THE SEDIMENT ACCUMULATION RECORD AT GLACIER-FED KINASKAN LAKE: IMPLICATIONS FOR CHANGES IN HOLOCENE HYDROLOGY AND CLIMATE IN THE UPPER ISKUT RIVER WATERSHED, BRITISH COLUMBIA

by

Monique M. Stewart

A thesis submitted in conformity with the requirements for the degree of Masters of Science Graduate Department of Geography University of Toronto

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While these forms may be included Bien que ces formulaires in the document page count, aient inclus dans la pagination, their removal does not represent il n'y aura aucun contenu manquant. any loss of content from the thesis. Canada THE SEDIMENT ACCUMULATION RECORD AT GLACIER-FED KINASKAN LAKE: IMPLICATIONS FOR CHANGES IN HOLOCENE HYDROLOGY AND CLIMATE IN THE UPPER ISKUT RIVER WATERSHED, BRITISH COLUMBIA

Ms. Monique M. Stewart

Masters of Science, Graduate Department of Geography

University of Toronto, 2008

ABSTRACT

Acoustic sub-bottom profiles, ~5 m vibra-cores and lake hydrology are used to reconstruct the

sediment accumulation history in Kinaskan Lake of the Upper Iskut River watershed in north­ western British Columbia. The lake is dominated by turbidity currents in proximal regions, producing well-laminated sediments that accumulate at -4.4 mm/year. Overflows and interflows distribute silt and clay sediments to distal lake areas and accumulation decreases to

~2 mm/year, forming well-defined couplets that are interpreted as varves. Sediment accumulation chronologies cover approximately the last 2000 years (Neoglacial interval) and

demonstrate distinct patterns of above-average (or increasing) sediment accumulation thought to be associated with wetter and/or warmer conditions. This high resolution accumulation

chronology is compared with lower resolution and longer duration environment histories from nearby lakes.

u ACKNOWLEDGEMENTS

I wish to acknowledge the support and assistance of numerous people and organizations, including a generous grant from NSERC, without which this project would not have been possible.

My thesis defense committee was composed of Dr. Joseph Desloges, Dr. Sarah Finkelstein and Dr. Bill Gough. Dr. Joseph Desloges, my thesis supervisor and instructor, provided direction and guidance in the field (note: this was my first camping experience), in the laboratory and throughout the writing process. I will forever be indebted to Dr. Desloges for offering me this incredible experience. Dr. Finkelstein, my supervisor as a teaching assistance for undergraduate field course GGR 390 and second defense committee member, graciously shared her paleoclimatology laboratory with me, while providing endless motivation. The third committee member, Dr. Gough, an expert in climatology, provided a unique perspective on my research and helped me to finalize this manuscript.

Field assistance came from Melanie Grubb (University of Northern British Columbia) and from Margaret Klocker, Park Operator and was appreciated.

Assistance in the laboratory results came from several sources. Mircea Pilaf helped open sediment samples. Toni Largo's analysis of 2004 Ekman cores from Kinaskan Lake, provided interesting comparison to the 2006 Ekman results, and both Toni Largo and Nicole Chow offered guidance in GIS applications. In several instances samples needed to be processed in external laboratories. The Geography Department at Queen's University and the Department of Chemical Engineering at the University of Toronto, offered laser particle size analysis. The samples sent to Queen's University were processed under the supervision of Dr. Robert Gilbert. At the University of Toronto, Sam Roshdi trained me on using the laser particle sizer so that I could conduct my own tests. For radiocarbon dates, samples of organic matter were sent to Dr. W.E. Kieser and Dr. R.P. Beukens at the Isotrace Laboratory of the Department of Physics at the University of Toronto. Geological provenance of sediment was determined using the XRF results generated by the laboratory of Dr. Mike Gorton, Department of Geology at the

in University of Toronto. Finally, Gerald Romme and Marcel Fortin provided GIS assistance from their offices in the Map Library at the University of Toronto.

Last, but not least, I would like to extend a warm "thank you" to my family; Clayton, Gabriele, Geneve and Ian, for always encouraging me to pursue my dreams, to my overseas partner, Cesar, for his patience, understanding and love, and to all of my wonderful friends.

IV TABLE OF CONTENTS

Abstract ii

Acknowledgements iii

Table of Contents v

List of Tables ix

List of Figures x

List of Appendices xiv

CHAPTER 1: INTRODUCTION 1

1.1 Introduction 1

1.2 Research Objectives 3

CHAPTER 2: LITERATURE REVIEW 4

2.1 Introduction 4

2.1.1 Pleistocene Glaciations 4

2.1.2 Holocene Climate Conditions and Neo-glaciations 13

2.2 Pro-glacial Lake Sedimentary Archives of Climate Change 19

2.2.1 Geomorphic Controls on Sedimentation 20

2.2.2 Climate Controls on Sedimentation 22

2.2.3 In-lake Patterns of Sedimentation 25

CHAPTER 3: STUDY AREA 35

3.1 Introduction 35

V 3.1.1 Setting 35

3.1.2 Morphology 37

3.1.3 Geology 37

3.1.4 Climate 41

3.1.5 Hydrology 44

3.1.6 Vegetation 48

CHAPTER 4: METHODS 49

4.1 Field Methods 49

4.1.1 Geophysical Methods 49

4.1.2 Sediment Sampling 49

4.1.3 Water Column Profiling 51

4.2 Laboratory Methods 51

4.2.1 Core Preparation 51

4.2.2 Lamina or Bed Thickness Measurements and Dating 51

4.2.3 Loss-on-ignition 52

4.2.4 Grain Size Analysis 53

4.2.5 X-Ray Fluorescence .54

CHAPTER 5: RESULTS 55

5.1 Introduction 55

5.2 Circulation of Kinaskan Lake 55

5.3 Attributes of the Upper Iskut River watershed from Aerial Imagery 59

VI 5.4 Acoustic Record 61

5.4.1 Bathymetry, Sediment Thickness and Character 61

5.5 Sedimentology of Kinaskan Lake ..68

5.5.1 Sediment Structure 68

5.5.2 Organic Content 75

5.5.3 Particle Size 82

5.5.4 X-Ray Fluorescence 86

5.6 Chronology 88

5.6.1 Couplet Counts 88

5.6.2 Radiocarbon Dates 89

CHAPTER 6: DISCUSSION 91

6.1 Introduction 91

6.2 Sedimentation in Kinaskan Lake 91

6.3 Long-term Changes in Lacustrine Sedimentation and Climate Conditions of

British Columbia: Contextualizing the Kinaskan Lake Sedimentary Record 97

6.3.1 Selected British Columbia Lakes for comparison of sedimentary records with Kinaskan Lake 98

6.3.2 Long-term Trends in Sedimentation and Climate Conditions of British

Columbia 103

CHAPTER7: CONCLUSION Ill

7.1 Summary of Findings Ill

Vll 7.1.1 Modern Patterns of Hydrology and Sedimentation Ill

7.1.2 Long-term Sedimentation 112

7.1.3 Sediment Structure and Composition 113

7.1.4 Findings of this Study 115

7.2 Future Directions 117

LITERATURE CITED 118

APPENDICES 128

viii LIST OF TABLES

Number Page

4.1a Site characteristics and sample lengths for 2006 cores 50

4.1 b Site characteristics and sample lengths for 2004 cores 50

4.2 Characteristics of samples processed for XRF 54

5.1 Mean couplet thickness of sediments 69

5.2 Loss-on-ignition results for Kinaskan Lake cores 76

5.3 Grain size results for Kinaskan Lake cores 83

5.4 Location and characteristics of granules and pebbles found in core KIN

309 85

5.5 Percent of element or compound composing sediment samples taken from five depths in KIN 311 87

5.6 Total couplet counts in each core .89

5.6 Radiocarbon sample details for Kinaskan Lake 89

5.7 Radiocarbon results for Kinaskan Lake 90

6.1 Sites of the Northern British Columbia lakes chosen for comparison with

Kinaskan Lake 97

IX LIST OF FIGURES

Number Page

2.1 Map showing the extent of Glacial Lake Stikine 6

2.2 Map showing the pattern of Cordilleran Ice Sheet movement 8

2.3 Map showing reconstruction of Cordilleran Ice Sheet flow 9

2.4 Map showing Cordilleran Ice Sheet flow 10

2.5 Map of major British Columbian physiographic regions 11

2.6 Map showing the location of Skinny Lake, Upper Iskut River watershed 15

2.7 Site specific climate conditions of the Holocene 16

2.8 a & b (a) Timing and duration of glacier expansions during the first millennium in

British Columbia, the Yukon and Alaska, (b) Map of locations for the expanding glaciers in British Columbia, the Yukon and Alaska 17

2.9 External and internal controls on sedimentation of glacial lakes 21

2.10 Structure and flow patterns of a turbidity current 28

3.1 Map of the Upper Iskut River watershed and sub-watersheds 36

3.2 Map showing geology of the Upper Iskut River watershed 39

3.3 Map showing local geology of the Upper Iskut River watershed 40

3.4 Map showing the location of Dease Lake and Cassiar weather stations relative to Kinaskan Lake 41

3.5 Monthly temperature normals for Dease Lake 41

3.6 Monthly precipitation normals for Dease Lake 42

3.7 Monthly rainfall normals for Dease Lake .42

x 3.8 Monthly snowfall normals for Dease Lake 42

3.9 Monthly precipitation normals for Cassiar 42

3.10 Monthly rainfall normals for Cassiar 43

3.11 Monthly snowfall normals for Cassiar 43

3.12 Map showing drainage network and glacier cover of the Upper Iskut River watershed 46

3.13 Hydrological Regimes of sites along the Iskut River 47

5.1 CTD profiles of Kinaskan Lake from June 26, 2006 55

5.2 CTD profiles of Kinaskan Lake from July 26,2004 57

5.3 a & b (a) Aerial photo A20117-91 of Tsatia Creek and (b) Aerial Photo A20059-

9 of Todagin Creek 60

5.4 Map showing the bathymetry of Kinaskan Lake, core sites and locations of selected acoustic profiles 62

5.5 Acoustic image taken across-lake representing one of the thick accumulations of sediment. Location is shown in Figures 5.4 and 5.8 64

5.6 Acoustic image taken across-lake representing one of the thick accumulations of sediment. Location is shown in Figures 5.4 and 5.8 65

5.7 Acoustic image taken across-lake representing one of the thick accumulations of sediment. Location is shown in Figures 5.4 and 5.8 66

