650 m in the last ~10 Ma.

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7 Potential Drivers of Late Miocene Incision

Late Cenozoic climate change, base-level fall during drainage basin integration, and differential rock uplift have all been proposed as possible drivers of late Miocene exhumation along the western slope of the Colorado Rockies. The combination of channel profile analysis and new constraints on the timing and magnitude of fluvial incision presented here demonstrate that 1) channel profile steepness of major river systems decreases from south to north along the western slope (Figure 17), 2) differences in profile steepness are independent of both discharge

(c.f., Pederson and Tressler, 2012) (Figure 8) and substrate lithology (Figure 17), and 3) that the steepest rivers have experienced the greatest amount of late Cenozoic incision (Figure 17 and

Figure 16). In this section, I consider what potential driving mechanisms best explain the correspondence of steep channels and deep incision across the study area.

7.1 Enhanced Fluvial Incision in the Late Miocene

One of the potential explanations for driving late Miocene incision along the western slope of the Rockies is the possibility that climatic changes during the late Miocene enhanced the potential for fluvial transport, either through an increase in storminess (e.g., Molnar, 2004) or increased mean discharge from snow melt (Pelletier, 2009). Apparent increases in global sedimentation rates between 2-4 Ma have been cited as evidence for an increase in the efficacy of fluvial erosion (e.g., Zhang et al., 2001), although these findings have recently been called into question (Willenbring and von Blanckenburg, 2010).

Along the western slope, evidence for increases in incision rate is limited, however.

Along the Colorado River, the key marker often cited as evidence for an increase in Pliocene incision rates is the basalt flow at Gobbler’s Knob (Kunk et al., 2002). As argued previously

(Aslan et al., 2008; 2010), the relationship of this flow to the position of the river is uncertain and 65

the high incision rate (~240 m/Myr, Kunk et al., 2002) may be an artifact. In contrast, if one considers basalt flows of known association to the position of river gravels and inset fluvial terrace/fan complexes (Figure 15C), incision rates along the Colorado River appear to be relatively constant with time. Along northern rivers, markers of younger age are sparse, but the data appear to admit relatively constant incision along these rivers as well during the past ~10 Ma

(Figure 15A, B). Although it is possible that slightly elevated rates of incision over the past ca.

640 Ka (Dethier, 2001) reflect climatic enhancement, overall, these results do not seem to require late Cenozoic increases in incision rate.

Regional patterns in the amount of fluvial incision and channel steepness further argue against climate change as the primary driver of incision along the western slope. The Colorado

River has experienced the greatest magnitude of incision in the last ~10 Ma, but yet remains the steepest of the rivers. Models of profile response to climate change (e.g., Wobus et al., 2010) suggest that increases in discharge and/or storminess which drive fluvial systems toward more efficient transport are associated with reduction of steady-state channel gradients (Whipple and

Tucker, 1999; Wobus et al., 2010). Thus, it remains difficult to reconcile steep channels and large magnitude incision along the Colorado River; some additional process must maintain steep gradients in the face of incision. Moreover, it seems even more unlikely that climatic enhancement of fluvial systems could simultaneously drive ~1500 m of incision along the

Colorado River while only resulting in ~500 m of erosion along the Little Snake River. These rivers are only a few hundred kilometers apart, have headwaters at broadly similar elevations, and experience similar discharge-area relations today. Overall, these results appear to rule out climate change as a significant driver of spatial variations in late Miocene incision and present-day profile shape in western Colorado.

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7.2. Base-level Fall and Transient Incision Associated with Basin Integration

Transient incision in response to drainage integration has long been thought to be a primary mechanism driving incision and canyon development across the Colorado Plateau (e.g.,

Hunt, 1956; Pederson, 2002). Although the present position of Grand Canyon appears to exploit an ancient paleocanyon (e.g., Flowers et al., 2007; Wernicke, 2011; Flowers et al., 2012), evidence that final integration of the Colorado River through the Grand Canyon to the Gulf of

California occurred around ~6 Ma is unequivocal (e.g., Lucchitta, 1990; Dorsey et al., 2007).

