VOLCANIC INFLUENCE OVER FLUVIAL SEDIMENTATION IN THE CRETACEOUS McDERMOTT MEMBER, ANIMAS FORMATION, SOUTHWESTERN COLORADO

Colleen O’Shea

A Thesis

Submitted to the Graduate College of Bowling Green State University in partial fulfillment of the requirements for the degree of

MASTER OF SCIENCE

August: 2009

Committee:

James Evans, advisor

Kurt Panter, co-advisor

John Farver

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Abstract

James Evans, advisor

Volcanic processes during and after an eruption can impact adjacent fluvial systems by

high influx rates of volcaniclastic sediment, drainage disruption, formation and failure of natural

dams, changes in channel geometry and changes in channel pattern. Depending on the magnitude

and frequency of disruptive events, the fluvial system might “recover” over a period of years or

might change to some other morphology. The goal of this study is to evaluate the preservation

potential of volcanic features in the fluvial environment and assess fluvial system recovery in a

probable ancient analog of a fluvial-volcanic system. The McDermott Member is the lower

member of the Late Cretaceous - Tertiary Animas Formation in SW Colorado. Field studies were based on a southwest-northeast transect of six measured sections near Durango, Colorado. In the

field, 13 lithofacies have been identified including various types of sandstones, conglomerates, and mudrocks interbedded with lahars, mildly reworked tuff, and primary pyroclastic units.

Subsequent microfacies analysis suggests the lahar lithofacies can be subdivided into three types based on clast composition and matrix color, this might indicate different volcanic sources or sequential changes in the volcanic center. In addition, microfacies analysis of the primary pyroclastic units suggests both surge and block-and-ash types are present. Several trends can be noted: (1) there is an overall fining-upward trend seen throughout the McDermott Member in a transition from lahars and fluvial conglomerates at the base to isolated sandstone-rich channels and extensive mudrock deposits near the top of the unit, (2) there is a lateral trend change in the grain size and thickness of lahars, and the thickness of well-developed fluvial deposits, and (3) there is evidence of drainage disruption following the deposition of lahar deposits by the iii

deposition of fine-grained silt/mudstones and sandstones which may indicate the creation of

natural dams. Stratigraphic relationships and paleoflow directions suggest lahars and other volcanic sediments were transported SSE from the La Plata Mountains through one or more paleovalleys. In comparing modern and ancient volcanically-influenced environments, it is evident the thin sheets of volcanic ash which dominate modern environments have low preservation potential. Rather, the volcanic signature in the geologic record consists of: (1) lahars interbedded with fluvial deposits, (2) volcaniclastic sediment such as reworked tuff or tuffaceous sandstone, and (3) paleocurrent and provenance data.

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Acknowledgements

I’d like to thank Phil Wheeler, the Terry Palmer Ranch, and Steve Whiteman from the

Southern Ute Indian Tribe for allowing me access to the McDermott Member from their

properties and/or gates. Thank you to both Dr. Jim Evans and Megan Castles who aided me in

field work in Durango. This research could not have been done without financial support from

the Bowling Green State University Geology Department and from the Katzner and University

Bookstore Funds for Graduate Research and Professional Development Fund. I’d also like to

thank my committee, Dr. Jim Evans, Dr. Kurt Panter, and Dr. John Farver for all their input and time.

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TABLE OF CONTENTS

Page

INTRODUCTION ...... 1

CHAPTER I. GEOLOGICAL BACKGROUND...... 11

CHAPTER II. METHODS ...... 19

CHAPTER III. RESULTS...... 21

Facies Analysis...... 21

Petrographical Observations...... 30

Paleogeography ...... 32

CHAPTER IV. DISCUSSION...... 37

Depositional Environments...... 37

Volcanic Inputs on Fluvial Environment...... 39

Signature of Volcanic Events in the Geologic Record ...... 44

CHAPTER V. SUMMARY & CONCLUSIONS...... 49

Summary ...... 49

Conclusions ...... 50

REFERENCES ...... 78

APPENDIX A. SAMPLE LOCATION DESCRIPTIONS ...... 84

APPENDIX B. PETROGRAPHY DATA...... 87

APPENDIX C. PALEOCURRENT DATA ...... 89 vi

LIST OF FIGURES

Figure Page

1 Proximal and Distal Volcaniclastic Environments ...... 52

2 Study Area Map of Durango, CO ...... 53

3 Cretaceous and Tertiary Rocks of Durango, CO ...... 54

4 Biostratigraphy of Late Cretaceous Rocks in Study Area ...... 55

5 Previous Paleocurrent Data of Animas Formation ...... 56

6 Geologic Map of Bedrock Units in Durango, CO ...... 57

7 Clast Count Histograms ...... 58

8 Stratigraphic Column of Section 1...... 59

9 Stratigraphic Column of Section 2...... 60

10 Stratigraphic Column of Section 3...... 61

11 Stratigraphic Column of Section 4...... 62

12 Stratigraphic Column of Section 5...... 63

13 Stratigraphic Column of Section 6...... 64

14 Photos of McDermott Member Lithofacies ...... 65

15 Photos of McDermott Member Lithofacies 2 ...... 66

16 Photos of McDermott Member Thin Sections...... 67

17 Petrography of McDermott Member ...... 68

18 Paleocurrent Rose Diagram of McDermott Member...... 69

19 Vertical Stratigraphic Columns of Sections 2 and 4...... 70

20 Photomosaic of Outcrop at Section 6...... 71

21 Representation of Fluvial Recovery within the McDermott Member ...... 72 vii

22 3-D Representation of Lahar Deposition near Volcanic Center ...... 73

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LIST OF TABLES

Table Page

1 Palynomorph Taxons Used for Dating in CO ...... 74

2 Fossils Used for Dating in McDermott Member ...... 75

3 Radiometric Ages of Volcanic and Plutonic Rocks...... 76

4 Lithofacies of McDermott Member...... 76

5 Statistical Analysis Data of Petrographic Data...... 77

1

INTRODUCTION

Volcanic Processes

It is widely acknowledged that volcanic processes can effect adjacent depositional

environments. For example, volcaniclastic deposits up to 3 km thick found within the Cascade

Range in California suggest large sediment loads imposed on fluvial systems by volcanic

processes (Smith, 1987). As another instance, the remobilization of pyroclastic materials by fluvial processes following the 1800 Taupo eruption in New Zealand caused shallow reworking of fluvial channels and deep incision of hillslope rills and gullies (Manville et al., 2004). A further example, the Pliocence Mushono tephra beds in Japan, suggests that overloading of fluvial systems by pyroclastic debris led to the development of low-sinuosity channel systems

(Kataoka, 2005).

Volcanic environments can be divided into proximal areas (affected by pyroclastic surges, blast and shock waves, bombs, flows or lahars) and distal areas (affected by lahars and volcanic ash) (Fig. 1). Some of these volcanically derived processes, such as lahars, can have great impacts on surrounding environments, in particular river systems.

Lahars can be defined as rapidly flowing, volcaniclastic, sediment gravity flows originating on or near a volcano during or shortly following an eruption (Giordano et al., 2002).

Lahars can occur as either “primary” or “secondary” flows. Primary lahars are those that are directly triggered by an eruption, while secondary lahars occur post-eruption and can be caused by intense rainfall on hillslopes that have been devegetated and/or accumulated large amounts of volcanic ash (Canuti et al., 2002).

Primary and secondary lahars have different properties evident in their deposits. For example, a series of secondary lahars generated by rain can form sediment aprons that onlap 2 older (pre-event) erosional surfaces cut into the original material (Giordano et al., 2002). These aprons are typically composed of several, low-aspect-ratio (ratio of width-to-height) aggradational deposits (Giordano et al., 2002). In contrast, primary lahar deposits display vertical and lateral grading of grain-size which reflects rapid processes of coupling/decoupling of water and particles and the amount of water available (Giordano et al., 2002). Such deposits tend to be more lobate in shape, and have high-aspect ratios (Giordano et al., 2002).

When an eruption occurs, ash can destroy vegetation, smooth topography, and form a thin, impermeable crust (Canuti et al., 2002). The ash layer can thus reduce rainfall infiltration, cause accelerated runoff and erosion, and generate secondary lahars, as was seen at Mount

Pinatubo in the Philippines during 1991 (Canuti et al., 2002). In addition, the introduction of large quantities of ash into a river changes sediment loads, flow behavior, and may damage wildlife and aquatic flora. Secondary lahars may be triggered by rainfall on slopes that have been steepened and/or devegetated by ashfall deposits. Once triggered, large quantities of ash are eroded, mixed with water, and flow down stream channels. The flows are capable of eroding rocky debris from streambeds. For example, one of the lahar deposits associated with the

Pichincha volcano in Ecuador, which was deposited within the last 10,000 years, contains a significant abundance of entrained lithics (fluvial sediment) which substantially increased the volume of the lahar (Canuti et al., 2002).

Volcaniclastic Effects on Adjacent Environments

Volcanic centers are commonly surrounded by fluvial systems. The impact of volcanic eruptions on fluvial systems include: flooding, drainage disruption including the formation and failure of natural dams, modifications of channel morphology and type, and high sediment loads. 3

Drainage basins impacted by volcanic eruptions can produce extremely high sediment yields

because of the combination of increased run-off and increased soil erosion from higher slopes,

which may be enhanced by damage to stabilizing vegetation (Hayes et al., 2002). The post-

eruption transport of high volumes of sediment can create extremely hazardous conditions such

as unstable hillslopes with potential landslides. Pyroclastic surges and lahars can alter the

geometry of the fluvial drainage system (Hayes et al., 2002), for example, surplus volcaniclastic

sediment introduced into a fluvial system may change a meandering stream to a braided stream

(Kataoka, 2005). At the terminus of a river system, the increased sediment load can result in

delta progradation (Kataoka, 2005). Post-eruption pulses of volcaniclastic sediments with high

sediment yields can also affect distal regions of relevant river systems by controlling the

depositional pattern on alluvial plains and deltas (Kataoka, 2005). Typically, the impact of lahars

persists many years after an eruption. For example, suspended sediment loads in the area

draining the Mount St. Helens lahar remained two orders of magnitude greater than pre-eruption

loads for almost 20 years after the 1980 eruption (Gran and Montgomery, 2005).

Over a time interval, volcanically-influenced fluvial deposits may display evidence of

syneruption, inter-eruption and post-eruption periods. Syneruption intervals are characterized by

high sediment loads and peak discharges. They are relatively short in duration compared to inter-

eruption periods (Smith, 1991). Syneruption periods are characterized by laterally extensive

flood deposits as well as interbedded debris flow deposits, or debrites. The volcanic component

of these deposits tends to have a monolithic composition and is more sand-rich (Smith, 1991).

Syneruption deposits are bounded by numerous erosional surfaces and may overlie previously

developed paleosols (Smith, 1991). Post-eruption periods are transitional between syneruption and inter-eruption periods. These transition periods may last several decades and constitute 4 fluvial “recovery” intervals. Inter-eruption periods are those where volcanism has little effect on fluvial systems, and typically these are longer in duration than syneruption periods (Smith,

1991). Inter-eruption period fluvial deposits have a more diverse composition and are characterized by more gravel-bedload facies (Smith, 1991). The lateral extent of fluvial deposition during an inter-eruption period is governed by the rate of lateral channel migration

(Smith, 1991).

The relationships of syneruption and inter-eruption deposits can give clues to subsidence rates as well as the eruption frequency of volcanic source areas (Smith, 1991). Major volcanic eruptions will initiate the syneruption aggradation of monolithologic, sand-rich flood deposits with interbedded debrites, which can have a combined thickness of meters to tens of meters

(Smith, 1991). These deposits may also be interbedded with pyroclastic fall or flow deposits

(Smith, 1991). When the hydraulic system begins to recover from syneruption disturbances, channels are incised into the syneruption deposits, and the mobilized sediment is transported to distal regions (Smith, 1991). Subsequently, the fluvial system resumes transport of material from non-volcanic sources, regrades slopes, and returns to transporting gravel bedload (Smith, 1991).