5.8 Map showing the spatial distribution of sediment thicknesses and the locations of transects shown in Figures 5.5 to 5.7 67

5.9 Mean couplet thicknesses relative to distance from the Kinaskan Lake -

Tatogga Lake delta 69

xi 5.10 a, b & c (a) KIN 309, (b) KIN 310 and (c) KIN 311 Couplet thicknesses versus depth in core 70, 71

5.10 d KIN 309, KIN 311 and KIN 310 Couplet thickness versus couplet count

(smoothed with a 25 year filter) 74

5.11 Photo showing the couplet (rhythmic) structure of sediments in KIN 310 72

5.12 Photo showing the couplet (rhythmic) structure of sediments in KIN 311 72

5.13 a & b Photos showing the sedimentary structures found in KIN 309; (a) finely laminated couplets and (b) massive deposits of clayey-silt 73

5.14 Mean Loss-on-ignition results relative to distance from the Kinaskan Lake -

Tatogga Lake delta 77

5.15 a, b & c (a) KIN 310, (b) KIN 309 and (c) KIN 311 Loss-on-ignition results versus depth in core 79, 80

5.16 a, b & c (a) KIN 309, (b) KIN 310 and (c) KIN 311 Loss-on-ignition results presented as a standardized sedimentation index versus depth in core 81, 82

5.17 Mean grain size results with distance from the Kinaskan Lake - Tatogga Lake delta 83

5.18 KIN 309, KIN 310 and KIN 311D50 Grain size with depth 84

5.19 a & b Photos illustrating (a) the site of a massive clayey-silt deposition in KIN

309 at 3.5 m depth and (b) showing one of the large clasts within this section 85

5.20 Geochemical make-up of KIN 311 88

6.1 a, b, c & d Photos of sediment structure for (a) KIN 310 (depth 3.12 m - 3.18 m),

(b) E4-06 (depth 0.05 m - 0.11 m), (c) Atlin Lake gravity core (note: photo in grey

xii scale), showing varved deposits and (d) Green Lake core (depth 1.34 m - 1.38 m, note: photo of thin section) showing annual laminae 93

6.2 a, b, c & d Photos of sediment structure for (a) KIN 309 (depth 0.9 m - 1.5 m),

(b) KIN 311 (depth 0.5 m - 0.6 m), (c) Atlin Lake gravity core showing annual varves from a low sediment input period (note: photo in grey scale), and in Green Lake core

(varves from 1900 to 1944) 95

6.3 Map showing lakes whose sediment will be compared with results from Kinaskan

Lake 98

6.4 Comparison of Loss-on-ignition Results between Several British Columbia Lakes

(including Kinaskan) 101

6.5 Comparison of Sedimentation Rates between Several British Columbia Lakes

(including Kinaskan) 102

Xlll LIST OF APPENDICES

Letter Page

A. All oblique (across-lake) transects and the long profile acoustic results for

Kinaskan Lake 112

B. Comparison Malvern and Coulter Particle Size Results 125

C. Stratigraphic Logs of KIN 309, KIN 310, KIN 311 along with Sediment

Structure and Composition Results 126

xtv Chapter 1 - Introduction

1.1 Introduction

Effects of climate change occur on various temporal and spatial scales. Timing of natural climate change has been related to external factors such as Milankovitch cycles and volcanism, and internal factors such as interactions between oceans and the atmosphere (Nesje and Dahl, 2003). Recently, it has been acknowledged by scientists and political leaders that anthropogenic activities are partly responsible for warming during the 20th and 21st centuries

(Meier, 1984; IPCC, 2007). To quantify the role of human activity in climate change, the magnitude, timing and spatial extent of major natural climate events of the past must be determined.

Historically, weather stations in Canada were positioned in populated regions, typically along the U.S. Border (The Atlas of Canada, 1974). Although these records provide high resolution data, their durations are not sufficient to show major periods of climate change.

Therefore, alternative sources of paleo-climate information have been used to highlight pre- historical climate conditions. These include documentary sources, tephrochronology, lichenometry, dendrochronology, stratigraphy and orientation of glacial landforms (i.e. moraine composition and arrangement), organic lake accumulations, peat bog deposits and glaciolacustrine sedimentary records, as methods used by scientists to identify past climate conditions (Osborn and Luckman, 1988).

Glaciolacustrine sediments can contain multi-millennia duration, high-resolution records of past climate conditions. Sediment structure and clastic composition reflect the condition of local glacier cover (tied to production and delivery of glacial sediment) which depends on ambient temperature and precipitation. Organic composition of lake sediments

1 2 provides information on autochthonous and allochthonous productivity (organic content and diatom species or counts) and types/quantities of vegetation in the watershed (pollen species or counts).

Limited information is known about the effect of climate change in northern British

Columbia due to the remote character of the region and the many fewer climate and hydrology measurements taken from there. An attempt to reconstruct the paleoclimate record for the

Upper Iskut River watershed has been undertaken using in-situ measurements and sediment samples from Kinaskan Lake, British Columbia. This site was chosen for a paleoclimate reconstruction for several reasons; (a) Kinaskan Lake has riot been the subject of any other paleoclimate study, (b) sediment samples from nearby alpine lakes (i.e. Skinny Lake and

Pyramid Lake) have contained up to 11.5 ka of reconstructed climate trends at varied resolutions, (c) the dimensions of Kinaskan Lake (area = 27.5 km2; max. depth = 135 m) ensure a long residence time for sediment and a high probability that fine sediments will settle rather than be removed, and (d) Kinaskan Lake's contributing basin is large (1,257 km2) and contains 0.75% glacier cover thus probably representing regional conditions. These glaciers directly control the hydrology and production/delivery of sediment to the lake (via Todagin

Creek). Therefore, changes in climate will be directly reflected in the sedimentary record.

When viewed in conjunction with the sedimentary record of other glacier-fed lakes of the British Columbian Cordillera, the results of this effort to reconstruct paleoclimates provide further information about the timing and extent of several major episodes of climate change, supplemented by evidence of local geomorphic events, over the late Holocene. Furthermore, this study demonstrates the necessity for continued exploration of paleoclimate archives stored in remote sites. 3

1.2 Research Objectives

This purpose of this research project is to understand depositional patterns and processes in Kinaskan Lake to create inferences about Holocene conditions in the Upper Iskut River watershed of Northern British Columbia. The primary goals include:

i. To use a variety of different methods (geophysical, water column properties, sediment

cores etc.) and results from other studies, to infer spatial variability in lake depositional

processes.

ii. To use the record reconstructed from several long cores to determine temporal

variability in depositional processes.

iii. To use the record in the cores to infer watershed and climate controls on the sediment

flux in the Upper Iskut River watershed.

iv. To compare and contrast the results of this analysis with findings from other studies to

infer Holocene depositional variability and possible regional climate controls. Chapter 2 - Literature Review

2.1 Introduction

Long and short-term climate changes, such as those responsible for glacial cycles, control the hydrology and sedimentation patterns of the Upper Iskut River watershed.

Episodes of glaciation in the British Columbian Cordillera began approximately four million years ago. At least twelve episodes of glaciation occurred over the past 1.65 Ma, a period of geological time known as the Quaternary. During this period, local expansion of alpine glaciers and the formation of confluent ice-masses molded the British Columbian landscape

(Ryder and Maynard, 1991). More recently, the widespread Cordilleran Ice Sheet formed during the late Pleistocene epoch. Since the decay of the Cordilleran Ice Sheet, a series of minor glacial cycles have occurred.

This chapter begins with a review of Pleistocene glacial events in the Canadian

Cordillera. Where possible, this is supplemented with evidence taken from sites near the

Upper Iskut River watershed. This is followed by a review of Holocene climate events up to the present. Finally, lake sedimentation processes and the use of lake sediments as climate indicators are outlined.

2.1.1 Pleistocene Glaciations

The Pleistocene began approximately 1.65 Ma BP. It has been subdivided into the early and middle Pleistocene by the Matuyama-Brunhes magnetic polarity reversal at -788 ka BP, and into the late Pleistocene by the marine oxygen isotope record at 132 ka BP

(Richmond and Fullerton, 1986; Easterbrook, 1999).

Unlike glacial cycles of the early to middle Pleistocene, evidence of the most recent advance in the Canadian Cordillera is abundant. Multiple dates have been suggested for the

4 5

start of this Fraser Glaciation in the British Columbian Cordillera. These dates, which range from 29 ka BP to 25 ka BP, mark the descent of snowline and advance of interior Cordilleran alpine and Coast fjord glaciers (Eyles and Clague, 1991; Stumpf et al., 2000; Ryder et al,

1991). Several scholars have attempted to highlight the main phases of Cordillera Ice Sheet formation. To accomplish this, stratigraphic evidence (i.e. glaciolacustrine silts and tills), glacial features (i.e. striae, erratics and moraines) and artifacts (i.e. animal bones, antlers, and preserved wood), have been taken into consideration (Prest, 1984; Ryder and Maynard,

1991). According to Flint (1957), Kerr (1934), and Davis and Mathews (1944) four phases of glacier growth led to the establishment of the Cordilleran Ice Sheet in the late Pleistocene.

The first phase was characterized by formation of small valley glaciers on the moist Coast and dry Rocky Mountains (Flint, 1957). These glaciers occupied cirques, smoothed aretes, and abraded spurs. During the second phase, the Intense Alpine Phase, these glaciers combined to form a dendritic network of valley glaciers (Flint, 1957; Ryder and Maynard,

1991; Ryder et al., 1991). The merger of these ice bodies put pressure on local asthenospheric materials which created peripheral fore-bulges and drove complex patterns of ice flow (Ryder et al., 1991). In some instances, the extension of their ice tongues into lowlands and valleys caused glacial lakes to form. For instance, in the Stikine Valley, a lake formed where advancing Coast Mountain glaciers dammed the westward flow of the Stikine

River (Figure 2.1). Glacial Lake Stikine spanned a distance of 200 km and reached elevations

of 900 m a.s.l. (Ryder and Maynard, 1991; Spooner and Osborn, 2000). Figure 2.1 Map showing the extent of Glacial Lake Stikine (Ryder and Maynard, 1991) In some cases, glaciers advanced into the deep water of these lakes. However, Ryder and

Maynard (1991) argue that the area covered by proglacial lakes was small relative to the

dimensions of glaciers. Therefore growth of the glacier was probably not stimulated by

calving along its small floating margin.