Given that incision along the western slope appears to initiate prior to this time - at ~10 Ma within the Colorado River (Aslan et al., 2010; Karlstrom et al., 2012; this study) and within northern rivers at ~8-10 Ma (this study) - transient incision associated with the final integration of

Grand Canyon cannot be the primary driver for the initiation of incision along the western slope.

Rather, the data presented here favor scenarios that consider that the upstream effects of transient incision associated integration of the Colorado River through Grand Canyon have not yet progressed upstream of the middle reaches of the Colorado River (Wolkowinsky and Granger,

2004; Karlstrom et al., 2008; Cook et al., 2009; Darling et al., 2012).

These results do not rule out a role for a possible, older drainage integration event upstream of Lee’s Ferry, however. The presence of ~1500 meters of relief that developed between 35 Ma and 16 Ma in the southern Colorado Plateau (Flowers et al., 2007; Cather et al.,

2008) may point to the need for a paleo-drainage divide somewhere to the south of the present day Book Cliffs. It is possible that breaching of that divide led to incision along the upper

Colorado River and Green River systems. However, if this scenario occurred, integration must have pre-dated final integration of the Colorado River through Grand Canyon at ca. 6 Ma.

Although data from this study seem to rule out incision driven by integration through Grand

Canyon, they leave open the intriguing possibility that integration of the upper Colorado River

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was achieved through a complicated, and long-lived series of integration events; possibly related to the upstream retreat of knickpoints from the lake Bidahochi area between ~16-10 Ma. This is an intriguing possibility and will require additional future study.

Because not much is known about the timing of breeching across the Book Cliffs and the integration of the Green River to the Colorado, the data presented here could be explained by an early integration event in which the Colorado River along the western slope was a fully integrated drainage, pre-dating an integrated Green River. It has been hypothesized that the Green River was relatively recently integrated into the Colorado watershed across the Uinta Mountains

(Hansen, 1986); recent dating of high terraces in the Green River basin, downstream of this point, suggest this event occurred sometime after ~8 Ma (Hansen, 1986) and well before 1.2 Ma

(Darling et al., 2012). It is likely that this integration even explains the ~100-200 m of relief across the knickzone along the Yampa River profile (Figure 9E). Prominent knickzones are observed along the Green River at Desolation Canyon in the Book Cliffs, through the Uinta

Mountains in Lodore Canyon, and along the Yampa River through Dinosaur Canyon (Figure 16).

Both knickzones along the Green River, at Desolation Canyon (Pederson et al., 2012) and Lodore

Canyon (Darling et al., 2012), have been interpreted as transient features that separate regions of downstream incision from upstream reaches that have not yet experienced deep incision.

Importantly, given the modern drainage configuration, the hypothesis that differences in the amount of incision along the Colorado River (~1500 m) and the White/Yampa/Little Snake

(~500-900 m) reflect a wave of incision that has propagated upstream along the Colorado River, but has not yet reached the northern tributaries (e.g., Pederson and Tressler, 2012) requires that transient incision is stalled across the knickzone along the Green River through the Book Cliffs

(Desolation and Grey Canyons, Figure 16). There are two problems with this hypothesis. First, the drop in elevation along the Green River through the Book Cliffs is < 200 m, and thus there is not enough relief along the steepened reach of the profile to explain a differential incision of the

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magnitude required (~600-1000 m). One would also have to claim that ~1000 m of material had been excavated from across of the Unita basin, north of the Book Cliffs.

The second problem with the hypothesis of base level fall is that it requires nearly simultaneous incision in both the headwaters of the Colorado River as well as in the Little Snake

River. As shown here, the best estimates of the onset of fluvial incision along both systems are

~8-10 Ma. If differences in incision reflect base level fall, it would require a scenario by which nearly instantaneous propagation of an initial wave of incision made its way throughout the entire system. For unknown reasons, this incision would have had to continue, at relatively constant rate, along the Colorado River, but stalled at the Book Cliffs along the Green River.