There have been many studies on the recovery of modern fluvial systems where changes in the sediment yield and affected areas can be measured. In these studies, the “rebound effect”

(time interval it can take for river systems to regain their pre-eruption conditions) can be quantified. It has been shown that the rebound effect depends on the volume and distribution of volcanic deposits, local climate, pre-eruption topography, and local hydrology (Manville, 2004).

Complications may include the creation and/or failure of natural dams from the extreme influx of sediment, rilling of surfaces, or superposition of a previous drainage pattern with a newly established one (Manville, 2004). In some cases, climate seasonality (especially the contrast 5 between rainy and dry seasons) can also complicate the rebound effect. For example, the 1964 rainy season following the 1963 eruption of Irazú Volcano in Costa Rica caused more than 90 debris flows (Gran and Montgomery, 2005). The composition and grain size of volcaniclastic inputs can also complicate the rebound effect. For instance, studies on the Pasig-Potrero River after the eruption of Mt. Pinatubo found that sediment inputs were still in decline a decade after the eruption based on sediment composition data. The volcaniclastic sediments of the 1991- eruption were more sand and pumice-rich than the original riverbed deposits. The post-eruption decline in pumice content indicates that hillslope inputs to the system are declining over time

(Gran and Montgomery, 2005). The observations led to a 3-phase fluvial recovery conceptual model (Gran and Montgomery, 2005). Phase one is characterized by fluvial instability and input from numerous lahars. Phase two is the fluvial recovery phase, which is characterized by a change from smooth and unarmored channel beds to coarse-grained, armored channel beds with reduced sediment loads. The final phase is ecological recovery, which is tied to pace of channel recovery, because stable beds help maintain stable species abundance and diversity.

In modern volcaniclastic systems, observations indicate that fluvial reworking of volcaniclastic deposits begins almost immediately after an event and can result in erosional surfaces. Distal volcaniclastic deposits will be interbedded with interdistributary or floodplain facies which may indicate an attempt to re-establish former fluvial conditions similar to that of pre-eruption (Kataoka, 2005). Although some conditions begin to recover soon after the eruption, other conditions may take much longer to recover, for instance the return of organisms.

Ecological recovery may not begin until the channel has somewhat stabilized with decreased sediment transport rates and the development of an armored channel bed (Gran & Montgomery

2005). After the 1991 eruption of Mount Pinatubo, for example, channels with more stable beds 6 and lower suspended sediment supported the return of vegetation, algae, and other aquatic organisms (Gran and Montgomery, 2005). Even after the channel armors and the ecological recovery begins, lateral migration rates can remain high if the sediment transportation rate is still high as well (Gran & Montgomery 2005).

Distinguishing Lahars, Debris Flows, and Hyperconcentrated Flood Flows

The mobilization of large sediment loads can be classified into several types of flows and their subsequent deposits based on their clast size, clay content, or fluid content (Giordano et al.,

2002). Sediment gravity flows are mass transport processes which involve gravity-driven movement of mixtures of solid particles and some type of fluid down a sloped surface (Hampton,

1972). These flows will deposit sediment as the applied shear stress drops below the yield strength of the moving material (Lowe, 1982). Four main end-member types have been identified based on flow rheology and particle support mechanism and include debris flows, grain flows, turbidity currents, and fluidized or liquefied flows (Lowe, 1982). Dilation is often required to trigger a sediment gravity flow, especially around volcanic centers. This can be accomplished by earthquakes or by thermal doming. These flows will move downslope and eventually frictionally freeze when the internal strength of the flow is re-established.

Debris flows

Cohesive debris flows, or mudflows, are sediment gravity flows supported by matrix

(Lowe, 1982). Typically debris flows are a flowing mixture of debris and water having a sediment concentration between 70-90% by mass (Capra and Macías, 2002). Subaerial cohesive debris flows produce deposits (debrites) which are massive, matrix-supported mixtures of gravel- sized clasts within a finer matrix (Giordano et al., 2002). If the debris flow is laminar, the 7 deposits have planar, non-erosive bases and move larger clasts as plug flows. Turbulent debris flows may be able to suspend clasts larger than those which could be suspended by the matrix and buoyancy alone (Lowe, 1982). Turbulent debris flow deposits differ from those of laminar flows by having basal scouring, possible grading, and clast orientation (Lowe, 1982). An individual debris flow may change from laminar to turbulent or vice versa, due to change in flow velocity, change in volume or thickness of material, or change in viscosity.

Cohesive debris flows have been shown to effectively transport sediment in both subaerial and subaqueous environments (Lowe, 1982). Subaqueous cohesive debris flows can occur where flows enter rivers or deep lakes where it may change into a turbidity current where the sediment support mechanism is turbulence rather than matrix strength (Lowe, 1982).

A type of highly concentrated debris flow is called a debris avalanche. A debris avalanche is a rapidly moving incoherent, unsorted, mixture of rock and soil mobilized by gravity (Capra and Macías, 2002). In volcanic areas, they are typically created by the collapse of part of the volcano. A debris avalanche may trigger debris flows or hyperconcentrated flood flows. These deposits also tend to have shattered clasts with a well-developed “jigsaw puzzle” texture where open fractures are filled with matrix (Capra and Macías, 2002). Typically, hummocky topography is produced by a series of debris avalanches (Guarin et al., 2004).

Lahars

Although the term ‘lahar’ has had various interpretations over the years, this paper will adopt the simple definition of “debris flows of volcanic origin” used by Canuti et al. (2002).

Facies associations of lahar deposits may vary depending on topography, type and magnitude of eruption, and proximity to water sources such as a river (Giordano et al., 2002). Lahar deposits 8

are typically matrix-supported, poorly sorted, coarse-grained, and rich in volcanic material.

Lahars may be turbulent or laminar, subaerial or subaqueous.

Mount Merapi, located in Central Java, has become quite well known for its lahars which have occurred continuously over Merapi’s historical record. At least 23 eruptions at Merapi over

the last 500 years have produced lahars which cover an estimated area of about 286 km2

(Lavigne et al., 2000). Many of these lahars were produced during the rainy season of November to April (Lavigne et al., 2000). Mapping these lahars in the 13 drainage basins surrounding

Merapi started in the late-1970s and is still continuing today (Lavinge et al, 2000).

Hyperconcentrated Flood Flows

A hyperconcentrated flood flow is a fluid flow where the high sediment concentration

(40-80%) typically causes hindered settling effects (Jakob and Hingr, 2005). Hyperconcentrated flood flow deposits (inundites) are generally massive or crudely stratified, coarse-grained, and may be normally or inversely graded (Guarin et al., 2004). These deposits are frequently produced by the dilution of one or more debris flows in a stream channel. This can be inferred by the observation of a proximal-to-distal increase in thickness of hyperconcentrated flood flow deposits at the expense of debris-flow deposits (Smith, 1987).

Observed deposits suggest the dilution of debris flows to hyperconcentrated flood flows during flooding on the North Fork Toutle River after the eruption of Mount St. Helens on March

19, 1982 (Pierson and Scott, 1985). Melting of snow and ice produced large volumes of runoff during the eruption, which mixed with volcanic debris and created lahars that flowed 27 km down the North Fork Toutle River (Pierson and Scott, 1985). Although the lahar to hyperconcentrated flow transformation was not observed in the field, the change from typical debrites (very poorly sorted, matrix-supported material containing dark volcanic clasts) to typical 9

inudites (poorly sorted, horizontally stratified, sandier, clast-supported, open framework gravels)

can be observed in the region 27 to 43 km downstream from the crater (Pierson and Scott, 1985).

The hyperconcentrated flood flows are estimated to have traveled at an average of 4.6 m/s

(Pierson and Scott, 1985).

Purpose

Most of the studies described above have been on modern fluvial systems and their

reactions to high volcaniclastic sediment influx after an eruption. These modern studies look at

changes in river morphology, bedload vs. suspended load rates, and recovery of the fluvial

system. Fewer studies have looked for evidence of volcanic influence on fluvial systems in the

rock record. Such studies focus necessarily on the preservation potential of the volcanic

signature. Evidence to indicate volcanic effects include high rates of aggradation or incision,

channel morphology changes, and changes in paleocurrents. The preservation potential of

volcanically-influenced deposits (those from both syneruption and inter-eruption periods) would

ultimately depend on the creation of accommodation space by base-level adjustments, tectonic

uplift or subsidence, or changes in sediment supply.

Because some volcanic centers may be active over long periods of time (1000s of years),

it is also important to look at volcanic influences in the geologic record to determine long-term hazards. Understanding volcanic effects on fluvial systems over geologic time can assess long- term recovery of such fluvial systems as well as potentially show evidence of volcanic events whose primary deposits were not preserved. Determining how well the rock record records volcanic-fluvial interactions may give some insight on what volcanic effects are actually preserved from the past. 10

The goal of this study is to determine if such volcanic influence on fluvial systems is recorded within the McDermott Member of the Animas Formation, near Durango, Colorado. The

McDermott Member is Maastrichtian in age (between 65-70 Ma) and varies in thickness within the Durango area from 78 – 107 m (Kirkham and Navarre, 2003). The McDermott Member is limited to the northern part of the San Juan Basin and is composed of purple, coarse-grained volcaniclastic rocks including tuffaceous sandstones, shales, and conglomerates (Brister and

Chapin, 1994). Particular outcrops contain purple diamicts containing volcanic material and have been interpreted as lahars (Kirkham and Navarre, 2003). The volcanic detritus found within these units has been attributed to volcanism likely associated with the La Plata Mountains found to the northwest during the Laramide Orogeny (Brister and Chapin, 1994).

The McDermott Member is particularly interesting because it contains both lahars and fluvial deposits. This suggests the unit provides insight to the preservation potential of volcanic influence on fluvial systems. This study will use the McDermott Member as a test case to reconstruct volcanic-fluvial interactions and determine not only the preservation potential of this interaction but also the magnitude that volcanic events had on the surrounding area. With the combination of recorded lithology changes, paleocurrent data, and incisional surfaces it may be possible to describe the paleo-landscape and changes to the environment that occurred during the deposition of the McDermott Member.

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GEOLOGICAL BACKGROUND

Location and Tectonic Setting

The study area (Fig. 2) is located southwest of Durango, Colorado in the northern portion of the San Juan Basin, southwest of the San Juan Volcanic Field. The San Juan Basin is located on the east-central area of the Colorado Plateau (Laubach and Tremian, 1994). This structural

basin is Cretaceous – Tertiary in age however it contains rocks ranging from Precambrian to

Cenozoic in age (Laubach and Tremain, 1994). The area of the present San Juan Basin was on

the margin of the Western Interior Seaway during the Cretaceous (Laubach and Tremain, 1994).

During the Sevier Orogeny (mid-late Cretaceous), the western margin of the Western Interior

Basin was subsiding due to thrust sheet loading from the adjacent Cordilleran overthrust belt

(Laubach and Tremain, 1994). During the Laramide Orogeny (late Cretaceous – Eocene), the

Western Interior Basin was segmented into uplifts and intermontane basins such as the San Juan

Basin (Brister and Chapin, 1994). The San Juan sag, located within the San Juan volcanic field,

was formed by asymmetric anticlines bounded by reverse faults (Gries et al., 1997). This feature

was most likely created because Laramide block uplifts were accompanied by westward-directed

compression. The study area also contains the Hogback monocline, which is interpreted as

another major Laramide structure that trends southwest-northeast and dips south - southeast

(Kirkham and Navarre, 2003).

Regional Stratigraphy

Precambrian tectonic activity in the Southern Rocky Mountain region created faults and

shear zones which were remobilized during the Phanerozoic (Brister and Chapin, 1994). During

the late Paleozoic, the Colorado Orogeny produced the ancestral Rocky Mountains (uplifted 12 highlands) which were stripped of their earlier Phanerozoic deposits and then blanketed with

Mesozoic sedimentary rocks (Brister and Chapin, 1994). Cretaceous nearshore clastic units such as the Dakota Sandstone, Mancos Shale, the Mesa Verde Group, Lewis Shale, Pictured Cliffs

Sandstone, Fruitland Formation, and the Kirkland Shale demonstrate at least 3 major transgression – regression episodes along with several less extensive cycles (Fassett, 1974).