The third phase, known as the Mountain Ice Sheet Phase, involved the extension of

valley glaciers into piedmont glaciers between 21 ka and 17 ka BP (Flint, 1957). Where

piedmonts of the Cassiar Mountains, Skeena Mountains and the Stikine Plateau coalesced

with ice domes in the Coast and Cassiar mountains, an ice sheet began to develop. At the

height of this fourth phase, the Continental Ice Sheet Phase, a 2.3 km thick mass of ice

capped valleys (Flint, 1957; Ryder and Maynard, 1991; Ryder et al., 1991). From this

reservoir, outlet glaciers pushed westward through valley conduits in the Coast Mountains

(Flint, 1957). 7

Alternatively, three phases of Cordilleran Ice Sheet formation and decay have been suggested by Stumpf et al. (2000). The first phase, beginning around 25 ka BP, was triggered by falling temperatures and increasing precipitation along the Pacific Coast. Glaciers extended from accumulation centers in the Skeena, Coast, Hazelton, and Omineca Ranges, either westward to the Pacific or east-southeastward to the Nechako Plateau. Because the glacier extensions were relatively thin, their movement was restricted by local topography

(Stumpf etal., 2000).

Glaciation reached its maximum during the second phase, with coalescing valley glaciers forming an ice sheet which reached 2500 m a.s.l. (Stumpf et al., 2000). Several accumulation centers were located along an ice divide which ranged from the Nechako

Plateau to 300 km inland from the Pacific coastline. From these centers, ice flowed westward through coastal fjords and spilled onto the continental shelf, and eastward into the interior. At this point, the ice sheet was sufficiently thick that it could flow over most topography and upslope (Stumpf et al., 2000).

A shift in climate conditions ushered in the third phase. Ice domes in the interior of

British Columbia retreated towards the Skeena, Coast, Hazelton, and Omineca Ranges.

Concurrently, the westward flow of ice stopped when the surface of the ice sheet fell below the elevation of mountain passes. The rise of the equilibrium line to above the ice surface caused stagnation and decay of ice lobes in valleys (Stumpf et al., 2000).

In addition to varied theories of Cordilleran Ice Sheet formation, multiple patterns of

ice flow have been proposed. During the late 19th century, Dawson mapped hypothetical boundaries of the Cordilleran Ice Sheet with ice flowing radially outward from a central

divide between 55° and 59° N (Jackson and Clague, 1991). At the start of the 20th century, 8

Kerr, and Davis and Matthews published their own theories of ice flow (Jackson and Clague,

1991). In each instance, a focus was placed on the radial flow of ice from a central dome

rather than an ice divide. Flint (1957) built on the theories of his predecessors, suggesting the

presence of a major ice divide further south, at 53 °N. He argued that from a central dome,

located along the divide, outlet glaciers pushed through the Coast Ranges. Between the

Mountain Ice Sheet Phase and the Continental Ice Sheet Phase, the movement of ice from the

divide shifted eastward from the peaks of the Coast Ranges due to greater accumulation, and

westward from the crest of the Rocky Mountains (Flint, 1957; Clague, 1989). More recently,

Ryder and Maynard (1991; Figure 2.2) suggested that the Cordilleran Ice Sheet was

dominated by an ice dome in the Skeena Mountains, from where ice flowed in a radial

direction. They also highlighted ice divides in the Skeena Mountains, the Coast Mountains

and on the East and West sides of the Stikine Plateau.

Approximate "A" location of Kinaskan Lake

>J /,*! 0,\Jffl \ MOUNTAINS -.."VV ,

5J <[Tt Iir.jll(ni ipjtiO ::i*^~-'h ,,-v Figure 2.2 Map showing the pattern of Cordilleran Ice Sheet movement (modified from Ryder and Maynard, 1991) Alternatively, Ryder et al. (1991; Figure 2.3) argued for the presence of one ice divide at -52 °N, from where ice flowed relatively directly northward or southward.

Approximate -^ location of KinaskanLake

Figure 2.3 Map showing reconstruction of Cordilleran Ice Sheet flow (modified from Ryder et al., 1991)

Specifically, Ryder et al. (1991) argued that there is insufficient evidence of a radial ice-flow

pattern as would be associated with an ice dome.

Finally, Jackson and Clague (1991) offered a diagram of ice flow which combined

patterns suggested by Crandell (1965), Lemke et al. (1965), Richmond et al. (1965), Prest et

al. (1968), Hamilton and Thorson (1983), Porter et al. (1983) and Wilson (1958). In the

image, several domes provide circuitous flows of ice (Figure 2.4; Figure 2.5). 10

Approximate \k location of KinaskanLake

Figure 2.4 Map showing Cordilleran Ice Sheet flow(modifie d from Jackson and Clague, 1991) 11

Figure 2.5 Map of major British Columbian physiographic regions (modified from Mathews, 1986 in Ryder and Maynard, 1991)

Despite the ongoing debate, several major ice streams in the Cordilleran Ice Sheet have been universally acknowledged. One stream flowed northward through the southern

Cassiar Mountains, bearing east at the Liard Lowlands. An additional stream, originating north of the Finlay River, moved into the Liard Lowlands. South of the Finlay River, ice moved eastward into the Rocky Mountain trench. A fourth stream moved northwest into the

Yukon, forming the Cassiar Lobe. Moving south, multiple streams crossed the Nass

Depression and the valleys of Takla and Babine Lake. Finally, from the ice accumulation zone in the northern Cassiar Mountains, an ice stream moved southwest towards the Teslin ice stream and northeast towards the Liard stream (Ryder and Maynard, 1991; Ryder et al.,

1991; Figures 2.2 , 2.3, 2.4 and 2.5).

After approximately 10,000 years of glaciation in the mountains, and 8,000 years of ice cover, a shift in climate conditions caused glacial decay along the coastline (13.5 ka to

11.5 ka BP) (Stumpf et al., 2000; Ryder et al., 1991). Eustatic sea level rise, combined with 12 subsidence of coastal land under the weight of overlying glaciers, caused sections of the ice sheet to extend into the ocean and calve. As coastal areas were unloaded, land began to rebound. Because this occurred in an irregular fashion, crustal deformation did not occur uniformly (Ryder et al., 1991). This was closely followed by deglaciation of the British

Columbian interior. First, high elevations became ice-free through downwasting. In the mountains, a layer of silty-sand till with occasional angular clasts was deposited. On nearby plateaus, a thick layer of till was laid down and glaciofluvial sediment was transformed into

drumlins and rock knob tails by the active ice flow. Eskers, outwash and kame terraces also appeared (Ryder et al., 1991). As glaciers in higher terrain receded, mounds of stagnant ice were left to decay in valleys and lowlands (Ryder et al., 1991; Johnson, 1992). According to the amount of debris supported by ice, the interaction between the ice and debris, and the

availability of melt-water, a variety of stagnant ice landforms developed (Gravenor and

Kupsch, 1958; Eyles et al., 1999; Shaw, 1994; Winters, 1961). Accumulation of supra-glacial

debris (loose gravel and cobbles) along zones of ablation resulted in hummocky moraines where debris collapsed through openings in the melting ice (Eyles et al., 1999; Gravenor and

Kupsch, 1958). Pressing of sub-glacial sediment into basal crevasses of the glacier by the

weight of overlying ice created ridges and troughs which are expressed on the landscape as

rim ridges and terrace ridges (moraine plateaux or plains plateaux), ice-squeezed plains

ridges, till-cored or completely till esker ridges, minor moraine ridges and drumlinoid ridges

(Eyles et al., 1999; Gravenor and Kupsch, 1958; Stalker, 1960). Examples of these stagnant

ice features are ubiquitous across the Upper Iskut River watershed.

Before its completion, this Fraser deglaciation was interrupted by several pauses and

readvances. In addition to the Sumas advance in the Fraser Lowland, recessional moraines 13 have been identified in the Omineca Mountain valleys and it has been suggested that cirques in the Atlin region were periodically reoccupied (Ryder and Maynard, 1991; Ryder et al.,

1991). In the Upper Iskut River watershed, little information is available about the

Quaternary history (Ryder, 1987; Spooner et al., 2002). It has been suggested that an ice divide in the Cordilleran Ice Sheet may have existed near the watershed (Ryder and

Maynard, 1991; Spooner et al., 2002). Deglaciation around the Upper Iskut River began in the plateaus and left stagnant ice in the valleys. By 9.7 ka BP, alpine regions were ice free.

Several hundred years later, large valleys were free of ice (Spooner and Osborn, 2000;

Spooner et al., 2002; Mazzucchi, 2000).

2.1.2 Holocene Climate Conditions and Neo-glaciations

The Holocene, which encompasses approximately the past 10,000 years, is

characterized by series of climate warming and cooling periods. Using oxygen isotope

records from Greenland ice cores and ocean sediments, North Atlantic ice-rafting debris

marine records and cosmogenic nuclide records as a guide, various cyclic frequencies of

climate change; 400-500 years, 900-1,150 years, 1,470 ± 500 years, and 2,500 years, have been suggested (Viau et al., 2006; Nesje and Dahl, 2003). Causes of these cycles include

external forcing factors, such as Milankovitch cycles and volcanism, and internal forcing

factors such as ocean-atmospheric interactions (Nesje and Dahl, 2003). Due to asynchronies

between glacial behavior of different regions, difficulties generating accurate dates for glacial

activity, and leading or lagging glacial response times to climate change, chronicling the

Holocene glacier and climate change history in British Columbia is extremely complex.

Focusing on the Cordillera, we can identify a series of minor glacial advances and retreats

that characterize the impact of Holocene climate change in this region. These records have 14 been established using radiocarbon dates of end and lateral moraines, documentary sources, lichenometry, dendrochronology, tephrochronology, organic lake deposits, peat bogs and glaciolacustrine sedimentary records (Osborn and Luckman, 1988).

An episode of warm and dry climate conditions dominated the first two millennia of the Holocene (Osborn and Luckman, 1988; Deevy and Flint, 1957; Viau et al., 2006). The timing and extent of this Holocene Thermal Maximum (Hypsithermal) varied greatly. At

Skinny Lake (and Kinaskan Lake), the Hypsithermal began 7.2 ka BP and terminated 4.8|ka

BP (Spooner et al., 2002). Warm climate conditions of this interval across North America

caused a portion of the Laurentide Ice-Sheet to collapse at approximately 8.2 ka BP. This

created a pulse of fresh melt-water which flowed into the North Atlantic and disrupted the

formation of North Atlantic Deep Water and thermohaline circulation (Viau et al., 2006).

Falling temperatures and an increase in precipitation, tied to decreased insolation,

marked the close of the Hypsithermal and the start of the Neoglaciation (Mann et al., 1998).