7.3 Differential Rock Uplift and Tilting along the Western Slope

Because climatically enhanced incision or basin integration does not seem sufficient to explain the patterns of fluvial incision and channel steepness along the western slope of the

Colorado Rockies, it appears to leave open the possibility of differential rock uplift between the range and the Colorado Plateau (e.g., Karlstrom et al., 2012; Darling et al., 2012). The relationship of steep channels and large magnitude incision, certainly, is consistent with this hypothesis, as we expect such relationships in systems adjusted to spatial variations in rock uplift

(e.g., Kirby et al., 2003). In the Colorado Rockies, moreover, the spatial correspondence between steep, rapidly incising rivers and low seismic velocity mantle (Karlstrom et al., 2012) suggests an association between the ‘Aspen anomaly’ and late Miocene exhumation.

North-south variations in both channel steepness and the magnitude of late Cenozoic incision from the Colorado to the Little Snake River appear to correlate well with the margin of the ‘Aspen anomaly’ (Figure 1). The Colorado River appears to have incised ~1500 m (perhaps decreasing upstream to ~1100 m, Kunk et al., 2002), the White and Yampa rivers have incised

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~700 – 900 m, and the Little Snake River has incised ~450 -650 m (Figure 16). All of this incision likely occurred during the past ~8-10 Ma. Additionally, the lower reach of the Colorado

River displays the highest values of normalized steepness, ~90, and the Little Snake River displays the lowest values of normalized steepness, ~ 40; the White River and Yampa River show intermediate values of normalized steepness (Figure 17, Table 2). This suggests that as these rivers begin to progressively overlie lower velocity mantle at depth they display an increase in incision and modern channel steepness. Qualitatively, at least, this is what one would expect if low velocity mantle is associated with a broad welt of surface deformation; tilting should occur toward the west-northwest across the flank of the anomaly.

The form of channel profiles along the western slope potentially also suggest an adjustment to headwater rock uplift. The Colorado River displays a steep lower reach that transitions to a lower gradient reach above Gore Canyon (Table 2). This transition, interestingly, occurs where the Colorado River crosses the largest gradient in P-wave velocities at 100 km depth (Figure 1). Likewise, the White River and Yampa River may also be responding to a forcing in their headwaters. Examination of the local normalized steepness values (Figure 9E and

9F) reveals increases in local ksn towards the headwaters of these rivers. In fact, in the uppermost,

~50 km, of their profiles, both the Yampa River and White River display much higher values of local normalized steepness, average of ~90 m0.9 (Figure 9E and 9F). As noted previously, although these steep reaches could be the result of coarse debris being shed off the Flat Tops, it is an intriguing possibility that they in fact represent a signal of greater differential rock uplift toward their headwaters, toward the center of the Aspen Anomaly. I suggest that it may be that the channel profiles are still responding to a pulse of uplift that occurred in the last 10 Ma and have not yet fully adjusted.

As knowledge of how the ‘Aspen anomaly’ varied in magnitude or position over time is uncertain, causal correlations between record of incision and the underlying mantle structure are

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speculative. I acknowledge that isostatic rebound is a potentially important component of the modern surface elevation (e.g., Molnar and England, 1990; Pederson et al., 2002). Although rebound is undoubtedly an important component of any surface uplift along the western flank, the results of this study suggest that the initial driver of ca. 10 Ma incision prior to any associated isostatic rebound likely involved a component of tectonically driven rock uplift. It is possible, that the initial signal of uplift was relatively short in duration and that isostatic rebound has been contributing to overall surface elevations as incision and mass removal progresses. However, studies that model the predicted isostatic response of the region have suggested that the headwaters of the Colorado River show high ‘residual incision’ (magnitude of modeled isostatic uplift subtracted from documented fluvial incision) relative to further downstream near the confluence with the Green (Karlstrom et al., 2012) further suggesting that there has been uplift in excess of rebound along the Colorado-Rocky Mountain Plateau transition. Overall, the results of this study appear to be consistent with the potential for several hundred meters of rock uplift in excess of isostatic adjustment; similar to the magnitude and wavelength observed along the eastern slope of the Rockies (e.g., McMillan et al., 2002; Leonard, 2002).

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8 Conclusions

New geologic data from the headwaters of the White, Yampa, and Little Snake rivers constrain the magnitude and timing of fluvial incision along the western slope of the Colorado

Rockies, combined with analysis of channel profiles, provide new insights into the history and potential drivers of late Miocene exhumation and lead to the following conclusions.