Figure 3 shows a representative stratigraphic column of Cretaceous and Tertiary units found in the area near Durango, Colorado.

Late Mesozoic – Tertiary uplifts created a broad feature called the San Juan Dome. The first effects of the San Juan Dome uplift upon sedimentation in the area occurred before volcanism with the thicker deposition of the Fruitland Formation and Kirtland Shale along the northwestern portion of the San Juan Basin where the Hogback monocline is found (Brister and

Chapin, 1994). Paleoflow directions in the Fruitland Formation and Kirtland Shale changed from southwest to the north. The Fruitland Formation contains coal-bearing continental sediments which were deposited as the Western Interior Seaway coastline was retreating to the northeast

(Scott et al., 1994). The first input of volcaniclastic clasts began within the upper Kirtland Shale during the Maastichtian (Brister and Chapin, 1994). Volcanic influences can further be found within the Late Cretaceous McDermott Member and the upper member of the Animas

Formation. The landscape during the Late Cretaceous in this region may be thought of as a low- relief surface covered by a thin alluvial-volcanic blanket (Brister and Chapin, 1994). The Eocene

San Jose Formation is the youngest sedimentary unit in the San Juan Basin and is composed of sandy fluvial sediments (Fassett, 1985). Volcanism increased into the Tertiary, including the eruption of voluminous volcanic rocks to form the San Juan volcanic field during the Oligocene

(Laubach and Tremain, 1994). 13

Animas Formation

Stratigraphy and Age

The Animas Formation is limited to the northern part of the San Juan Basin and grades

laterally into the Nacimiento Formation in New Mexico (Fassett, 1974). Although the name

Animas Formation was first described by Whitman Cross in 1892, John Reeside gave the lower

member a separate name, the McDermott Member, in 1924 (Reeside, 1924).

The McDermott Member is Maastrichtian in age based on palynomorphs (Fig. 4, Table 1) which were collected from beds about 3 m above its base (Kirkham and Navarre, 2003;

Newman, 1987). Reeside (1924) also described fragments of dinosaur bone, turtle bones, and wood within the McDermott Member which support a Late Cretaceous age (Table 2). The Upper

Member of the Animas Formation is considered to be Paleocene in age based on palynomorphs

(Newman, 1987) found in the upper member of the Animas Formation about 30.5 m above the

McDermott Member (Brister and Chapin, 1994). Paleocene plant fossils were also identified by

Knowlton (1924) as well as log imprints (Reeside, 1924).

There is an intraformational unconformity between the McDermott Member and the

Upper Member. Reeside (1924) described the contact between the two members at one outcrop near the Animas River to be an angular discordance. The Upper Member is disconformably overlain by the Blanco Basin Formation and the Conejos Formation in some places (Brister and

Chapin, 1994). In contrast to the red and purple color of the McDermott Member, the Upper

Member consists of drab reddish-brown and green colors indicating reducing conditions

(Kirkham and Navarre, 2003). The McDermott Member has been measured to be up to 107 m in thickness (Sikkink, 1987). The Upper Member has been reported to be about 338 m thick based 14 on a section measured along the Animas River (Reeside, 1924). Both the lower McDermott

Member and the Upper Member display a general fining-upwards sequence.

Lithology and Diagenesis

McDermott Member (Lower Member)

The McDermott Member can be observed along the southeast margin of the San Juan

Dome as distinctive red-to-purple diamicts (Kirkham and Navarre, 2003). Previous workers have interpreted these coarse-grained, poorly sorted, matrix-supported deposits within the McDermott

Member as lahars (Kirkham and Navarre, 2003). The lahars contain andesite porphyry clasts

(Kirkham and Navarre, 2003; Brister and Chapin, 1994) and less common quartz monzodiorite or quartz monzogabbro clasts (Kirkham and Navarre, 2003). The source of the clasts has been interpreted as the La Plata Mountains which lie to the northwest (Brister and Chapin, 1994;

Sikkink, 1987).

Other deposits found in the McDermott Member include coarse-grained volcaniclastic sediments such as purple tuffaceous sandstones, conglomerates, and shales (Brister and Chapin,

1994). Typically, conglomerates grade upwards into grey-white coarse tuffaceous sandstones and purple-grey tuffaceous shales (Reeside, 1924). Sikkink (1987) also identified a paleosol near the top of the McDermott Member based on root structures, burrows, and organic material.

Beds of tan-brown to greenish-brown sedimentary breccia occur in the McDermott

Member as well as the purple deposits. Some workers, Zapp (1949) and Carroll (1997), interpret these deposits as tongues of Upper Animas-like sediments within the basal McDermott Member while others attribute the color differences as the result of chemical alteration (Kirkham and

Navarre, 2003).

15

Upper Member

The Upper Member generally displays a fining-upward sequence with conglomerates and sandstones found in the basal section and sandstones, shales, and sparse coal beds found in the

upper section (Kirkham and Navarre, 2003). Organic matter is common in the Upper Member as

beds of lignitic coal, carbonized leaf matter, and recycled coal from the Fruitland Formation

(Brister and Chapin, 1994). Conglomerate clasts are mostly weathered andesite, quartz, quartzite,

and chert (Reeside, 1924). Sandstones and pebbly sandstones found in this unit contain clasts

from Cretaceous rocks and Precambrian basement (Brister and Chapin, 1994). Unlike the

underlying McDermott Member, the Upper Member lacks lahar deposits (Brister and Chapin,

1994).

Depositional Environment

McDermott Member

The coarse-grained diamict deposits found in the basal McDermott Member have been

interpreted as cohesive debrites, or lahars, because they are coarse-grained, poorly sorted,

matrix-supported deposits containing volcanic clasts (Kirkham and Navarre, 2003). The conglomerates, sandstones, siltstones and shales found in the McDermott Member have been

interpreted as fluvial deposits because of trough cross-bedding sandstone with cut-and-fill

structures and interbedded mudstones are often cut by these channel sandstones (Sikkink, 1987).

The top and base of the member also contains well-rounded quartz-chert-quartzite pebble

conglomerates, whose origins have been attributed to erosion of exposed Dakota Formation

(Brister and Chapin, 1994).

One bed in the basal strata found in the Basin Mountain quadrangle area contains euhedral plagioclase crystals and a lack of sedimentary matrix (Kirkham and Navarre, 2003). 16

Kirkham and Navarre (2003) interpreted this as a sill or lava flow. Similarly, Gries et al. (1997)

located and identified several sills within underlying shale units below the Animas Formation

and interpreted features in a well drilled through the San Juan Volcanic Field as a sill in the

Animas Formation.

Upper Member

The ribbonlike and sheetlike sandstone bodies and the massive shales and mudstones of

the Upper Member have been interpreted as fluvial and floodplain deposits (Brister and Chapin,

1994; Kirkham and Navarre, 2003). Based on isopach maps (Fassett, 1985), and paleoflow data

(Sikkink, 1987), Brister and Chapin (1994) interpreted the Upper Member as an alluvial fan

environment in which streams were flowing to the south or southwest away from the San Juan

Mountains. Conglomerate clasts are andesitic and of volcanic origin which also links the Upper

Member to the San Juan Mountain as the source area (Reeside, 1924; Fassett, 1985).

Paleocurrents and Provenance

McDermott Member

The McDermott Member contains a range of detritus from Precambrian basement,

Mesozoic sedimentary sources, and volcanic sources (Brister and Chapin, 1994). It has been

postulated that the source area for the McDermott Member was the La Plata Mountains (Fassett,

1985; Sikkink, 1987) to the northwest because of paleoflow data as well as isopach map data.

Paleocurrent data from previous studies of the McDermott Member and the Upper Member are

shown in Figure 5. Combinations of this data led to the idea that the McDermott Member was

deposited as a coarse volcaniclastic apron even though only a remnant of this apron is preserved

in the northwest part of the San Juan basin today (Fassett, 1985; Sikkink, 1987).

17

Upper Member

Conglomerate clasts within the Upper Member are primarily volcanic in origin, however, non-volcanic clasts are more common in the top of the Upper Member (Kirkham and Navarre,

2003). There is a change from volcanic-dominated to arkosic-dominated composition moving upwards within the Upper Member (Brister and Chapin, 1994). It has been postulated by Fassett

(1985) that Laramide volcanic centers in the San Juan Mountains were the source area for the volcanic material found in the Upper Member (Kirkham and Navarre, 2003). The volcanic detritus is interpreted to be from unroofing of surrounding highlands based on preliminary field and petrographic observations (Brister, 1990; Brister and Chapin, 1994).

Regional Volcanic History

Cretaceous – early Tertiary volcanism

The only documented surviving occurrence of extrusive volcanic rocks associated with the Laramide orogeny is the Cimarron Ridge Formation on the northwest edge of the San Juan

Mountains (Brister and Chapin, 1994). These volcanic rocks are rhyodacitic in composition

(Brister and Chapin, 1994). It has been concluded that the Cimarron Ridge Formation overlies the Kirtland Shale and older units and because the older units were previously folded (Brister and Chapin, 1994). The La Plata Mountains, which lie 13 – 29 km northwest of Durango, were created by a series of intrusions (laccolith, stocks, sills, and dikes) that caused domal uplift of

Pennsylvanian to Upper Cretaceous sedimentary rocks (Eckel, 1937). Fassett (1985) postulated that the volcanic event which provided material for the McDermott Member was probably short- lived and local because the resulting thin but widespread coarse volcaniclastic apron of the

McDermott Member extends only a short distance around the vent area. The volcanism 18

coincided with regional uplift, resulting in doming and erosion exposing the core of the laccolith

(Brister and Chapin, 1994). Radiometric dating has been done on volcanic rocks exposed near

the San Juan Basin as well as plutonic rocks for these igneous bodies which may represent the

roots of extinct volcanoes. Table 3 provides a list of radiometrically dated material relating to the

San Juan Basin region. In particular, K-Ar and fission tracks dating have been done on diorite-

monzonite porphyry sills, laccoliths, dikes, and stocks of the La Plata Mountains which reveal

ages from 70-67 Ma (Mutschler et al., 1997).

Mid-late Cenozoic volcanism

In the late Cenozoic, basaltic volcanism is widespread but unevenly distributed throughout western Colorado (Budahn et al., 2002). The volcanic fields in Colorado are relatively small in comparison to others associated with the Rio Grand Rift, Colorado Plateau and others in the western US (Budahn et al., 2002). The Oligocene San Juan volcanic field is one of the largest erosional remnants of extensive intermediate – silicic magmatism found in the

North American Cordillera (Budahn et al., 2002; Parat et al., 2005). These late Cenozoic volcanic rocks range from 1000-3000 m thick in the San Juan sag region (Gries et al., 1997).

Volcanism in this area began about 35 Ma with the eruption of andesite lavas and volcanic breccias of the precaldera Conejos Formation (Parat et al., 2005). Activity shifted to large eruptions of dacitic and rhyolitic ash-flow tuffs at about 29 Ma (Parat et al., 2005). At least 17 major ash flow-tuffs erupted during this time from a series of adjacent calderas within the volcanic field (Parat et al., 2005). A significant change in volcanism occurred throughout the region at about 26 Ma, with a transition to eruptions of basalt and rhyolite (Parat et al., 2005).

19

METHODS

Field Work

Field work was conducted in various locations within the Durango, Colorado area from

June 18, 2008 through June 30, 2008. Outcrops of the McDermott Member were found in La

Plata County within the Basin Mountain USGS 7.5’ quadrangle and the Durango East USGS 7.5’ quadrangle. Six stratigraphic sections were compiled ranging from a few meters to almost 64 meters in thickness. The six section locations are numbered and shown in Figure 6.