Around 6.8 ka BP, the Crowfoot Glacier began to expand in the Canadian Rockies, leaving

behind rock glacier deposits and cirque moraines. Sections of till were capped by tephra from

a Mazama eruption at 6.8 ka BP (Osborn and Luckman, 1988). Between 6 ka and 3 ka BP,

cirque moraines were created by advancing glaciers at Dunn Peak and in the "Garibaldi

Phase" of the Coast Mountains (Osborn and Luckman, 1988). Around 4 ka BP, increased

snowfall over the Upper Iskut River watershed buried caribou antlers (Ryder, 1987). Ryder

(1987) suggests that the region remained cool or moist until recently, preserving the antlers

until their recent discovery at the confluence of the Iskut and Ningunsaw Rivers. A study of

lake sediments from Skinny Lake, also located in the Iskut region, supports the theory of

climate cooling and increased precipitation starting at 4.8 kaBP (Spooner et al., 2002; Figure 15

2.6; Figure 2.7). Outside of the Upper Iskut River watershed, similar conditions have been noted. In the Purcell Mountains, advance of the Bugaboo Glacier left behind tills (Nesje and

Dahl, 2003). In the Mount Tiedemann area, glaciers reached their maximum Neoglacial extents at -2940 14C years BP (Nesje and Dahl, 2003; Spooner et al., 2002). Additionally, in the northern Coast Mountains of British Columbia, the Frank Mackie Glacier and Berendon

Glacier are believed to have advanced at 2.8 ka BP (Clague and Mathewes, 1996).

** Location of KinaskanLake

Figure 2.6 Map showing the location of Skinny Lake, Upper Iskut River Watershed (modified from Spooner et al., 2002) 16

Figure 2.7 Site specific climate conditions of the Holocene (T = Temperature; P = Precipitation) (modified from Spooner et al., 2002)

Between the close of the Neoglacial and the start of the Medieval Warming Period (~

AD 700), glaciers positioned along a -2,000 km stretch of coastal mountains belonging to

British Columbia, the Yukon Territories, and Alaska, experienced a short-lived expansion.

Evidence for this is in the Upper Iskut River watershed where Forest Kerr Glacier overrode a mature forest between 100 and 500 AD (Nesje and Dahl, 2003; Reyes et al., 2006) (Figure

2.8 a &b). 17

500 e.C./A.D. 300 to» woo I ' ' • • I ' ' • "• I "" • i ' • • • i Alaska / 7 ^-•ww^ ^«£ & 'WL 10 / f 1 fiSf ^ NXII y Yukon 1 Territay ) BriclBt CS3* v / AS / V^a \ «^/V tv MrarMM • w C3^ / Friink Msck'n 1-Lillooe1.Bridge 2-Miserable 7 vo6 Toddj, 3-Tiedemann Vi 4-Frank Mackie Pacific jWHj Suiprtul So 5-Todd.Surprise Ocean Eh 3 6-ForestKerr jp British FoilM Kert r-Beare £* Columbia 8-SJierWan 9-Niana "\^/ / 10-Kuskulana E*2 Srmidan ism* 11-Copper 1i .-""i* 12-Tebenkof, Ntona Barftett 13-Grewingk i&i*h I 14-Dinglestadt i Kuitaitema I j 11 Copper • |^^_400^ Taberfwf wi BHtMt r— Gmvingk a. OinglesMI Mud Glacier, Great Glacier and Flood Glacier 500 BC.IKD 500 1000 CaUbnud yaor A.O.* C

Figure 2.8 a & b (a) Timing and duration of glacier expansions during the first millennium in British Columbia, the Yukon and Alaska (Reyes et al., 2006), (b) Map of locations for the expanding glaciers in British Columbia, the Yukon and Alaska (modified fromReye s et al., 2006).

The term, "Little Ice Age" (LIA) was introduced by Matthes (1939) in reference to glacial advances which followed the Medieval Warming Period (Medieval Climate

Optimum) (Luckman, 2000). Since then, the term Neoglacial has replaced the "Little Ice

Age" in its original definition and LIA now refers to Late Neoglacial advances (Luckman,

2000). Synchronous glacial advances in the Canadian Rockies and the British Colombian

Cordillera occurred between the 12th and 13* centuries, during the 18th century and during the 19th century. Regional dendrochronology records show that glacial advances followed periods of increased precipitation and low temperatures (Luckman, 2000). Within the Rocky 18

Mountains, dendrochronology, radiocarbon dates, lichenometry and documentation have been used to isolate several clusters of years during which glaciers reached their LIA maxima: 1700 to 1725 AD and 1825 to 1850 AD (Luckman, 2000).

In the Stikine-Iskut region of British Columbia, terminal moraines at the three largest glaciers: Great, Flood and Mud, date from the 17th to the 18th centuries (Ryder, 1987). Since reaching their late-Neoglacial maximum positions, recessional moraines have formed.

According to the age of forest stands on the moraines, dates from overridden tree stumps, organic materials from local soil horizons and till stratigraphy, the moraines formed recently.

At More Creek Basin, two glaciers have left morainal ridges and fore-fields which contain conifers and willows. The ages of these trees suggest germination in the late 19th to early 20th century (Ryder, 1987). At all three of the glaciers, an inner moraine bears a young forest, with the oldest hemlock and spruce identified as 200 years old. The outer moraine has a mature coniferous forest which colonized in the late 17th to early 18th century (Ryder, 1987).

In the northern Coast Mountains of British Columbia, glaciers reached their LIA maximum in the 17th century (Clague and Mathewes, 1996). This was followed by several brief readvances suggested by several small moraines and thin silt deposits in Berendon Fen

(Clague and Mathewes, 1996). Alternatively, dates for moraines in the Canadian Rockies range from the 16th to the 19th centuries suggesting a series of glacier advances culminating in slightly asynchronous glacier maxima (Clague and Mathewes, 1996).

At the start of the 20* century, glacier recession in the British Columbia Cordillera was rapid. From 1950 to 1970 the rate slowed. By the late 20th century, glaciers had lost over

25% of their area (Luckman, 2000). Glacier advances were due to increased winter precipitation and/or decreased summer temperatures (Osborn and Luckman, 1988). However, 19 other factors could also be important. A study of the Grand Pacific Glacier and the Melbern

Glacier by Clague and Evans (1993) revealed a trend of advancement by the Grand Pacific

Glacier but a retreat by the Melbern Glacier during the past century. It is believed that a shift in the ice divide between the two glaciers, related to differences in the steepness of each glacier and the size of their calving front, caused this inequality (Clague and Evans, 1993).

Alpine environments, such as the Upper Iskut River watershed, are highly sensitive to climate warming. Due to their remoteness, archives of climate change stored in the lake sediments, landscape and vegetation of these sites have been neglected. It is important to understand the local and large-scale effects of climate change. Warmer conditions promote glacial melt, mobilization of sediment and shallowing of the permafrost active layer. Where combined with greater precipitation, flooding and mass-wasting can occur. Within the Upper

Iskut River watershed, these events are possibly recorded in lake sediments as massive deposits of poorly sorted sediment. At Pyramid Lake, north of Kinaskan Lake, a series of minerogenic layers were deposited by hillslope failures between 5.1 and 4.4 ka BP. These have been tied to changes in atmospheric circulation and increased precipitation (Mazzucchi et al., 2003). On a global scale, glacial melt outside of Greenland and Antarctica has caused one-third to one-half of sea level rise in the 20th century (Meier, 1984; Meier and Wahr,

2002; Clague and Evans, 1993). Of this amount, more than one-third originated from glaciers along the Gulf of Alaska (Clague and Evans, 1993).

2.2 Pro-glacial Lake Sedimentary Archives of Climate Change

Where supply of sediment is not a limiting factor, lake sediments, particularly those extracted from glacier-fed lakes, can contain archives of climate changes and extreme events on millennial timescales with annual resolution. The volume of sediment entering a lake 20 depends on geomorphic controls, hydroclimate, and the propinquity of depositional sites and glaciers (Luckman, 2000).

2.2.1 Geomorphic Controls on Sedimentation

Lacustrine sediment can be characterized as clastic, chemical or biogenic. Clastic sediments are discrete particles of allochthonous material such as sand, silt and clay, whose composition can be classified by mineralogy and lithology (Schnurrenberger et al., 2003).

For a given lake, the proportion of clastic sediment depends on supply and transport.

Lakes receive clastic sediment through various processes. Sorted clastic sediment enters through aeolian transport or as the suspended, saltated or bed-load of rivers, taking the form of turbidity and overflow currents. Unsorted clasts enter via ice-rafting and coastal colluvial action (Winsemann et al., 2004; Hodder et al., 2007). Fluxes in clastic inputs from contributing watersheds are related to supply controls such as deflation, regional denudation, mass-wasting, vegetative growth or reduction, and tectonic activity (Drewry, 1986; Hodder et al., 2007). In glaciated regions, sediment supply is highly linked to the availability of loose till, outwash and ice-rafted debris. Consequently, paleolimnological analyses use particle size and geological origins to interpret historical transport and deposition conditions (Singer et al.,

1988; Komar et al., 1984; Swift et al., 1972; McCave et al, 2006; Hakenson and Jansson,

1983; Gilbert and Butler, 2004; Mazzucchi et al., 2003; Gilbert et al., 1997). Sediment load is a function of precipitation, glacier and snow ablation, and other hydrologic conditions within the drainage basin which include glacier extent (Luckman, 2000). Particle size measurements provide clues of transport energy, and mineralogy of a deposit suggests provenance and thus the distance and energy involved in transportation (Schnurrenberger et al., 2003; Menounos et al., 2006). 21

The sedimentation pattern within a lake is determined by input locations, basin morphology, water circulation, biota and chemistry. Morphological characteristics of a lake that can affect depositional patterns are depth (e.g. sills, deep basins etc.), the basin area to lake area ratio, and the lake area (Figure 2.9) (Desloges and Gilbert, 1994).