1. The Colorado River is approximately twice as steep as the Little Snake River, whereas

the Yampa and White rivers are intermediate between these two end members.

2. These differences in steepness persist within rock units of similar erodibility (this study)

and do not appear to reflect differences in discharge (c.f., Pederson and Tressler, 2012).

3. Incision along the White, Yampa and Little Snake rivers post-dates 10 Ma and most

likely pre-dates 6 Ma.

4. Variations in modern profile shape coincide with spatial variations in the magnitude of

late Cenozoic incision and exhumation. Incision along the Colorado River began ~8-10

Ma and the river has incised at least ~1500 m since then. Over the same time period,

incision along northern rivers appears to have only reached ~700-900 m along the

White/Yampa and ~450-650 m along the Little Snake.

This combination of regional patterns in channel steepness and the magnitude of late Cenozoic erosion along the western slope of the Rockies appears to be best explained by differential rock uplift associated with low-velocity mantle beneath western Colorado. Incision associated with more efficient fluvial transport due to late Cenozoic climate change does not appear to explain the similar timing but spatially variable incision. Relative base-level fall during basin integration could explain the patterns, but then it must be an early event, pre-dating integration of the

Colorado River through Grand Canyon, and would seem to implicate a more complicate integration of the upper Colorado than previously recognized. Collectively, the association

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between steep channels, deep exhumation, specific zones of high steepness along the Colorado,

White, and Yampa rivers, and low velocity mantle at depth appears to implicate differential rock uplift during the past ~10 Ma driven by changes in the buoyancy structure of the mantle lithosphere as the best explanation for late Miocene – recent incision in the upper reaches of the

Colorado River and tributaries of the Green along the western slope of the Rockies.

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Appendix: 40Ar/39Ar Analytical Methods and Results

The 39Ar/40Ar age determinations for this study were provided by Matt Heizler at New

Mexico Tech University. The following detailed description of methods and analytical techniques used were also provide by Matt Heizler and are a modified excerpt from the New

Mexico Geochronology Research Laboratory internal report #: NMGRL-IR-771:

Groundmass concentrates were prepared from basaltic samples by choosing fragments visibly free of phenocrysts whereas biotite or sanidine was obtained by standard mineral separation procedures. The prepared samples were irradiated in three batches; either for 10 hours or for one hour at the USGS TRIGA reactor in Denver, CO along with the standard Fish Canyon tuff sanidine as a neutron flux monitor. Most samples were analyzed by the step-heating method using a defocused CO2 laser to heat the samples (Tables A-1 – A-4). The age of the Sand

Mountain Sample was determined by probability distribution of individual sanidine grain total fusion ages (Figure A-5). A summary of the preferred eruption ages along with a listing of the analytical methods is provided in Table A-1 and Table A-2 and the general operational details for the NMGRL can be found at internet site http://geoinfo.nmt.edu/publications/openfile/argon/home/html.

The sample age spectra are defined by 8 to 12 heating steps and each sample provides either a plateau or isochron age that range between ~4.6 and 12.6 Ma (Tables A-1 – A-4; Figures.

A-1 – A-4). Groundmass samples typically record age spectra (Figures A-1 and A-2) with an initial non-radiogenic step that is often discordant (younger and older) from the remaining steps that are themselves somewhat scattered. Isochron analysis demonstrates that for many age spectra the discordance is explained by trapped excess 40Ar (Figures A-3 and A-4). The preferred age for each sample is given by the method (weighted mean or isochron) that in most cases yielded the lowest MSWD for the chosen steps and contained the great part of the spectrum. This is

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summarized in Table A-1 and Table A-2 and labeled either plateau or isochron on each age spectrum (Figures A-1 and A-2). Regression values for the isochrones are given by the York

(1969) method. The biotite spectra are overall flat, however isochron data suggest minor excess argon contamination and therefore the isochron age is chosen as the preferred age.

For most age spectrum analyses the majority of gas released yields well-defined plateau and/or isochron results and therefore the preferred ages are confidently assigned as eruption ages.

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