At each location, GPS coordinates were recorded and photos were taken of the outcrops and of specific features and lithologies within them. Stratigraphic sections were measured taking into attention changes in lithology (composition, grain size, and sedimentary structures) of each bed. The stratigraphic sections were also classified into different lithofacies based on clast size, texture, and sedimentary structures which are shown in Table 4. Consecutive pictures were taken of one particular outcrop, section 6, for compilation of a photomosaic.

The attitude of cross-bedding was measured at six locations. The resulting measurements were then corrected for bedding strike and dip with a stereonet. Ninety-eight (98) rock samples were taken from a number of units within the stratigraphic sections as well as some grab samples from other locations.

Clast counts, which identified the clast size and lithology, from 100 clasts in an individual bed were conducted from lahar deposits at 5 locations within the field area. Figure 7 displays grain-size histograms resulting from these clast counts.

20

Lab Work

Samples collected from the field area were slabbed and subsequent rock chips were

created. From these chips, 43 thin sections were created of which 11 were used for point counts.

Slides chosen for point counts include both clasts and matrix from several lahars as well as

sandstones and the pyroclastic flow deposit. Because the rocks were so feldspar-rich, feldspars

were stained with sodium cobaltinitrite and potassium rhodizonate using the method of

Houghton (1980). This stain turns plagioclase red and potassium feldspar yellow. The thin sections that were not used for point counts were used to better understand lithology and grain size.

Measured stratigraphic sections were redrafted using Adobe Illustrator ®. The resulting stratigraphic sections are shown in Figures 8 through 13. A photo mosaic of section 6 was created using PanaVue Image Assembler ®.

21

RESULTS

Facies Analysis

Facies analysis is based on a unique combination of a lithology, physical and biological

structures, and chemical constituents of a rock (Walker and James, 1992). The purpose of facies

analysis is to use the unique characteristics of a rock or package (sequence) of rocks to determine

the depositional environment and potentially show how it changed over time. Thirteen (13)

different lithofacies were identified and described within the McDermott Member.

Lithofacies 1: Matrix Supported, Massive Diamict or Lahar (Dmmv)

Lithofacies 1 consists of subangular to subrounded, poorly sorted, pebble-to-boulder sized clasts supported by a finer-grained (silt/clay) matrix. The facies is internally massive with clasts dominated by volcanic material. Clast size ranges from <1 cm to almost 1 meter in diameter. These units have generally non-erosive basal contacts however the top of at least one appears to be erosional (Fig. 14a). Individual beds range in thickness from about 2 meters to almost 23 meters, however, it can be difficult to recognize multistory units, so the thickness of individual beds may be less. Beds are also generally tabular in dimension. These beds can range from purple to grey to tan in color.

The poor sorting, absence of stratification, clast size, and non-erosive bases indicate that these beds represent cohesive debris flow deposits or debrites. The internal massive structure and non-erosive bases also indicate that the debris flows which laid down these deposits were subaerial, laminar flows. Such deposits have been previously described by Smith (1987), Lowe

(1982), and Runkle (1990). Because volcanic material is the primary clast component these deposits are interpreted to be lahars.

22

Lithofacies 2: Massive or Crudely Bedded Gravel Conglomerate (Gm)

Lithofacies 2 consists of moderately/well sorted, massive to horizontally bedded,

conglomerate that is typically associated with various sandstones in fining-upwards sequences

(Fig. 14b). The beds are typically 1.5 meters in thickness. The conglomerates form in most

typically tabular beds but also in lenticular forms and some beds which pinched out laterally.

Although no imbrication was evident, these conglomerates often contained intraclasts and are commonly overlain by beds of cross-bedded sandstone. Thin conglomerate also forms the base of small scale cut-and-fill structures over scoured bases.

The conglomerates are interpreted as fluvial deposits because they are stratified, have erosional bases, and are associated with cross-bedded sandstone. The thicker gravels that form tabular beds are interpreted as bedforms (mid-channel gravel bars) using the criteria of Miall

(1978). The thin occurrences of conglomerate beds associated with cut-and-fill structures

represent pools and riffles. The fining-upward sequences represent bar-top deposition of mid-

channel bars (Church and Jones, 1982). The general association of all these is consistent with

coarse-grained, bedload streams with pool and riffle morphology.

Lithofacies 3: Massive Sandstone (Sm)

Lithofacies 3 consists of massive fine- to coarse-grained, moderate to well-sorted

sandstone. These beds can be up to 2 m in thickness, but are generally less than 1 meter thick.

These massive sandstones can be interbedded with cross-bedded or laminated sandstones and can

be part of fining-upward sequences. The sandstones appear massive because sedimentary

structures are not visible.

These massive sandstones are interpreted as parts of mid-channel sand bars because they

are associated with fining-upward sequences as well as cross-bedded and laminated sandstones 23

which Miall (1978) has interpreted to be various bar forms which display lower-flow regime

conditions. Seeing that the massive sandstones are commonly interbedded with cross-bedded

deposits, it is more likely that homogeneity of sand grains prevented any internal stratification from being recognized, rather than rapid deposition, fluid escape, or extensive bioturbation destratifying the unit.

Lithofacies 4: Planar laminated sandstone (Sl)

Lithofacies 4 consists of planar laminated, fine- to coarse-grained, moderate to well- sorted sandstone which displayed various bed thicknesses (Fig. 14c). Laminated sandstone beds are typically ~ 20 cm in thickness and are commonly found interbedded with cross-bedded sandstones and towards the top of fining-upward sequences.

These planar laminated sandstones are like those described by Miall (1978) and have been interpreted as tops of mid-channel sand bars representing upper-flow regime conditions. As it is common to see planar laminated beds overlain by cross-bedded sandstones, this is an indicator of changing flow strength or velocity or the migration of sand bars to a section of the

channel where flow velocity is decreased.

Lithofacies 5: Trough cross-bedded sandstone (St)

Lithofacies 5 consists of trough cross-bedded, fine to coarse-grained, moderate to well-

sorted sandstone beds (Fig. 14d). The angle of cross-bedding ranged from 4° to 20°, and set thickness ranged from 10 – 75 cm. These beds are typically found directly below or above planar

laminated deposits, associated with planar-tabular cross-bedded sandstone, and forming parts of

fining-upward sequences.

Similar trough cross-bedded deposits were interpreted by Miall (1978) to represent 3-D

dunes or that form under lower-flow regime conditions. The association with other cross-bedded 24

deposits as well as their position in fining-upwards deposits indicates these as part of sand bars

(Miall, 1977).

Lithofacies 6: Planar tabular cross-bedded sandstone (Sp)

Lithofacies 6 consist of planar tabular cross-bedded, fine- to coarse-grained, moderate to

well-sorted sandstone beds (Fig. 14e). The thickness of these deposits range from 20 - 30 cm and

they are commonly found above or below laminated sandstones as well as be part of fining

upward sequences, particularly towards the top.

Beds displaying planar-tabular cross-bedding have been interpreted by Miall (1978) to be

2-D sand bars representing lower-flow regime conditions. Galloway (1985) and Walker and

James (1992) describe that planar tabular cross-bedding can occur in finer-grained bar sequences

typically overlying trough cross-bedding representing dune migration.

Lithofacies 7: Sandstone with Intraclasts (Se)

Lithofacies 7 consists of fine- to coarse-grained, moderate to well-sorted sandstone with

mudstone intraclasts. Beds with intraclasts are relatively thin (20 cm) and are found below beds

displaying cross-bedding and above gravel conglomerate beds.

Sandstone beds bearing intraclasts have been interpreted by Miall (1978) to represent

scour fills. Pieces of fine-grained beds, such as floodplain or muddy cutbank deposits which have

undercut or eroded, are eroded and carried to a different location, typically to an adjacent pool,

and deposited as intraclasts. They can be found overlying conglomerates at the channel base, and

overlain by cross-bedded deposits by the continuing migration of dunes down channel. The occurrence of sandstone with intraclasts deposited above conglomerate beds may suggest the intraclasts were brought in during flood conditions.

25

Lithofacies 8: Massive siltstone (Fm)

These fine-grained, silt-sized deposits are well-sorted and generally massive (Fig. 14f).

Thickness ranges from 5 cm to ~ 1 meter. Massive siltstones may occur at the top of fining-

upward sequences or interbedded with massive mudstone deposits. These deposits

predominantly occur at the top of Section 2 where they may display coloration banding (Fig.

15a) throughout.

Similar massive, siltstone deposits have been interpreted by Miall (1978) to represent

overbank or drape deposits. Galloway (1985) also describes such deposits to represent a channel

plug. Sudden abandonment of a channel by a process such as avulsion, may stop the influx of

bed-load sediment and the channel may then be filled by finer-grained suspended load material

(Galloway, 1985) which can be destratified by bioturbation leaving them massive. In the

McDermott Member, an alternate hypothesis is that these are infills of river channel following the creation of natural dams by lahars. In other words, the fine-grained muds and silts may represent deposition within a temporary lake created by such a dam. A study done on the Cowlitz

River in Washington, discovered exposures of 3 debris-flow deposits containing remobilized volcaniclastic sediment from eruptions which that occurred about 2,500 – 3,000 years BP (Chan et al., 2006). Each debris-flow deposit is overlain by 35 cm of horizontally laminated sand and silt that was interpreted to have accumulated in ponded water of the blocked Cowlitz River

(Chan et al., 2006). In the McDermott Member, both Section 2 and 4 include massive and laminated siltstones and mudstones that overlie thick lahar deposits.

Capra (2007) compiled many published instances of volcanic natural dam creation and the length of time it took for them to fail. In particular, during the 1991 eruption of Mt. Pinatubo, a lahar created a natural dam 24 meters tall which failed in less than a month and created a 26

temporary lake containing a volume of 0.075 km3 (Capra, 2007). Another example occurred after

the 1980 eruption of Mt. Hood where a lahar created a temporary dam 10 meters tall which failed

after only 12 minutes (Capra, 2007). Failure of a natural dam can create an outburst or a

secondary debris flow. In such cases, part or all of the fine-grained channel-fill sediment might be removed or at least show incision into other sediments. In some cases, these secondary flows or floods are more voluminous and extensive than the primary volcanic deposits which constructed the dams (Capra, 2007). In the McDermott Member, Section 4 contains multistory lahar deposits with erosional or incisional surfaces exhibited in at least one. This incisional surface may have previously been covered with fine-grained mudstone and siltstone but was not preserved because of an outburst flood. Incision may have occurred from the erosion of not only the fine-grained material, but also the top boundary of the lahar deposit.

The massive siltstones in the McDermott Member often show liesegangue banding (Fig.

15a). This is a secondary feature attributed to hydrothermal alteration. The presence of liesegangue banding indicates later migration of hydrothermal fluids, probably related to the

Oligocene San Juan volcanic field.

Lithofacies 9: Laminated siltstone and mudstone (Fl)

Lithofacies 9 consists of planar laminated, heterolithic, mud to silt-sized deposits also known as mudstone-siltstone rhythmites. They are commonly interbedded with massive siltstone

deposits. Individual rhythmites are < 1 cm thick and packages are up to 25 cm in thickness.

Mudstone-siltstone rhythmites have been interpreted by Miall (1978) to represent

individual overbank or waning flood deposits. These deposits can be found in many

environments but have the highest preservation potential in floodplains, abandoned channels, or

pooled channel deposition behind a natural dam created by the lahar deposits. 27

Lithofacies 10: Massive mudstone (Mm)

This lithofacies consists of massive, well-sorted, mud-sized deposits. Thickness ranges from 5 cm to almost 2 meters. Massive mudstones occur predominantly overlying lithofacies 1 as well as interbedded with fine-grained sandstone especially in Section 4.

As in lithofacies 8 and 9, these massive mudstone deposits can be interpreted as overbank deposits, (Miall, 1978) infilling of abandoned channels, or as fine-grained river channel fill behind natural dams. The last explanation is most likely when the massive mudstones are found overlying lahar deposits as in Section 4. When a natural dam is formed, a lake is commonly created and as infilling occurs, deposition of fine-grained, muddy layers, possibly interbedded with coarser layers may be observed (Capra, 2007).