Boundary conditions External Forcing of the basin - weather (temperature. - lithology/siructure precipitation) - morphometric • hydrology (average and • glacial and nival extreme runofl regimes} - vegetallonal - human Internal Forcing -tectonics • votcanlsm -^•IWater and Sediment YJiidl-

Lacustrine Basin -morphology - limnology (mixis. circulation, throughput)

Glaciolacustrine Depositional| Processes * Characteristics of the Deposits

Figure 2.9 External and internal controls on sedimentation of glacial lakes (from Desloges and Gilbert, 1994) Where lakes have maximum depths extending into an aphotic hypolimnion, sediment

focusing is followed by minimal post-depositional disturbance. As a result, deep lakes have

the potential for preserving a continuous (undisturbed) supply of sediment during late glacial

and Holocene environmental conditions (Desloges and Gilbert, 1994). Where sub-basins are

created by the presence of a sill, sediment-laden underflows can sometimes pass from the

proximal to the distal basin. This occurs where: (a) the thickness of the underflow exceeds

the sill height, (b) the inertia of the underflow is sufficient to drive it upslope, or (c) where

interaction of the underflow with up-lake compensation currents adds buoyancy to the turbid

mass and the turbidity current spills over the sill (Chikita et al., 1996).

Lake and watershed area impact the quantity and distribution of sediment.

Specifically, a ratio between basin area and lake area changes in concert with up-basin 22 storage effects such as in floodplains, lakes, wetlands etc. As the ratio decreases, there are more direct exchanges of material between the receiving lake and its hill-slope and glacier systems (Desloges and Gilbert, 1994). For example, Menounos et al. (2006) found an increase in specific sediment yield at Green Lake, British Columbia up to a basin scale of 3 x

104 km2. After this threshold, there were greater inputs of reworked Quaternary sediments relative to those formed by primary denudation (Menounos et al., 2006).

Lakes with a large surface area (>10 km2) are effective sediment traps. Large lakes typically have deep lake floors which are protected from intensive circulation and slope failures. Therefore, fine sediments accumulate relatively free of disturbance (Embleton and

King, 1968; Gilbert et al., 2006). Conversely, sedimentation can be limited if lake size increases to the point where the basin is unable to provide sufficient load to form distinct deposits, unless there is adequate local glacier coverage to drive erosion and seasonal runoff and the formation of distinct annual deposits (Desloges and Gilbert, 1994, 1998).

2.2.2 Climate Controls on Sedimentation

Paleoclimatological reconstructions ideally require long-term, undisturbed sedimentary sequences with annual to sub-annual sediment inputs. Furthermore, they need a contributing watershed which is sensitive to atmospheric controls (Menounos et al., 2006).

Glacier-fed lakes receive sediment inputs in which the volume and characteristics of the sediments are related to non-linear relationships and feedback mechanisms between climate, glaciers, fluvial activity, geomorphic controls, terrestrial biota and lacustrine processes

(Hodder et al., 2007).

Climate controls on glacially-dominated systems include seasonal peaks of runoff which transport high volumes of material into the lake basin. These deposit a layer of light- 23 colored sediments during the summer and dark colored sediments during the winter

(Desloges and Gilbert, 1994). Varved sediment in glaciated watersheds of western Canada typically contains a record of 100 to 1,000 years (Hodder et al., 2007). Where protected from disturbance, annual accumulations form a two lamina ("couplet") or three lamina sedimentary unit (Desloges and Gilbert, 1994; Ojala and Saarnisto, 1999). A couplet is composed of: (a) silts and clays deposited under calm lake conditions characteristic of autumn and winter and (b) fine sand and silt transported by peak flows of nival melt and precipitation events during the spring and summer. Contributions of silt to the lake occur during the summer when glacial melt-water enters a warming lake as an interflow (Desloges and Gilbert, 1994). Conversely, a three lamina pattern is evident in Holocene sediments from

Finish lake Valkiajarvi (Ojala and Saarnisto, 1999). During spring floods, a layer of clastic matter is deposited. Sediment grades upwards into an organic-rich lamina, created between late spring and early autumn. In winter, the lake is ice-covered and fine suspended sediment is deposited in a low energy setting (Ojala and Saarnisto, 1999).

Non-annual rhythmites form due to a mix of climate and non-climate driven variables. The collapse of glacial dams and subsequent release of anomalous discharges, summer and autumnal storm runoff, mass wasting events, and warm or cold episodes can all occur irregularly (Embleton and King, 1968; Luckman, 2000; Desloges and Gilbert, 1994).

Glaciers are a major sediment source to lakes in mountain environments (Drewry,

1986; Menounos et al., 2006). Glaciers create and mobilize sediment through ice shear and entrainment by subglacial melt-water. Along glacial margins, sediment-enriched water is released. This typically peaks at the start of the glacial melt season. In addition to the close relationship between glacial melt and sediment release, it has been suggested that greater 24 glacial extent can be inferred from thicker varves (Hodder et al., 2007). This becomes complicated where extreme events, including jokulhlaups, introduce large volumes of surging water into a lake. In a study of Chilko Lake sedimentation, Desloges and Gilbert

(1998) attributed extreme variability between sediment yield and inferred glacial cover to episodes of similar outburst floods. Another consideration occurs where glacial tongues extend into a lake followed by calving and rafting of glacial ice debris which produce anomalously coarse and/or thick deposits (Hodder et al., 2007; Embleton and King, 1968).

Glacier erosion is closely related to the lithology of the substrate. Easily eroded bedrock and unconsolidated surface material can form waves or "superslugs" which slowly progresses through glaciated drainage basins. From these slugs, sediment is slowly released, adding to yields from contemporary glaciers and enhancing accumulation rates in downstream water-bodies. While a "superslug" can be centered at any site in the watershed, storage of sediment in glacier forefields can have an opposite effect on sedimentation rates

(Menounos et al., 2005). Furthermore, up-stream lakes can trap sediment indefinitely

(Toniolo and Schultz, 2005).

Although fluvial systems in glaciated watersheds can store sediment or delay transport, they are a major mobilizer of glacier sediment and are controlled by glacier hydrology (Drewry, 1986). Discharge from glacier melt can vary on diurnal, sub-seasonal, seasonal, and annual cycles (Drewry, 1986). Specifically, the melt rate is tied to air temperature and the timing and amount of precipitation. Snow fall during the winter does not melt until spring and summer. Late in the summer, the rate of nival ablation appears to increase because snow is water-saturated. This produces a greater amount of melt-water per volume of snow (Hodder et al., 2007). An additional pattern of discharge is the tendency of 25 low flows to occur as a series of multi-year low flows and for high flows to cluster into a multi-year series of high flows (Hodder et al., 2007). As the melt-water flows through a valley system, channels, floodplains and lakes can trap sediment (Drewry, 1986; Hodder et al., 2007). Where this sediment is subsequently released, it appears as a lag in sediment transport. If the mass is not released, truncation in the sediment series occurs (Hodder et al.,

2007). This suggests that there may be a complex year-to-year "memory" in the accumulation series.

Terrestrial biota, including humans, can influence rates of sedimentation in several ways. Dense covers of vegetation can intercept precipitation and hold soil. Deforestation by humans combined with grazing and trampling by animals can facilitate erosion (Hodder et al., 2007). Fortunately for paleoclimatological research, glacial watersheds are typically more isolated from human disturbances than their tropical or temperate counterparts (Desloges and

Gilbert, 1994)

2.2.3 In-lake Patterns of Sedimentation

Once sediments enter a lake, they can follow several paths towards depositional sites.

Investigations of these mechanisms and processes have taken two approaches. First, studies have interpreted and created inferences from the sedimentary properties of deposits in the lacustrine basin. Second, measurements of water column and sediment interactions have been conducted on a real-time basis.

Lacustrine sedimentation is tied to particle size and shape, density, concentration, viscosity and turbulence of incoming flows (Drewry, 1986). The role of particle size and shape is complicated by formation of clay and organic-rich aggregates, known as floes, and scavenging of small particles by larger ones (Hodder and Gilbert, 2007; Hakenson and 26

Jansson, 1983). In still and isopycnal water, a particle will settle from suspension when the net gravitational force related to the particle's submerged weight is balanced by the drag force. However, natural bodies of water are characterized by circulation and differences in water density which affect the distance traveled by particles before sedimentation and their rate of descent (Drewry, 1986; Hakenson and Jansson, 1983).

Currents are created by wind shearing against the lake surface, differences in barometric pressure and movement of calved ice blocks (Drewry, 1986). Where glaciers and ice caps are located near the lake, katabatic winds can create oscillating waves along the length of the water-body (Drewry, 1986). In the northern hemisphere, the Coriolis effect, which is caused by earth's rotation, can cause right-handed deflection of currents. This can result in greater sedimentation on western lake shores where fluvial inputs are from the north or greater sedimentation on eastern lake shores where fluvial inputs are from the south

(Smith, 1981; Drewry, 1986; Hakenson and Jansson, 1983). Examples of this phenomenon have been observed at Nicolay Lake, Canadian High Arctic, Atlin Lake, British Columbia and Meziadin Lake, British Columbia (Lamoureux, 1999; Serink, 2004; Gilbert and Butler,

2004). Bates (1953) noted that this effect decreases in significance as flow decelerates.

Alternatively, at Bowser Lake, British Columbia, high velocity turbidity currents were able to overcome this effect (Gilbert et al., 1997).

The three main variables which cause the density of inflowing stream water to differ from ambient lake conditions are: temperature, salinity and suspended sediment concentration (Smith, 1981). Depending on the density relationship between ambient water and the inflow, the inflow can be; hyperpycnal (underflow), homopycnal (interflow) or hypopycnal (overflow) (Drewry, 1986). According to Drewry (1986), the relationship 27 between these three types of inflows defines patterns of lacustrine sedimentation. However,

in each instance there is a systematic decrease in flow competence with distance from the

flow entry point. Where the density of incoming stream flow is significantly greater than lake

water, turbulent exchange and mixing between the two fluids is labored and the stream enters

the lake as a discrete gravity current (Bates, 1953; Drewry, 1986). Middleton and Hampton

(1976) suggest that hyperpycnal flows can take one of four forms: turbidity currents, grain

flows, liquefied/fluidized flows, and debris flows. The difference in hydrostatic pressure

between the each flow and ambient fluids causes a rising current (Allen, 1970; Mulder and

Alexander, 2001). As the flow extends deeper into the lake, the velocity gradient between the

two fluids decreases and distinct deposits are left on the lake bed (Bates, 1953). At Bow

Lake, Alberta, thick varves are interpreted to have resulted from a series of underflows

related to climate conditions and glacial melt (Smith, 1981).

Deposits from turbidity currents have been identified in lake sediments from Atlin

Lake, Bowser Lake and Meziadin Lake, all in British Columbia (Serink, 2004; Gilbert et al.,

2006; Gilbert et al., 1997; Gilbert and Butler, 2004). These cohesive plumes of water and

sediment (typical sediment concentration <9%) form where there are moderate to steep

slopes, deposits of unconsolidated material such as recently failed sediment with excess pore

pressures, overloading, and episodes of seismic activity (Allen, 1970; Bagnold, 1954, 1962).