Lithofacies 11: Massive crystal-ash tuff (Tc)

These deposits are massive, ungraded to normally graded, coarse-grained, moderately- sorted, crystal-rich tuff. Feldspar crystals within these beds are very angular but do not appear to be broken. The tuff also contains volcanic and sedimentary lithic fragments but no original glass.

In several cases, the tuffs transition upward from massive and coarse-grained to planar laminated and fine-grained. One particular sample, 08COS73, found in Section 1, contained a greater abundance of lithics (quartz grains) than the others.

These deposits are interpreted as slightly fluvially reworked tuff. The presence of lithics

(rock fragments as well as some subrounded quartz) shows the effect of fluvial reworking. Cole and DeCelles (1991) identified reworked tuffs whose crystal and vitric components did not vary from primary tuff, however, the reworked tuff contained a greater proportion of non-volcanic lithic fragments as well as evidence of fluvial sedimentary structures. Some tuff beds within the

McDermott Member exhibited planar laminated surfaces most likely indicating fluvial 28

reworking. Crystal-rich tuff was also studied by Runkel (1990) who indicated lithic rock fragments and traction-produced sedimentary structures to be signs of fluvial reworking.

Lithofacies 12: Cross-bedded to wavy-bedded crystal-ash tuff (Tcc)

Lithofacies 12 consists of one example of a very fine-grained, moderate to well-sorted,

crystal-rich tuff. This deposit displays low angle cross-bedding to wavy bedding and is ~ 40 cm

thick (Fig. 15b). This unit is very fine-grained, and contains euhedral crystals of both plagioclase feldspar and pyroxene (Fig. 15c). Estimates of pyroclast fragment sizes indicate 100% ash (<

2mm). Pyroclastic constituents are composed of 100% crystals with an extremely fine-grained opaque matrix, however, it is likely that the deposit once contained glass but was probably devitrified leaving only the crystal components.

Cross-bedded to wavy-bedded crystal tuff are interpreted as surge deposits. Distinctive features that support the surge deposit interpretation include low angle cross-bedding to wavy bedding and better sorting compared to pyroclastic flow deposits. Surge deposits described by

Crowe and Fisher (1973) exhibited low angle cross-bedding as well. Bull and Cas (2000) found a surge deposit with similar euhedral crystals supported in an abundant fine-grained matrix.

Finally, in the McDermott Member, the surge deposit is overlain by a pyroclastic flow deposit. It is common to find pyroclastic deposits, such as pyroclastic flow, surge, and ash-fall deposits, stacked upon each other in the field (Cas and Wright, 1987). More specifically, according to the pyroclast size and constituents, this unit can be classified as a cross-bedded to wavy bedded crystal-ash tuff according to the classification scheme of Schmid (1981).

Lithofacies 13: Lithic-crystal lapilli tuff (TLL)

This lithofacies consists of a coarse-grained, poorly sorted, matrix supported, lithic tuff

(Fig. 15b). The composition of clasts and matrix is homogeneous overall. The matrix is 29

composed of ash-sized fragments rather than crystals. Lava clasts range from lapilli to block- sized (Fig. 15d) and typically become smaller towards the top of the deposit. This fining- upwards of larger clasts is known as coarse-tail grading (Cas and Wright, 1987). The bed is ~ 5 meters thick and overlies the probable surge deposit discussed above. Compared to lahar deposits, the unit is more consolidated. Original glass has most likely been lost so estimates of vitric constituents are interpretive and only applicable to the deposits as viewed today. Although there is a lack of visible glass fragments in this deposit, it is likely that the original glass has been devitrified. An estimate of pyroclast fragment sizes are as follows: 20% blocks (> 64 mm), 50% lapilli (2 – 64 mm), and 30% ash (< 2mm). Estimates for pyroclastic constituents are as follows:

60% crystals, 40% lithics, and 0% vitric.

This lithofacies is interpreted as a type of pyroclastic flow called a block-and-ash deposit

(Cas and Wright, 1987). Block-and-ash pyroclastic flows are topographically controlled and form unsorted deposits with an ash matrix, containing large unvesiculated clasts with a monolithic composition. These particular types of pyroclastic flow deposits are typically unwelded (Cas and Wright, 1987). Homogeneous clast composition and gas segregation pipes are a few features which can help to distinguish block-and ash deposits from debris flows or avalanches (Cas and Wright, 1987). Although clast composition was homogeneous in this deposit, no gas segregation pipes were observed. According to the pyroclast size and component percentages, this unit can be classified as a lithic-crystal lapilli-tuff according to the classification scheme of Schmid (1981).

30

Petrographical Observations

The mineral constituents of the McDermott Member were determined by point counts done on various samples as well as hand samples and other thin section observations. The main objective in studying the petrography of samples was to determine their relationship to each other and source area variation. Eleven samples were chosen for point counting based on degree of preservation of identifiable mineral grains as well as location within stratigraphic sections.

Seven samples of lahar matrix, reworked tuff, and fluvial sandstone were point counted. Three lahar clast samples and two samples from the pyroclastic flow unit were point counted as well. A summary of statistical information for the samples used for point counts is listed in Table 5.

The McDermott Member is dominated by Na-rich plagioclase feldspar. Whether it is sandstone clasts, lahar clasts, lahar matrix, or pyroclastic flow deposits, they all contain a large abundance of abraded and rounded plagioclase grains or subhedral to euhedral plagioclase crystals. This study found both zoned and unzoned, euhedral plagioclase crystals (Fig. 16a), similar to previous studies by Sikkink (1987). Within some volcanic clasts, the feldspar crystals exhibited a pilotaxitic texture, meaning that the crystals have a preferred alignment direction

(Fig. 16b) which may be relict flow banding. Samples of the McDermott Member also contained pyroxenes (Fig. 16c). Clinopyroxenes dominated but some orthopyroxenes were identified as well. Previous studies identified clinopyroxene and augite in volcanic clasts from the McDermott

Member (Sikkink, 1987; Wegert, 2008). Amphibole was identified, however was minor and typically replaced. Several samples such as 08COS75, a sample from a pyroclastic flow deposit, contained minor apatite as well. Most samples also contained smaller opaque minerals throughout, most likely magnetite. 31

Most samples can be classified as moderately altered. The samples in this study display chlorite rims around grains as well as sericitic alteration. Sericite not only replaced crystals (Fig.

16d) but also developed as an intergrowth between crystals (Fig. 16e). Wegert (2008) also identified a low grade phyllitic alteration within his samples. Although all original volcanic glass has been replaced, several samples exhibited ghost glass shards (Fig. 16 g and h) which had been preserved by an opaque mineral. Other samples exhibit orange albite alteration (Fig. 16f) which can coat the grains or fill in between them. This type of alteration has been found to replace potassium feldspar as well as pumice clasts (Gifkins et al., 2005).

The point count data from the lahar matrix, reworked tuff, and sandstone samples was plotted on a QFL diagram in Figure 17a. Most samples plot within the lower left hand corner indicating basement uplift petrofacies based on the classifications of Dickinson et al. (1983).

However, two tan lahar matrix samples (08COS83 and 08COS92) indicate transitional arc petrofacies. Both of these samples had a greater percentage of lithic fragments than other samples. Lithic clasts were typically volcanic but 08COS92 contained a small percentage of sedimentary lithics as well. These petrofacies simply indicate that material from lahar matrix has several sources. The lahars picked up material from previously eroded basement rocks as well as feldspar-rich volcanic material.

The lahar clast and pyroclastic flow deposit point count data was plotted on a QAPF diagram (Fig. 17b). Although previous workers such as Wegert (2008) have identified volcanic clasts within the McDermott Member as trachyandesite based on whole rock composition, the resulting data from point counts plots the clasts as well as pyroclastic flow material in the basalt or andesite zone of the QAPF diagram (Fig. 17b). Based on conventional optical measurements the composition of the plagioclase is sodium-rich (Ab75-86) which supports an andesitic 32

composition rather than basaltic. The sodium content may actually support a more dacitic

composition but there is an obvious lack of potassium feldspars identified by the staining.

Paleogeography

Paleocurrent Data

Cross-bedding found in rock beds provides us with the ability to calculate paleoflow

direction of ancient river systems. Figure 18 contains a rose diagram indicating various

paleocurrent data. Paleocurrent data from section 6 indicates a paleoflow direction to the

southeast and southwest which mimics paleoflow data recorded by Brister and Chapin (1994).

On the other hand, cross-bedding within sections 2 and 4 resulted in a paleoflow direction to the

northeast and northwest. This change in paleoflow direction may indicate a temporary flow

direction change within the river system but more likely indicates something less dramatic such

as cross-bedding which was flanking a bar so river flow moved around it. This change in flow

direction did not appear to occur following the immediate deposition of a lahar deposit so a

paleoflow direction change as a result of damming seems unlikely. It should also be noted that

the paleocurrent data is not statistically significant because of the limited number of measurements taken and/or available.

Lateral Facies Relationships

Sections 2 and 4 demonstrate lateral facies relationships in terms of the thickness of lahar deposits and fluvial facies types and sequences (Fig. 19). Section 4 contains multistory, vertically stacked lahar deposits (facies Dmmv) ~ 39 meters thick capped by ~13 meters of fine-grained mudstones (Mm) and sandstones (Sm). These are interpreted as ponded channel facies due to the decrease in grain size as well as the lack of sedimentary structures. The deposition of these fine- 33

grained deposits overlying lahar deposits (Dmmv) is an indicator of drainage disruption. The

potential creation of a natural dam by lahar deposits (Dmmv) allows for the deposition of fine- grained deposits without interference from normal fluvial flow conditions. These are overlain by fluvial channel deposits consisting of ~4 meters of cross-bedded and laminated sandstones

(facies St and Sl). Section 2 is similar to Section 4. It contains a total of ~ 17 meters of lahar deposits (Dmmv) which are not consecutively stacked but are separated by several meters of cross-bedded and laminated sandstones (Sp and Sl). Overlying the thicker set of lahar deposits are fine-grained siltstone and mudstone (Fm and Fl) and sandstones (Sm and Sl). These are interpreted as overbank or waning flood deposits and again can be deposited as a consequence of river damming. These fine-grained deposits are overlain by repeated sections of fining upward conglomerates (Gm) and cross-bedded and laminated sandstones (Sl, Sp, Sm). These facies are interpreted as various bar forms which represent a braided river environment. Section 2 is then capped by ~6 meters of massive siltstone (Fm) which is interpreted as floodplain deposits because of grain size and relationship to underlying channel deposits. Figure 19 shows a tentative lithocorrelation between Sections 2 and 4. It is quite evident while both sections are similar, that section 4 is dominated by thicker lahar deposits while section 2 is dominated by fluvial channel deposits.

Other differences between the two sections include the occurrence of cross-bedded sandstones (Sp and Sl) in between the lahar deposits in section 2. Furthermore, section 2 contains a greater abundance and thickness of fluvial channel deposits overlying the lahar deposits.

Similarities between the two sections include the presence of fine-grained siltstone and mudstone (Fm, Fl, and Mm) deposited on top of lahar deposits (Dmmv). This is evidence of ponding or lacustrine-type environments created by the damming of the river system. The fine- 34

grained deposits are infilling abandoned river channels and record drainage disruption. Overlying

these fine-grained deposits is a return to more typical fluvial channel deposits (Gm, St, Sl, Sp,

and Sm). This suggests two things: (1) both sections show evidence of drainage disruption and

fluvial recovery and (2) Section 4 shows greater effects of volcanic processes.

Volcanic influence of Sections 1, 2, and 4 may indicate that these represent paleovalleys.

The presence of paleovalleys may not be obvious because they are a large scale feature, and sedimentary rock filling incised bedrock representing a valley probably won’t be visible on an outcrop scale. However, the combination of the sections reveals that some contain volcanic- dominated facies such as primary pyroclastic deposits and lahar deposits (Sections 1, 2, and 4) and some sections contain only fluvially dominated deposits (Sections 3, 5, and 6). The differences in deposits may represent a change from paleovalley fill to valley margin assuming that they were deposited at roughly the same time.

The size of clasts located within the lahar deposits did not vary much over distance.

Figure 7 provides histograms of the longest axis of 100 clasts from several different lahar deposits from the southwest to the northeast of the study area. All 5 measured lahar clast groups were dominated by clasts whose longest axis ranged from 1 to 10.5 cm. Some clast exceeded 50 cm but in general, the clast size was limited to below 20.9 cm.

Proximal to Distal Facies Relationships

In general, proximal-source deposits are dominated by mass-flow processes which produce debris flow deposits and by primary volcanic deposits such as pyroclastic tuff. Distal- source deposits are dominated by fluvial deposits and are less influenced by primary volcanic processes. Sections 4 and 6 provide the greatest difference in proximal to distal-type deposit differences. Section 4 is dominated by multi-story packages of lahar facies (Dmmv) and has a 35 limited thickness of fluvially-influenced facies (Sl, Sp, St). Section 6, on the other hand, is composed of all fluvial facies (Gm, Sm, Sl, Sp, St) and contains no proximal source deposits such as lahar deposits (Dmmv) or pyroclastic facies (TLL, Tcc). Section 5 follows the distal- source type deposits of Section 6 containing fluvial facies (Gm, Sl, Fm). This change from lahar- dominated (Section 4) to fluvial-dominated (Sections 5 and 6) may indicate a spatial difference where Section 4 is closer to the source area than sections 5 and 6.

Alluvial Architecture

The 3-D alluvial architecture of fluvial deposits in the McDermott Member was studied using a photomosaic of the outcrop at Section 6 (Fig. 20). This outcrop contained both channel

(lithofacies Gm, St, Sl, Sp, Sm) and finer-grained deposits (lithofacies Fl). The bar forms here are dominated by sandstone commonly displaying cross-bedding (lithofacies St and Sp) and planar laminations (Sl). This braided river system contains abundant sand bars, some amalgamated to each other especially in the top right region of the outcrop. Large gravel bars are prominent towards the right of the outcrop. Convex shapes outline some bar forms as well as the large channel bottom outlined in the bottom of the outcrop. Bars are established where material is deposited from the sediment typically in areas such as the apex of channel bends, places where the channel widens, and channel junctions (Church and Jones, 1982). Areas between bar forms are labeled as well characterized as bar margin avalanche facies which are angled away from the bar. On top of several bar forms is finer-grained material which is characterized as the bar platform. Some small bar forms are displayed as discrete forms, such as the amalgamated units to the upper right of the photo. Sometimes the small bars will accrete to major forms and anneal into their fabric so that individual bar forms are no longer recognizable (Bluck, 1982). Sand bars in the center of the outcrop may be more difficult to identify individually. The spatial position of 36

bar forms may vary, in that planar laminated surfaces represent of view of bar head on, while

cross-bedded deposits represent a side view. Paleoflow measurements indicate that flow was oblique to the outcrop so it does not represent a perfect cross-sectional view of the river system.

37

DISCUSSION

Depositional Environments

Fluvial

The McDermott Member contains many examples of fluvial lithofacies. Outcrops consist of various types of coarse-grained, cross-bedded sandstones, conglomerates, and fine-grained siltstones which represent mid-channel bars, channel fills, and overbank deposits. This type of fluvial complex can be interpreted as a bed-load channel(s). Bed-load channels have generally straight to sinuous channel patterns, lateral erosion and deposition, and contain coarse-grained channel fill with little suspended-load material (Galloway, 1985). Fluvial architecture can be interpreted as a braided river complex because of the coarse-grained deposits, multiple adjacent gravel and sand bars, and cut-and-fill structures. Although the McDermott Member exhibits many instances of fining upward, these sequences are related to bar top deposition of channel bars rather than by lateral accretion of point bars such as in meandering streams. There is a general lack of lateral accretion surfaces as well, which would indicate a meandering fluvial system rather than braided. A braided fluvial style, found here, is also most typical for tectonically active or volcanic environments (Walker and James, 1992).

Lacustrine

Several places within the McDermott Member, fine-grained siltstones, mudstones, and sandstones (lithofacies Fm, Fl, Mm, Sm) are found immediately overlying lahar deposits

(Dmmv). The lahar deposits can act as a natural dam which interrupts drainage disruption within the fluvial system and may create a temporary lacustrine environment. This allows for the settling out of fine-grained material. Capra (2007) explains that creation of temporary natural 38

dams or impoundments is a commonplace occurrence surrounding volcanic centers. A variety of

eruption-induced deposits such as debris avalanches, lahars, and pyroclastic flows have been documented both the modern and ancient record to have created natural dams which, in many

cases, created lakes. Although some of these natural dams withstood over time, some failed after

minutes, days, or years. Capra (2007) explains that natural dam stability depends on (1) the dam

volume, (2) the material type, and (3) the inflow rate.

Furthermore, many natural dams, such as those created by debris flows, may lead to the

generation of secondary debris flows or of outburst floods. These may be recorded by the

dilution of a debris flow into a hyperconcentrated flood-flow downstream of the dam (Capra,

2007) or by incisional surfaces where the fine-grained, infilled river channel sediment has been

eroded away. Within the McDermott Member, the fine-grained, temporary lacustrine deposits are

preserved in several locations such as Sections 2 and 4. There are potential areas, however,

where the evidence of this fine-grained material has been eroded away. Within Section 4, there is

an incisional surface which separates two lahar deposits (Dmmv) at the 6 meter interval (Fig.

19). Here, there is erosion into the underlying lahar deposit which may be the result of fluvial

erosion or by an outburst flood which has not only eroded the fine-grained sediments, but further

eroded into the lahar deposit.

Volcaniclastic

Proximal to distal volcaniclastic deposits can be classified as such based on the

predominance of primary deposits and debris flow deposits to that of fluvial deposits. For

example, deposits considered proximal to their volcanic source area can contain thicker deposits

of primary volcanic deposits (lava, ash-fall, surge and flow tuff) and thicker debris flow deposits

as well (Runkle, 1990). Deposits of intermediate thickness consist of hyperconcentrated flood- 39

flow deposits with an increase in fluvial interaction (Runkle, 1990). Distal source deposits are

dominated by fluvial sedimentation rather than mass-flow (Runkle, 1990). Debris flow deposits

and primary deposits tend to thin away from the source area so that they form a wedge-like shaped deposit (Smith, 1991). Proximal – distal relationships within the McDermott Member

may not encompass all sections, but a proximal – distal facies change is seen in sections in the

northeast of the study area. Previous studies done on the McDermott Member indicate a source area to the northwest based on paleoflow data, isopach maps, etc. This would indicate that the collection of stratigraphic sections from this study would follow a lateral transect where all sections are generally the same distance from the source area. Although Sections 4, 5, and 6 are close together, there is a change from lahar dominated deposits (Dmmv) within Section 4, to fluvially dominated deposits (Gm, Sm, Sp, Sl, St) in Sections 5 and 6. This may indicate that

Section 4 represents proximal facies and Sections 5 and 6 represent distal facies in relationship to the volcanic source area. Other possibilities include these sections representing separate drainage, a difference in time period thus reflecting differences in volcanic influence, or preservation disparity.

Volcanic Impacts on Fluvial Environment

Drainage Disruption

Drainage disruption of a fluvial system in response to volcanic-induced processes is a common occurrence. Within the McDermott Member, normal fluvial conditions are seen before lahar deposition with the deposition of various cross-bedded and laminated sandstones (Gm, St,

Sl, Sp). Drainage disruption is then seen several times following the deposition of thick lahar deposits (Dmmv) which overly the fluvial deposits. Rather than an immediate change back to 40

normal fluvial sediments, there are significant deposits of fine-grained siltstone and mudstones

(Fm, Fl, Mm) which may be the result of natural dams caused by the lahar deposition. The dams

may create lacustrine environments or temporary ponding of the fluvial channels which allow for

the deposition of the fine-grained sediments. Further disruption of drainage can occur with the failure of the natural dam which may occur as an outburst flood or secondary debris flow.

Deposits produced during syneruption periods by processes like lahars can be considered geologically instantaneous (Smith, 1991). Such processes can occur numerous times over short periods of time such as years or decades and their occurrences are visible in our historical record.

Furthermore, the fine-grained sediment which has been found to infill temporary lacustrine environments created by natural dams can also have a rapid rate of deposition. One particular ancient example, the Nevado de Colima (Mexico), contained thick lacustrine deposits which was

indicative of a lake which lasted for some time before it overflowed although inflow rates

indicated infilling occurred over only a few days (Capra, 2007). The 1.8 ka eruption of Taupo in

New Zealand also resulted in ignimbrite-created natural dams which allowed for the rapid

accumulation of locally thick lacustrine deposits which formed as a result of the drainage

disruption (Manville, 2001). Fluvial recovery models created today such as that of Gran &

Montgomery (2005), can track river recovery over years to decades. Overall, it appears that the

processes described above dealing with lahar deposition, the creation of natural dams and the

subsequent creation of temporary lakes, and fluvial recovery can occur on a geologically short

time scale. This potentially means that certain sections within the McDermott Member represent

“snapshots” of geologic time which may represent years or decades rather than thousands or

millions of years.

41

Excess Sediment Loads

Following an eruption, sediment input can increase dramatically as ash-fall loaded

terraces and upland areas are eroded and from the destruction of stabilizing vegetation (Gran &

Montgomery 2005). A river’s morphology may change to compensate for the high sediment loads. Initially, finer-grained materials are preferentially mobilized but over time, the bed coarsens and pebble clusters may form (Gran & Montgomery 2005). Over time, the pebble clusters amalgamate to form armored gravel bars which aids in further decline of sediment transport rates (Gran & Montgomery 2005). In the McDermott Member, individual pebble clusters are not identifiable, but following the introduction of lahar material does eventually result in the coarsening of the fluvial channel. Following the deposition of fine-grained material in response to natural dam creation by lahar deposits (Dmmv), fluvial channel deposits appear to regain their coarser-grained fabric (Fig. 9). Displayed in the upper half of Section 2, conglomerate and sandstone facies (Gm, Sp, Sm, St, Sl) are deposited and coarsen upsection.

This is an indicator of erosion and removal of the fine-grained material which also may point to decreasing sediment input. Particular rivers may aggrade vigorously in response to sediment overload followed by downcutting into the underlying sediments or bedrock. Although the

McDermott Member does not exhibit major downcut surfaces, the aggradation of sediments followed by the return to coarser channel sands and removal of fines displays the paleofluvial system’s response to the excess sediment input. Furthermore, the fluvial morphology in the

McDermott Member does not change from meandering streams to more braided streams following increased sediment influx, but does appear to return to coarse-grained, braided stream morphology which is displayed before the excess sediment input. 42

As previously mentioned, the McDermott Member shows evidence of paleovalleys as well as fluvially dominated areas, potentially marginal to the paleovalleys. Distinguishing

between certain sections can provide a view of the paleogeography of the area. Section 1 (Fig. 8) contains primary volcanic deposits whose original transport is topographically controlled.

Although the outcrop itself was found on a present-day ridge, the surge and flow deposits would have been deposited in valleys and/or depressions and therefore, probably represent a paleovalley. Lahars are similar in that they travel furthest down valleys where their deposits are found at their thickest. For instance, Mount Merapi contains multiple river channels flanking its slopes which feed lahar deposits downvalley, which in turn is why all the lahar sensors on the volcano are located on the rivers such as K. Boyong and K. Bebeng (Lavinge et al, 2000).

Fluvially dominated sections (Sections 3, 5, and 6) may represent areas between valleys or along

their margins. The fact that lahar deposits (Dmmv) in particular, are much thicker in section 4

(Fig. 11) may indicate a deeper valley. If sections 2 and 4 represent different valleys, of different

depths, then the thickness of such deposits would be different. On the other hand, section 2 may

represent a wider valley such that all lahar deposits were still distributed as laterally extensive

sheets but are thinner because the same amount of material is covering a wider area. If this was

the case, then it may have been easier for the fluvial system to recover from high sedimentation

influx conditions. This may account for the thicker succession of channel fluvial deposits in

section 2.