Steady-uniform turbidity currents receive a fixed supply of dense fluid from upstream (Allen,

1970; Best et al., 2005; Nemec, 1990). When this fluid is released from a confining canyon

or channel walls and reaches an open and fiat area, it spreads laterally along the bed. Event-

driven turbidity currents can originate from wave forcing, tidal action, storms, wind-driven 28 circulation, high river discharge events and retrogressive slope failure (Best et al., 2005;

Middleton and Hampton, 1976; Mulder and Alexander, 2001).

Turbidity currents typically have a distinctive head, body and tail. The head is slow moving, thick and lobate, with an overhanging front due to frictional resistance at the base of the current (Middleton and Hampton, 1976). As water fills spaces between lobes of the head, sediment concentrations are diluted (Allen, 1970; Middleton and Hampton, 1976).

Turbulence off of the neck creates eddies of suspended sediment and water that escape across the flow ceiling and base (Middleton and Hampton, 1976; Figure 2.10).

ft/J Schematic Subdivision of a Turtridtty-Currftnt Figure 2.10 Structure and flow patterns of a turbidity current (Middleton and Hampton, 1976) The head can be an active agent of erosion depending on the shear stress of the turbidity current, the type of substrate and the amount of time the head is in contact with a certain area of substrate (Allen, 1970). The body has a relatively uniform thickness and flow velocity while the tail is characterized by rapidly decreasing flow and sediment concentration

(Allen, 1970; Middleton and Hampton, 1976).

Erosion by turbidity currents can leave scour marks and tool marks on the lake bed.

Scour marks, including flutes, form when an eddy is trapped below the flow. As it spins on a 29 quasi-horizontal axis oblique to the flow, a scour is drilled into the substrate (Middleton and

Hampton, 1976). Tool marks, including grooves, striations, prod marks and bounce marks, form when clasts are dragged linearly across the lake bed. The size of the clast depends on the velocity of the current (Middleton and Hampton, 1976). Deposits by turbidity currents can liquefy where sediments become less dense with depth. If the viscosities of the two materials are almost equal, symmetrical sinuous warping can occur within the deposits. If the viscosity of the lower material is greater than the upper material, a pattern of blunt mounds of sand interrupted by "flames" of upward reaching mud can form (Middleton and Hampton,

1976). Deposits can be sub-categorized as distribution graded (characterized by decreasing particle size from the base to the top of the deposit) or as coarse-tailed (sorting limited to coarse particles)(Middleton and Hampton, 1976).

The author found no incidence of grain flows in the literature of glacier-fed lakes in

British Columbia. These flows can occur in a viscous regime, with the fluid responsible for maintaining dispersive pressures, in an inertial regime, where grain inertia is paramount, or in a transitional regime (Bagnold, 1954; Middleton and Hampton, 1976). They develop on slopes which exceed an angle of initial yield, whose value depends on porosity, grain shape and the type of interstitial fluid. For viscous or transitional regime flows with a fluid that is denser than water, buoyancy reduces the effective weight of particles and the threshold angle of initial yield (Middleton and Hampton, 1976). As surface grains flow off the slope as a dilated mass, sub-surface grains assume a shallower angle (Leeder, 1982). During down- slope movement, equilibrium is reached between the weight of the particles and the transfer of grain momentum onto the shear plane. This is responsible for the near-parabolic velocity 30 profile of actively shearing grains above the shear plain and cap of non-shearing grains

(Middleton and Hampton, 1976; Leeder, 1982; Lowe, 1976).

Grain flows leave most lake beds smooth in their wake. Occasional sole markings take the form of leaf-like marks, ropy sole marks, load structures and slide marks (Middleton and Hampton, 1976). Depositional structures form where driving stress is unable to maintain velocity. Deposits are laterally and vertically inversely graded. This may be due to dispersive pressures in the flow which push coarse grains to the flow surface where there is reduced shear strain (Bagnold, 1954; Leeder, 1982). Alternatively, large grains in the flow may act as a sieve, filtering fine grains which accumulate near the shear plain and pushing larger grains further upwards (Middleton, 1970; Leeder, 1982).

Middleton and Hampton (1976) define fluidized flows as mass movements that develop from liquefied sands ejecting pore water. Evidence of fluidized flows in literature of glacier-fed lakes in British Columbia also eluded the author. These transient masses have high viscosity and low strength, enabling them to travel down slopes as shallow as 3° - 10°.

Where pore fluids are expelled, the flowing mass solidifies from the bottom to the top

(Middleton and Hampton, 1976). Lowe (1976) distinguishes between fluidized flows, where sediment is fluidized by its own escaping interstitial fluids, and liquefied flows. He argues that fluidized flows cannot exist, but that liquefied flows are common. Liquefied flows form where liquefied sediment fails. This can be triggered by erosion and undercutting, earthquakes, collapse of steep slopes, seismic activity, excessive hydrostatic pressures from artesian waters, tidal fluxes, wave and current activity, and the release of interstitial water from rapidly deposited sand (Lowe, 1976). Because of the mass' low resistance to shear, it can flow down shallow slopes and transform topography while undergoing slight translation, 31 agitation and widening of the liquefied front. Sedimentary characteristics; texture, composition and diversity of particles, the bed slope form, the type of failure, the extent of liquefaction and the release of excess water determine the flow's velocity (Lowe, 1976).

Decay of a liquefied flow occurs where extreme turbulence changes a flow into a high-density turbidity current, or where weak turbulence causes coarser materials to settle out of suspension. This transforms the remains of the flow into a low-density turbidity current.

Deposits created by the transformation of a liquefied flow into turbidity currents include turbidites, fluxo-turbidites and thick beds of coarse debris (Lowe, 1976). Where liquefaction is heterogeneous due to the presence of clay, gas bubbles and organic matter, shear resistance and flow viscosity are enhanced (Lowe, 1976). As pore water diffuses out of the flow it can become siphoned into pipes at sufficient velocities to entrain particles (Leeder, 1982). When enough water has been removed from the flow, the entire mass solidifies. Typical deposits include dish structures, concave-upward layers of massive to lightly laminated sands and pillars, and vertical peaks of sand that pierce the sub-horizontal layers (Middleton and

Hampton, 1976; Leeder, 1982).

Sub-aqueous debris flows are cohesive slurries of water and sediment with at least as much sediment as water by volume and more sediment than water by weight. In Saanich

Inlet, British Columbia, sub-aqueous debris flows were triggered by an earthquake (Blais et al., 1997). These flows move readily under gravity, taking the form of a wave or surge and entrain local sediment through dispersion or traction (Toniolo et al., 2004; Ilstad et al., 2004; de Scally et al., 2001; Middleton and Hampton, 1976). Their composition depends on the materials available for mobilization. Coarse-grained flows rely on grain-to-grain interaction for movement, whereas fine-grained flows are controlled by electrostatic inter-particle forces 32

(Ilstad et al., 2004). Debris flow rheology is a function of the concentration of suspended

solids, cohesion, grain friction, particle sizes and shapes and pore pressure (Imran et al.,

2001; Toniolo et al., 2004; Mulder and Alexander, 2001).

Similar to turbidity currents, debris flows have a characteristic head, body, tail and

rigid plug. The head forms from the slurry pushing through ambient water (Marr et al.,

2001). For rapidly moving flows, the head slides upon a thin layer of water. Behind the head

trails a body and tail with decreasing sediment concentration. If a debris flow has low

viscosity, the head can stretch out the body and tail. Where the slurry is stiff, the head can

separate from the body (Toniolo et al., 2004; Ilstad et al., 2004). In some cases, the body rafts

a rigid plug either atop or just below the surface of the flow. The thickness of the plug is

directly related to the strength of the flow and inversely related to the slope angle and flow

density. If the plug expands to the thickness of the flow, the current will stop (Middleton and

Hampton, 1976).

The erosive potential of a debris flow depends on its rheology and whether it is

hydroplaning. Stiff flows, traveling across a smooth and wet lake bed, are liable to pull-

aparts and blocking. Pull-aparts occur when a debris flow loses friction at its base and its

body is torn into two parts. Blocking involves a separation slide ahead of the main flow. The

resulting flow segments are more effective agents of erosion than a coherent unit, creating

low-relief striations and grooves in the substrate (Ilstad et al., 2004; Middleton and Hampton,

1976). Ilstad et al. (2004) identified three additional erosive patterns of debris flows. The first

pattern involves movement across a bed where there is constant contact between grains and

the bed surface. The second situation has a partly liquefied flow with grains intermittently in

contact with the bed. Finally, the third situation involves blocking and hydroplaning. 33

Deposition from large debris flows can expose buried pipelines, dislodge offshore oil rigs, and trigger tsunamis (Toniolo et al., 2004). For debris flows with an approximately normal distribution of particle sizes, deposits near the source are composed of coarse particles. Towards the site of final settlement, deposits become increasingly fine and develop grain-size laminations with normal or inverse grading (Middleton and Hampton, 1976).

The effects of discrete gravity currents tend to be localized. Where gravity currents diffuse, beds of fine and well-sorted sediment have been found across distal regions (or regions up-lake from sills) of pro-glacial lakes (Embleton and King, 1968; Gilbert et al.,

1997). Regardless, most sedimentation in thermally and/or chemically stratified glacier-fed lakes results from a combination of interflows and overflows.

Glacier-fed lakes in the Canadian Cordillera are typically dimictic. They have thermal stratification during winter and summer, and experience turnover following break-up in the spring and again in the autumn (Smith, 1978). Interflows occur where the density of influent water is similar to the epilimnion but less than the metalimnion. The flow travels at an intermediate depth, mixing with lake water in a radial jet pattern and rapidly losing energy

2 (Drewry, 1986). Typically, these flow only travel a distance of 10 WX, (Wx = cross-sectional width)(Drewry, 1986; Bates, 1953).

In a similar process, overflows form where the density of inflowing water is less than density of the epilimnion. These tend to form during the spring and autumn when influent water has a low concentration of suspended sediment. Additionally, they can occur where upstream ponds have filtered the sediment load or where stream water has warmed relative to lake water (Drewry, 1986). The turbid flow expands at the surface of the epilimnion in a two-

3 dimensional jet to a distance of 10 WX. Winter overflows were been observed at Lillooet 34

Lake, British Columbia (Gilbert, 1975). These occurred because the density of inflowing water (lower sediment concentration and lower temperature than ambient water) was less than lake water (Gilbert, 1975). At Hector Lake, Alberta, Smith (1978) suggested that epilimnial currents were responsible for transportation of fine sediments, with down-lake katabatic winds driving their circulation. Chapter 3 - Study Area

3.1 Introduction

Kinaskan Lake was chosen as the focus of this study because it is a major sediment trap for material eroded from the Upper Iskut River watershed. It is one of a series of lakes under investigation in north-western British Columbia and the south-western Yukon

Territory relating watershed conditions and regional climate to sediment yield and lake-

specific accumulation chronologies.

3.1.1 Setting

Kinaskan Lake is located in the valley of the Upper Iskut River in north-western

British Columbia. The watershed of Kinaskan, covering 1,257 km2, includes sections of the

Northern and Central Plateaus and Mountains. West of the lake, the Klastine Plateau (1,800

m a.s.l) transitions into Tuktasayda Mountain (2,275 m a.s.l). East of the lake, the Spatsizi

Plateau (2,000 m a.s.l) grades into the Klappan Range. The highest peak in this range and the watershed is Tsatia Mountain (2,362 m a.s.l). Finally, the Tahltan Highland provides a

transitional zone between coastal mountains of the west and plateaus to the east.

A series of tributaries and lakes occupy the Upper Iskut River watershed (Figure 3.1).

Zetu Creek and Summit Creek are conduits for runoff and melt-water in the north-west

section. Zetu Creek feeds Kluachon Lake, whose outflow combines with discharge from

Summit Creek and then enters Eddontenajon Lake. In the north-east region, flow is

transported from Ealue Lake into lower Eddontenajon Lake via Coyote Creek. East of

Kinaskan Lake, Todagin Creek incorporates outflow from Kluea Lake, Todagin Lake and

Tsatia Creek. At Tatogga Lake, discharge from Eddontenajon Lake and Todagin Creek

combine before crossing the Tatogga-Kinaskan delta.

35 36 Upper Iskut River Watershed and Sub-Watersheds, British Columbia Summit Creek

Upper Isk^tRiver Watershed, British Columbia

Coyote Creek

Todagin Creek

Tsatia Creek 57° N s4 130° W"

Figure 3.1 Map of the Upper Iskut River watershed and sub-watersheds [Telegraph Creek ed. 3 (104G), 1989;

Spatsizi River ed. 4 (104H), 1990]. The dashed line represents an elevation of 1,524 m a.s.l., while the dotted line represents sub-watershed boundaries. 37

3.1.2 Morphology

Kinaskan Lake is 17 km long and occupies a narrow valley floor. It covers an area of

27.5 km2 and has an average width of 1.5 km. Although it is oriented in a north-east to south­ west direction, the lake bends into a dog-leg 6 km south of the delta. This leaves a one kilometer section of the lake oriented north-west to south-east.

Shoaling above the dog-leg divides the lake into two basins. The upper basin has a maximum depth of 125 m. At a distance of 3.3 km from the delta, the sill peaks at 58 m below summer lake level. The sill extends across 0.3 km of lake-bottom. Down-lake of the sill, the second basin has a maximum depth of 135 m.

Information contained in geological maps (section 3.1.3), aerial photos (section 5.3), acoustic records (section 5.4) and topographic maps provide limited information about the origins of the dog-leg and sill. However, a report by Souther (1972) suggests that the Iskut

River and Mess Creek valleys may be controlled by major fault zones and folding. Therefore, the author has inferred that these structural characteristics are related to bedrock structure and tectonics.

3.1.3 Geology

Ashe et al. (1995) and Massey et al. (2005) describe the regional geology. According to both sources, the major mountain ranges flanking the Kinaskan Lake valley contain a mix of sedimentary and igneous bedrock. North of Kinaskan, the Cassiar Mountains are of primarily marine sedimentary and volcanic origin, dioritic intrusive rock and a mix of limestone, marble and calcareous sedimentary units. West of the lake, the Spectrum range is dominated by volcaniclastic and alkaline volcanic rocks, marine sedimentary rocks, and a blend of dioritic, syenitic, monzonitic, felsitic and high-level quartz phyric intrusive 38 lithologies of the Spatsizi Formation and the Stuhini Group. East of the lake, the Klappan

Mountains are composed of Bowser Lake Group sediments of the middle to upper Jurassic.

These highly erosive sediments consist of laminated siltstone, mudstone, shale, fine-grained sandstone and chert-pebble conglomerate, with some calc-alkaline volcanics and rare deposits of coal.

The western lowlands, which extend from the base of the Spectrum range to Kinaskan

Lake, are composed of Late Triassic silicious siltstone-mudstone, chert and fine to medium- grained feldspathic wacke, and Lower Jurassic augite-phyric basalt, rhyolite, mafic and felsic epiclastic rocks, and siltstone. Between Todagin Creek and Kinaskan Lake, on the eastern lowlands, bedrock consist of Late Triassic plagioclase, hornblende-phyric volcanics and lithic wacke, along with Lower Jurassic basaltic conglomerate, rhyolite-bearing volcaniclastics, mafic and minor felsic epiclastic rocks and siltstone, and rhyolite (Ashe et al., 1995; Massey et al., 2005; Figure 3.2).

Locally, geological structures in the Upper Iskut River watershed control drainage patterns and introduce intrusive formations. West of Kinaskan Lake are two strike-slip faults of approximately 7 km length each. Along the faults are groats and a pluton of the Early

Jurassic. Just south of the pluton, two shorter faults are separated by an exposed ankerite alteration. North-west of Kinaskan Lake and south-west of Eddontenajen Lake, two faults approaching 4 km length steer a small stream directly into the Kinaskan-Tatogga delta. Two additional faults of similar length are found slightly north. Finally, a dendritic cluster of five faults cradles the Red Stock formation north-west of Kluea Lake (Ashe et al., 1995; Figure

3.3). Upper Iskut Watershed Bedrock Geology 39 —7—jyr1! *-v '—

^^C^

57° N

20 Kilometers 130° W"

Legend Igneous (volcanic) alkaline volcanic rocks Stream calc-alkaline volcanic rocks BUM Walerbody marine sedimentary and volcanic rocks Bedrock Type undivided volcanic rocks Sedimentary volcaniclastic rocks sandstone, sitstone, rare conglomerate I I Igneous (intrusive) mudstone, siltstone, shale tine clastic sedimentary rocks dioritic intrusive rocks conglomerate, coarse clastic sedimentary rocks granite, alkali feldspar granite intrusive rocks coarse clastic sedimentary rocks high level quartz phytic, felsitic intrusive rocks intermixed and varicoloured siltstone, sandstone and conglomerate, minor coal intrusive rocks, undivided laminated siltstone and fine-grained sandstone, chert pebble conglomerate monzodioritic to gabbroic intrusive rocks sandstone, siltstone, conglomerate quartz dioritic intrusive rocks undivided sedimentary rocks quartz monzonitic intrusive rocks limestone, marble, calcareous sedimentary rocks syenitic to monzonitic intrusive rocks

Figure 3,2 Map showing geology of the Upper Iskut River watershed (Massey et al, 2005; Canmap Water,

2006; Geogratis NTS 104G08, 2003; Geogratis NTS 104G09, 2003; Geobase NTS 104G08, n/a; Geobase NTS

104G09,n/a) 40

Earty Pliooant l*par Tim* to lamr Jrante It) MaMsMt CZ3

UMTrimtatt) on 3a nrInnr ito MUdki Annie CXI 10 | lujWnMi> II HTRimVE ROCKB

nnkcttiMf tfltston* ALTERATION (E«r MlMe?)

Figure 3.3 Map showing local geology of the Upper Iskut River watershed (Ashe et al., 1995) 41

3.1.4 Climate Environment Canada's climate stations at Dease Lake (58° 24' N, 130° W) and

Cassiar (59° 17' N, 129° 50' W) provide the longest temperature and precipitation records near the Upper Iskut River watershed (Figure 3.4, Figures 3.5-3.8 and Figures 3.9-3.11, respectively).

^ Provincial Parti A, Ox*** Lata Fate ^rowmaipn

Figure 3.4 Map showing the location of Dease Lake and Cassiar weather stations relative to Kinaskan Lake

(B.C. Parks) 1192340 DEASE LAKE 1OT1-1M0 Normals

Mean Maximum and Minimum Temperatures

| > Muan mEramumtemperaajTe —*— Mean minimum temperature |

25 20

s 10 JP-"" _~ •Nsfc S 5 _^ ,.-***" -^-*. X •^L V >- >-*^ ^~ ^>"^ V "*N. n «-"' -V s. •» -20 v_ -25 •*- JFMAMJ JASOND

Figure 3.5 Monthly temperature normals for Dease Lake (Environment Canada, 1961 - 1990) 1192S4*DEASE LAKE 1961-1M0 Normals

Monthly Precipitation

| 50 I 40

^ lllh II JFMAMJJASOND

Figure 3.6 Monthly precipitation normals for Dease Lake (Environment Canada, 1961-1990)

1102340 DEASE LAKE 1961-1M0 Normals

Monthly Rainfall

70 SO 50 f 40 J 30 20 10 0 1 E JFMAMJJASONllll D

Figure 3.7 Monthly rainfall normals for Dease Lake (Environment Canada, 1961-1990)

1192340 DEASE LAKE 1361-1W0 Normals

Monthly Snowfall

~z jn ji •2- tn I 25 " I • " 15 - I I I • J: 111 1 * lll1 1 1l - _ 1 JFMAMJJASOND

Figure 3.8 Monthly snowfall normals for Dease Lake (Environment Canada, 1961-1990)

1191440 CASSIA*.1461-1M0 Noimah

Monthly PradpttaHon

90 -p 80 - 70-- f 60 - I 50 -- if 40 - - nm. .4+ =tt f 30-- 1 20 - 10 - 0 -iiiiiiiii III JFMAMJJAS O N D Figure 3.9 Monthly precipitation normals for Cassiar (Environment Canada, 1961-1990) 43

1191440 CASSIAR1961-1990 Nonnals

Monthly Rainfall

70 60 f 50 f 40 | 30 20 10 0 t JFMAMJ •JASON D Figure 3.10 Monthly rainfall normals for Cassiar (Environment Canada, 1961-1990)

11*1440 CASSIAR 1961-1M0 Nonnals

Monthly Snowfal

I JFMAMJJASONI D

Figure 3.11 Monthly snowfall normals for Cassiar (Environment Canada, 1961-1990) Normals from both sites show a seasonal pattern of temperature and precipitation. In winter, minimum temperatures are associated with Arctic air while maximum temperatures are related to incursions of Pacific air. Precipitation in the winter months is low and takes the form of snow (Raphael, 2002; Environment Canada, 1961-1990). The transition from winter to summer (i.e. April and May) occurs slowly with temperatures in April remaining low due to high albedo and the persistence of Arctic air. Only 43% of precipitation during this period occurs as rainfall (Raphael, 2002). Summer is characterized by warming and high rainfall due to: (a.) increased solar radiation which drives convection of cooler, unstable air, (b.) the extension of the Hawaiian High northward which replaces cold lows causing storms to shift northward and (c.) cold air pockets which are detached from westerly flow and bring showers

and thunderstorms (Raphael, 2002). Among the -700 mm of total precipitation falling

annually at Cassiar, 275 mm occurs between May and September (Pojar and Meidinger, 44

199). Finally, the return of the Aleutian Low and rejuvenation of westerlies marks the peak of autumn and the slow return to cool winter conditions (Raphael, 2002).