Another potential reason for the differences in lahar deposit thicknesses in Sections 2 and

4 may be because they represent two separate volcanic events. A larger or stronger volcanic

eruption may have dislodged or discharged more material which traveled in the direction of 43

Section 4 whereas, processes which caused the deposition of lahar material at Section 2 may have been lower in intensity or duration.

Paleocurrents and Provenance

Section 6 provides paleoflow data which indicates a flow to the southeast and southwest.

This data follows the paleocurrent trends which have been identified by Brister and Chapin

(1994). As this potentially represents a source area to the northwest, it may be inferred that deposition of thick lahar deposits at these locations did not appear to interfere with paleoflow direction. The limited number of paleoflow measurements taken and/or available in these locations makes them not statistically significant. Several paleocurrent measurements resulted in a northeast/northwest direction but variation in paleoflow direction could indicate something such as the movement of water flow as it traveled around various bar forms which may result in a 180º variation in paleoflow direction (Miall, 2000). Small-scale cross-bedding may also be more influenced by bar form or turbulent eddies which skew the true overall river flow trend so that they may not match up exactly (Miall, 2000).

Clasts in the lahars show little variation with regards to composition and clast size.

Variation in the clasts within the lahars only occurs in the tan lahar from Section 4 (Fig. 11). The clasts in this unit appeared more abundant and more angular compared to other lahar deposits in

Section 4 and 2. Provenance variation may be attributed to this apparent abundance of juvenile material. Lahar matrix color varies as well in between Sections 2 and 4. The lahar deposits are distinguished by color which ranged from purple, grey, to tan. The color of their matrix may represent alternate source areas that sampled materials of different composition. The color difference may also be related to diagenesis. For many pre-Quaternary volcaniclastic sequences rock color is largely determined by diagenetic alteration of volcanic grains (Smith, 1991). 44

Although source area may not have changed overall for the McDermott Member, clasts within

the tan lahar deposit of Section 4 may have been derived from less reworked La Plata volcanics

compared to those of other lahar deposits. As the clast composition does not vary for other lahar

deposits, any provenance difference would most likely be minor and may be limited to a

difference in valley or area which the lahar initially travel.

Fluvial Recovery Trends

Present-day monitoring of fluvial systems before, during, and after an eruption, can provide quantitative data which can exhibit how a fluvial system recovers from excess sediment influx. Gran & Montgomery (2005) have explained fluvial recovery to include declining sediment input over time, coarsening of the river bed and formation of features such as pebble clusters and armor patches over time, and ecological revival. Within the McDermott Member, a change in lithofacies and grain size is evident over time to indicate fluvial recovery. A cycle of

(1) pre-eruption fluvial (Gm, Sl, St, Sp, Sm), (2) lahar deposition (Dmmv), (3) deposition of

fine-grained material (Fm, Fl, Mm), and (4) return to normal fluvial conditions (Gm, Sl, St, Sp,

Sm) is displayed several times within the McDermott Member in particular, in Sections 2 and 4.

A representation of this fluvial recovery process over time is displayed in Figure 21. Although

we cannot measure sediment input rates or see ecological recovery within the McDermott

Member, a cycle of facies changes allows for the interpretation of fluvial recovery.

Signature of Volcanic Events in Geologic Record

Pyroclastic Surges and Flows

Pyroclastic density currents, like surges and flows, are topographically controlled which can provide clues on the paleogeography of the study area. While pyroclastic flows are confined 45 to pre-existing valleys, higher energy pyroclastic surges have the ability to overtop hills and therefore their deposits may be more extensive than pyroclastic flow deposits. Preservation of such deposits may depend on their lateral extent and how quickly they themselves are buried.

Fluvial influence may rework primary pyroclastic deposits such that the original structures are lost and we are left with a sand-rich facies rather than the actual pyroclastic deposit (Smith,

1991). However, primary pyroclastic deposits may be found interbedded with syneruption deposits like debris flow deposits which are non-erosive and can potentially preserve these primary deposits. Primary pyroclastic deposits are only found in Section 1 of the study area. The primary volcanic deposits are not only thin but not laterally extensive. The two pyroclastic beds are interpreted to be a surge and a block-and-ash deposit. In general small pyroclastic surges or flows have a lateral extent that is generally limited to < 20 km from the vent (Cas and Wright,

1987). It has been documented, however, that large pyroclastic flows (forming ignimbrites) from calderas such as Valles (New Mexico), Aira and Aso (Japan) have runout extents up to 50 km away from the source (Sheridan, 1979). Although it may not be possible to distinguish a source vent for these particular deposits, locating other outcrops in the surrounding area may provide clues about the location of their source and extent.

Block-and-ash deposits, like the one found within the McDermott Member, are created by lava-dome collapse. Dome collapse can operate on a steep-sided andesite volcano due to either gravitational forces or from an explosively directed blast (Cas and Wright, 1987). When a dome collapse occurs, fragmented lava is generated and can travel downslope as a pyroclastic flow (Cas and Wright, 1987). Pyroclastic flows have high particle concentrations, are dominated by laminar flow mechanisms, and produce deposits which are poorly sorted (Cas and Wright,

1987). Pyroclastic flows which produce block-and-ash deposits are generally small in volume, 46

producing < 1 km3 of material during a single eruption (Cas and Wright, 1987). Pyroclastic

surges are commonly associated with pyroclastic flows and can be generated by a directional

blast, out of the head of a moving pyroclastic flow, or from small gravitational collapse at the margins of an eruption column (Cas and Wright, 1987). Surges are turbulent, more diluted than flows, and produce deposited which are well sorted along individual laminae and can be cross- bedded (Cas and Wright, 1987).

Lahars

Lahar deposits, like those found within the McDermott Member, are syneruption deposits which have a high preservation potential because they are laterally extensive and relatively thick

(Smith, 1991). While the chance for these deposits to be reworked and incised by fluvial processes may be high, lahar deposits can be well preserved into the rock record as they are within the McDermott Member. Lahar deposits also have the ability to record volcanic eruption events within the material that they carry. Lahars contain abundant volcanic material which can include lava clasts, pyroclastic rock fragments, and previously deposited volcanic material which the lahar may pick up while in transit. As shown in Figure 22, lahars have a tendency to travel down river valleys and may be deposited radially from the volcanic center. The extent and travel path may vary depending on amount of material mobilized, initial direction of dislodged material, and original location of loose material on or near the volcano.

Tuff

Tuff deposits created by the fallout of primary pyroclastic material from a volcanic eruption can be found widespread from the actual vent. Thickness of fallout material generally decreases exponentially from the source area (Walker and James, 1992). Ash-fall deposits will mantle the topography unlike pyroclastic flow and surge deposits which may be confined to 47

valleys (Cas and Wright, 1987). Although ash-fall deposits may be laterally extensive, they are

not generally thick like lahar deposits. The McDermott Member contains very minor primary tuff

deposits (TC, TLL, TCC). While the primary pyroclastic deposits (surge and flow) are preserved in

one small area, other tuff deposits within the McDermott Member appear to be slightly

reworked. Fluvial influence is evident within reworked tuff of the McDermott Member by

general lack of grading as well as sedimentary structures found within the tuff beds indicating

fluvial influence. While abundant primary tuff deposits may not be preserved, their signatures

are recorded in the McDermott Member within sediments like tuffaceous sandstones which

contain abundant tuff crystal material.

Volcaniclastic Rocks

Many rocks within the McDermott Member contain components of weathered volcanic

rocks or volcanic fragments and include tuffaceous sandstones and siltstones as well as the lahar

deposits and primary pyroclastic deposits. These tuffaceous rocks are included in the previously

mentioned fluvial facies (Sl, Sm, Sp, St, Fm) and can include weathered volcanic rock fragments

and crystals. The tuffaceous rocks may acquire their volcanic components from reworked primary ash-fall deposits or from the erosion of volcanic clasts like those found within the lahar

deposits (Dmmv). Tuffaceous sediments provide evidence of volcanic processes for which all primary deposits may be lost. Although the McDermott Member only contains minor tuff deposits, it does contain an abundance of tuffaceous material which is another way that the rock

record can preserve evidence of volcanic processes.

Diagenesis and Metamorphism

Volcanic sediments and primary pyroclastic deposits from the McDermott Member are

altered and contained various alteration minerals such as sericite and chlorite. Sericite, a very 48 fine-grained white mica, typically occurs as a replacement or alteration product of feldspar

(Raymond, 2002). The lack of potassium feldspar in most samples may due to their replacement to sericite. Chlorite also commonly replaces feldspar especially in the finer-grained matrix of these samples. The level of alteration can be described as moderate to strong because pseudomorphs are common especially for mafic minerals like clinopyroxenes. Sericite and chlorite are found replacing feldspar crystals and matrix, and any glass has been either completely destroyed or replaced by opaque minerals (Gifkins et al., 2005). Hydrothermal alteration is also evident within the McDermott Member by the occurrence of liesegangue banding. This banding is caused by the mineralization of mobilized minerals through dissolution and is visible evidence of alteration related to fluid flow (Fu et al., 1994).

49

SUMMARY & CONCLUSIONS

Summary

The McDermott Member is a volcanically influenced fluvial unit. It shows evidence of volcanic influence on fluvial sedimentation by ways of drainage disruption, creation of natural dams and probable avulsions. The McDermott Member represents a variety of fluvial channel and bar forms creating a braided river environment (Gm, Sm, Sl, Sp, St) with paleocurrent flow direction towards the southwest. This braided river morphology was temporarily interrupted several times by the deposition of thick lahar deposits (Dmmv). Lahar deposits create natural dams which create ponding of the fluvial system or a temporary lacustrine environment where fine-grained siltstone and mudstone is deposited (Fm, Fl, Mm). These fine-grained deposits are then overlain by coarser-grained channel sandstones (Sm, Sl, Sp, St) which represent fluvial recovery and a change back to normal fluvial conditions. The McDermott Member also contains primary volcanic deposits like pyroclastic flow and surge deposits but is dominated by fluvially reworked materials like lahar deposits and tuffaceous fluvial sediments. These types of deposits can record volcanic events even if the primary deposits are not preserved. Variation of sections laterally across the McDermott Member containing lahar deposits (Dmmv) and/or primary volcanic deposits (TLL, Tcc) represent paleovalley environments which ranged in depth or width.

Alternate sections are dominated by fluvial deposits (Gm, Sm, Sl, Sp, St) and represent areas adjacent or marginal to paleovalleys. The McDermott Member provides us with an excellent example of volcanic-fluvial interactions because it contains both preserved syneuption deposits

(lahars) and inter-eruption deposits (fluvial). The McDermott Member also offers a look at these interaction cycles over time by displaying multiple recovery instances in the rock record.

Furthermore, the McDermott Member potentially provides a view of a small window of time in 50 the Late Cretaceous because the length of time for lahar deposits to be laid down, the creation of temporary lakes, and fluvial recovery have been shown to occur on years to decades time scale rather than thousands to millions of years.

Conclusions

The McDermott Member provides the ability to observe volcanic events within the rock record even with a limited number of primary deposits. Lahar deposits, mildly reworked tuff, and tuffaceous sandstones are major components of the McDermott Member and represent volcanic signatures in the rock record. Fluvial reworking of primary volcanic deposits may lose the initial signature, but deposits such as tuffaceous sandstones and lahar deposits contain volcanic lithic fragments and/or crystals that somewhat preserve the volcanic event.

Preservation potential of syneruption and intereruption deposits varies depending on subsidence rates, base level adjustments, and eruption frequency (Smith, 1991). The lahar deposits, classified as syneruption deposits, have a high preservation potential because they are relatively thick and laterally extensive. Preservation potential of inter-eruption deposits like normal fluvial sediments is much lower if the deposition occurs on a tectonically stable area or one with low subsidence rate (Smith, 1991). In the case of the deposits in the McDermott

Member, both syneruption and inter-eruption deposits are preserved so we have the ability to see the effects of each type on each other. As in all depositional environments, several factors can contribute to the likelihood of deposit preservation and geometry. Eruption type and frequency can determine the relief of source area as well as sedimentation input rates. Sediment dispersal can be greatly controlled by tectonics and subsidence. Subsidence within the source area may favor particular storage areas like valleys, and starve adjacent areas of a volcaniclastic apron 51

(Smith, 1991). Subsidence and tectonic time scales, however, deal with a much longer time scale

(100,000’s to millions of years) than volcanic deposits (100’s to 10,000’s of years).

The McDermott Member also provides the ability to view fluvial recovery cycles within

the rock record. Observing these events and cycles in the rock record may provide an analog to

consider for modern processes. Rather than only considering the hazards of a volcanic eruption

on a short time scale, with the addition of rock record information, it can allow us to be more

prepared for extensive effects such as multiple eruptive events or the remobilization of sediments

deposited by an eruption. Although the rock record may not provide exact measurements such as

sedimentation rate or temporal length of fluvial recovery which can be measured in the modern

record, it provides an extended history of the volcanic center and potential extent of such effects.

It is important to not only study volcaniclastic deposits on an outcrop scale, but also more

specifically on a meter-to-meter scale. A broad classification of these deposits in only a general

way can potentially lose information about the history of the unit’s deposition. For instance,

Sikkink (1987) classified outcrops of the McDermott Member into three different facies: (1)

volcaniclastic conglomerate, (2) trough cross-bedded, volcaniclastics, and (3) siltstone and claystone. This general description of the McDermott Member may be useful for depositional environment interpretation and a sense of the overall fining upwards appearance of the member, but without a more detailed look, information is lost such as the repeated normal fluvial flow interruption, the creation of natural dams by the lahars, and evidence of fluvial recovery.

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Figure 1 – Example of a classification of proximal and distal volcaniclastic environments around a volcanic center based on types of deposits and processes. Adapted from Vessell and Davies (1981).

53

Figure 2 – Map of Southern Colorado displaying several structural, volcanic, and sedimentary provinces relevant to the study area. Adapted from Gries et al. (1997).

54

Figure 3 – Upper Cretaceous and Tertiary rock units found around the Durango, Colorado area.

55

Figure 4 – Biostratigraphy of Upper Campanian, lower Maastrichian, and lower Paleocene formations in southwestern Colorado. Adapted from Newman (1987).

56

A.

B.

Figure 5 - Previous paleocurrent data from the Animas Formation; (a) Brister and Chapin (1994) created paleocurrent rose diagrams for the McDermott Member (left) and the Upper Member taken from the San Juan Sag area, and (b) rose diagrams complied by Sikkink (1987) from paleocurrent data from the Animas Formation located at the north/northwest margin of the San Juan basin.

57

in UTM. Adapted from Kirkham stratigraphic sections measured in th e tions 2, 4, and 6. Coordinates are at Sec taken measurements nt ock units found south of Durango, CO. Locations

and Navarre (2003). Figure 6 – Geologic map of bedr field are labeled 1 – 6. Paleocurre 58

A B 100 100 90 90 80 80 70 70 60 60 50 50 40 40 Clast number

Clast number 30 30 20 20 10 10 0 0 0 1020304050 0 1020304050 Longest axis range (cm) Longest axis range (cm)

C D

100 100 90 90 80 80 70 70 60 60

50 50 40 40

30 number Clast

Clast number Clast 30

20 20 10 10 0 0 0 1020304050 0 1020304050 Longest axis range (cm) Longest axis range (cm)

E

100 90 80 70 60 50 40

Clast number 30 20 10 0 0 1020304050 Longest axis range (cm)

Figure 7 – Clast count histograms from various lahar deposits from different areas around Durango; (a) counts taken from purple lahar outcrop furthest to southwest near section 1, (b) counts taken from purple lahar outcrop at section 2, (c - e) counts taken at section 4 from (c) grey lahar, (d) purple lahar, and (e) tan lahar.

59 of a area. This section consists

hwest corner of the study

ction 1 located at the sout two primary pyroclastic units.

Figure 8 – Stratigraphic column representing Se Figure 8 – Stratigraphic conglomerate, reworked tuff, and

60

ts of thick lahar deposits (ex. 11-26 m) dded and fining upwards sequences).

Figure 9 – Stratigraphic column represen tative of Section 2. The section consis as well fluvial deposits (cross-be

61

Figure 10 – Stratigraphic column representing Section 3 located along the bank of the present day Animas River. This section is composed of generally purple fine-grained sandstone.

62

of the study area. This section and mudstone. The upper most beds are and mudstone. The cated in the northeast corner deposits. (ex. 4-43 m) followed by fine-grained silt representative of Section 4 lo luvial f developed - ell w y b

Figure 11 – Stratigraphic column characterized consists of multi-story lahar deposits 63

Figure 12 – Stratigraphic column representative of Section 5.

64

Figure 13 – Stratigraphic column representative of Section 6 in the northeast corner of the study area. Section 6 is characterized by cross-bedded and laminated fluvial deposits. 65

A B

C D

E F

Fl

Fm Fl Fm Fl Fm

Figure 14 – Photographs of various lithofacies found in the McDermott Member; Yellow lines highlight lithofacies separations or cross-bedding; (a) erosional contact between the upper purple lahar deposit and overlying grey lahar deposit (Dmm), (b) fining upwards sequence in a gravelly conglomerate (Gm), (c) planar laminated sandstone (Sl), (d) trough cross-bedded sandstone (St), (e) planar-tabular cross-bedded sandstone (Sp), and (f) interbedded massive and laminated siltstone (Fm/Fl).

file:///C|/Users/Shea/Documents/Figure14_ro.htm[6/22/2009 7:13:01 AM] 66

TLL

TCC

A B

C D

Figure 15 – (a) Coloration banding within siltstone/mudstone facies, (b) outcrop photo of pyroclastic flow deposit overlying a pyroclastic surge deposit, (c) euhedral plagioclase crystal (top) and euhedral pyroxene crystal (bottom) within surge deposit, (d) clast-rich fabric of the block-and-ash deposit.

67

A B

CPX S

C D

E F

GS

GS

G H

Figure 16 – Photographs of thin sections from the McDermott Member; (a) zoned and unzoned plagioclase, (b) pilotaxitic texture within feldspar crystals, (c) a clinopyroxene crystal (CPX) in the center of the photograph, (d) feathery yellow sericite (S) alteration of two crystals, (e) sericite intergrowth between three plagioclase crystals, (f) orange albite alteration found within a clast in a lahar deposit, (g) glass ghost shard (GS) in the center of a volcanic clast found within a lahar deposit, (h) magnified view of the glass ghost shard from (g) . 68

Note: All 5 samples b. oclastic flow unit (08COS75 and

08COS89, 08COS92), grey lahar matrix (08COS37), OS73); (b) QAPF plot of lahar clasts and pyroclastic deposits. Purple ey lahar clast (08COS38), and pyr a. 08COS76). ott Member; (a) QFL plot of various lahar matrix samples, reworked samples, lahar matrix QFL plot of various (a) Member; ott Figure 17 – Petrography of the McDerm tuff, and sandstone. Tan lahar matrix (08COS83, sandstone (08COS45), reworked tuff (08C lahar clasts (08COS19, 08COS46), gr

69

Figure 18 - Rose diagram displaying paleocurrent data from the McDermott Member; Vector mean is 220, vector magnitude is 0.47 with a circular standard deviation of 58.96, circular variance is 3476.3, and test of significance is 0.266. Paleocurrent data is not statistically significant because of the small number of measurements.

70

Figure 19 – Vertical Stratigraphic sections of Section 2 (left) and Section 4 (right). Beds outlined in red are fluvial deposits, deposits in yellow are lahar deposits, and beds outlined in green are fine-grained deposits overlying lahar deposits which represent drainage disruption. Thick black tie lines connect corresponding lahar deposits and reworked tuffs. Section 2 contains significantly thinner lahar deposits compared to Section 4. Section 2 also is dominated by well developed fluvial deposits in the upper half section. Both sections have fine-grained siltstones and sandstones overlying the thicker lahar deposits.

71

ify most that the flow was that the n, BP = bar platform, red asterisks ident the river features.

a perfect cross-section of Section 6; SB = sand bar, BM bar margi cture as a result of photo stitching. Paleocurrent data indicates cture as a result of photo stitching. Paleocurrent photo does not represent Figure 20 - Photomosaic and interpretation of areas of doubling features within the pi so this to the outcrop oblique likely 72

t1

t2

t3

t4

Figure 21 – Representation of fluvial recovery process shown in the McDermott Member; The sketches t1-t4 display: normal fluvial sedimentation near a volcanic center (orange) (t1), the deposition of thicker lahar deposits (purple) (t2), the creation of a natural dam by the lahar deposits and the subsequent deposition of fine-grained lacustrine deposits (red) upstream of the dam (t3), and finally the return to normal fluvial sedimentation (t4). Section 2 may be a representation of these series of deposits within the McDermott Member.

73

Figure 22 - 3-dimensional representation of lahar deposition from a volcanic center. The lahar deposits (purple) can be found in valleys and their extent and travel paths may vary distally from the volcanic source.

74

Table 1 – Alphabetic list of palynomorph taxons used for dating of strata in Colorado (Newman, 1987).

Aquilapollenites delicatus Syncolporites Aquilapollenites Thomsonipollis magnificus Balmeisporites Tilia danei Cranwellia Tilia wodehousei Gunnera microreticulata Triatriopollenites Ilexpollenites compactus Triatriopollenites trina: Alnus Trina Interpollis Tricolpites interangulus Kurtzipites Tricolporopollenites confossus Kuylisporites scatatus Triporopollenites rugatus Liliacidites complextus Triporopollenites scabroporus Momipites coryloides Trudopollis meekeri Momipites tenuipolis Ulmoideipites Myrtaceoipollenites peritus Umbosporites callosus Polyporopollenites Vacuopollis (Conclavipollis) wolfcreekensis Proteacidites Wodehouseia spinata Pseudoplicapollis (Sporopollis)

75

Table 2 – List of various plant and vertebrate fossils found within the McDermott Member used for dating of unit (Reeside, 1924)

Flora Onoclea neomexicana Quercus praeangustiloba Asplenium neomexicanum Ficus denveriana? Salpichlaena sp. Ficus eucalyptifolia Fern, genus? Ficus leei Sequoia acuminata? Lesquereux Ficus planicostata Cyperacites sp. Ficus sp. Canna? magnifolia Knowlton Cinnamomum linifolium Sabalites sp.? Grewiopsis sp. Palm, genus? Cissus coloradensis? Pistia corrugata Leguminosites? neomexicana Myrica? neomexicana Rhamnus cf. R. cleburni Pterospermites sp.

Dinosauria Kritosaurus navajovius Brown Monoclonius sp. Armoured dinosaur (Scelidosauridae) Carnivorous dinosaur suggesting Dryptosaurus and Dynamosaurus Chelonia Baena nodosa Gilmore Pisces Lepisosteus sp.

76 the northern San Juan Basin. r rock units within the McDermott Member Member the McDermott r rock units within nd plutonic rocks in proximity to – Radiometric ages of volcanic a Table 3 Table 4 – Lithofacies classification table fo

77

Table 5 – Statistical analysis of Petrography Data for the McDermott Member

78

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APPENDIX A

Sample Location Descriptions

Table A1 – Section location and rock unit source for samples used for point counts.

85

Table A2 – Location information of rock samples (1 – 56). GRAB indicates that the rock sample was a ‘grab’ sample not from a specific section.

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Table A3 – Location information for rock samples (57 – 98).

87

APPENDIX B

Petrography Data

Table B1 – Raw point count data for various lahar matrix, reworked tuff, and sandstone deposits.

88

Table B2 – Raw point count data for various lahar clasts and pyroclastic flow deposit.

89

APPENDIX C

Paleocurrent Data

Table C1- List of bedding and cross-bedding strike and dip measured in the field, cross-bedding measurement corrected with a steronet, and resulting paleocurrent direction.

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Table C2 – Statistical Analysis of Paleocurrent Data from the McDermott Member