Climate in the Upper Iskut River watershed is a function of many variables including: latitude, air masses, fronts, and mountain effects (Raphael, 2002). Latitude defines the

prevailing winds as well as the solar angle of incidence. Solar radiation is also affected by

elevation, mountain slope aspect, the time of year, the time of day, albedo, atmospheric

conditions, and ambient temperatures. Air masses are separated by three fronts: the Arctic

Front which ushers in cold and dry arctic air (cA), the Pacific Polar Front which presents

cool and wet air (mP), and the Maritime Front which leads warm and tropical air masses

(mT). For six months of the year, Pacific air masses dominate the climate. Storms are

moderated by weakened westerlies and a semi-permanent high pressure zone in the Pacific.

For the remainder of the year, arctic air masses drive circulation with the Rocky Mountains

preventing their westward flow (Raphael, 2002; Pqjar and Meidinger, 1991).

3.1.5 Hydrology

Todagin Creek is the primary source of water and sediment entering Kinaskan Lake.

This creek drains an area of 594 km2 which contains all of the Upper Iskut River watershed's

glacier cover (area = 9.43 km2). These glaciers occupy 1.6% of the Todagin Creek sub-

watershed and 0.75% of the Upper Iskut River watershed. Supplementing inputs from

Todagin Creek are minor amounts of water and sediment from nival melt and autumn

precipitation throughout the Upper Iskut River watershed (Figure 3.12).

Hydrometric data from three stations along the Iskut River provide the timing and

magnitude of discharge peaks. Discharge peaks are related to glacier melt, nival melt and 45 autumn precipitation (Figure 3.13). Downstream from the Upper Iskut River headwaters, the significance of glacier and nival melt relative to autumn precipitation decreases.

The first hydrometric station is located near the outflow of Kinaskan Lake

(08CG003). During the winter, discharge is limited to base-flow. Between May and

September, glacial and nival melt causes discharge to peak at 115 m3/s. Autumn precipitation is responsible for a renewed (albeit smaller) peak in discharge of 50 m3/s (Environment

Canada, 1964-1996).

Downstream from Kinaskan Lake, the Iskut River is fed by Snippaker Creek

(08CG004). Glacial and nival-melt creates up to 1,500 m3/s of discharge between May and

August. Highest flows occur later in the year when autumn rains bring total discharges to

1,000 - 4,750 m3/s (Environment Canada, 1966-1995).

The third hydrological station is on the Iskut River below Johnson River (08CG001).

From January to April, base-flow discharge is 250 m3/s. Between May and September,

glacial and nival melt introduces discharges of 2,000 m3/s. As sources of meltwater are

depleted, autumn precipitation leads to peak discharges of up to 7,000 m3/s (Environment

Canada, 1959-2006). 46

Upper Iskut Watershed Hydrology Legend Stream ; Wateibody i Permanent Snow Cover

SummitCreek

Zetu Creek" •

Coyote Creek

^Tpdagfii CreBfc-v

V

£f»

Tsatia Creek _• '*<--.-. '

>^?

57° N EaG>a*ii ^^ c=» JL -I—i—i—i—i J S 16 Kilometers 130° W

Figure 3.12 Map showing the drainage network and glacier cover of the Upper Iskut River watershed (Canmap

Water, 2006; Geogratis NTS 104G08, 2003; Geogratis NTS 104G09, 2003; Geobase NTS 104G08, n/a;

Geobase NTS 104G09, n/a) 47

Figure 3.13 Hydrological Regimes of sites along the Iskut River (Environment Canada, 1966-1995;

Environment Canada, 1959-2006; Environment Canada, 1964-1996). The red line represents data from a given year, the green line provides the maximum daily discharge, the blue line provides the minimum daily discharge and the yellow line provides the mean daily discharge. 48

3.1.6 Vegetation

The Upper Iskut River watershed is located within the Boreal Cordillera ecozone. The ecozone is over 50% forested with open to closed-canopy cover. In the north and eastern portion of the watershed, several types of spruce, willow and birch dominate. In the western region, Engelmann spruce, sub-alpine fir and sub-boreal spruce appear. Tree-line is at approximately 1,200 m and the watershed straddles two permafrost zones: Isolated Patches

(0 - 10%) and Sporadic Discontinuous (50 - 90%)(CCMF, 2002; Cassiar Forest, 2003;

Permafrost). Chapter 4 - Methods 4.1 Field Methods

A number of field approaches were taken to characterize the sedimentary

environment of Kinaskan Lake. These include geophysical measurements, coring and water

column profiling.

4.1.1 Geophysical Methods

Sub-bottom acoustic profiles were collected during the summer of 2004. A

Datasonics Dual-Frequency Chirp II system was used to collect sub-bottom sediment

characteristics along 50 km of survey line. The sub-bottom imagery has a maximum

resolution of 0.5 meters and under ideal conditions can penetrate up to 100 meters into the

silt and clay muds. Survey lines were conducted on oblique cross-lake transects and one long

profile ran down the center of the lake.

4.1.2 Sediment Sampling

During the 2006 field season, sediment coring sites were selected and a Rossfelder

submersible vibra-coring system was used to collect up to 6 m long cores. For each sediment

core, a six-meter-long aluminum pipe, with an opening diameter of 75 mm, was fitted with a

core catcher on one end. The opposite end was inserted into the vibra-core head and the unit was lowered to the lake bottom. At the campsite, cores were cut into shorter segments for

shipping. Four short Ekman cores were also taken in attempts to capture undisturbed surface layers. These were also stored in short pieces of aluminum pipe (Table 4.1 a).

49 50

Core Core Number Distance from Water Wet Sample Type (Waypoint) Site Coordinates Delta (km) Depth (m) Length (m) 57° 40' 19.73" N, 130° 6' Vibra 310 30.518" W 1.5 79 4.268 57° 40' 23.624" N, 130° 6' Ekman E4-06 29.22" W 1.7 78.4 0.1095 57° 38' 10.469" N, 130° T Ekman E3-06 13.086" W 6 109.3 0.0895 57° 35' 26.685" N, 130° 8' Ekman E2-06 42.375" W 11.5 130.5 0.086 57° 35' 32.655" N, 130° 8' Vibra 309 48.085" W 11.52 121.5 4.193 57° 34' 52.423" N, 130° 9' Vibra 311 18.195" W 12.8 130.9 4.606 57° 34' 53.461" N, 130° 9* Ekman El-06 7.553" W 12.9 132.3 0.081 r

Table 4.1 a Site characteristics and sample lengths for 2006 cores. An additional nine Ekman cores, which were taken during the field season of 2004, were processed (Table 4.1 b).

Core Core Number Distance from Water Wet Sample Type (Waypoint) Site Coordinates Delta (km) Depth (m) Length (m) 57° 40' 3.5" N, 130° Ekman 237 & 54.4" W 2.29 -80 0.085 57° 38' 55" N, 130° 80:Sx< Ekman 239 T 24" W 4.48 120 0.08 57° 38" 11" N, 130° Ekman 240 7- 14» w 5.84 -110 0.085 57° 36'54" N, 130° Ekman 242 7 11" W 8.29 >120 0.095 57° 35' 45" N, 130° Ekman 258 8' 30" W 10.72 >120 0.07 57° 34' 55" N, 130° Ekman 247. 9' 12" W 12.45 >120 0.0975 57° 34' 2" N, 130° 80 £x< Firman 250 10' 15" W 14.34 120 0.085 57033.22" N, 130° 80Sx< Ekman 252 10' 53" W 15.79 120 0.11 57° 32' 52" N, 130° Ekman 253 11' 13" W 16.66 40£x<80 0.095

Table 4.1 b Site characteristics and sample lengths for 2004 cores. 51

4.1.3 Water Column Profiling

Field measurements of temperature, turbidity and conductivity, were taken with a

Datasonde 4A Hydrolab in late July of 2004 and late June of 2006.

4.2 Laboratory Methods

Analysis of samples involved measurement of lamina and bed thickness, radiocarbon dating, Loss-on-ignition estimates for organic content, grain size analysis and X-Ray

Fluorescence for geochemical make-up.

4.2.1 Core Preparation

In the laboratory, core sections and Ekmans were stored upright for a week to promote segregation of interstitial water. At the end of the week, the upper plastic cap was removed and a siphon was used to remove any surface water.

The cores were split by incising the outer casing and using steel wire. A series of digital photographs was taken at various stages of drying to record lamina and bed structures

(Lamoureux, 1996). One half of the core was sealed for archival purposes, while the working half was left to dry in a controlled environment.

4.2.2 Lamina or Bed Thickness Measurements and Dating

When the 2006 sediment samples achieved a plastic-like texture, typically after two to three weeks of drying (due to the high clay content of the cores), the surface was scraped clean using a sharp razor blade. At this stage of drying, lamina and beds were easily distinguished by their differing color. Lamina and bed thicknesses were measured using an

Increment Measurement System which included a microscope with 40 x magnification and length measurements with precision of up to 0.01 mm. Similar tests were not conducted on the 2004 Ekman samples. 52

Couplet (pairs of light and dark lamina) counts were made to establish a preliminary chronology of sediment deposition. Additionally, four samples of organic matter were sent to the Isotrace Laboratory at the University of Toronto for radiocarbon dating.

4.2.3 Loss-on-ignition

The content of organic and carbonate matter in the vibra-cores was estimated through loss-on-ignition (LOI) at five centimeter intervals along the length of the cores. For the 2006

Ekman cores, three samples were taken; one at the surface, one at 3 cm depth and one at 6 cm depth.

Preparation involved oven-drying the samples over 12-24 hours at 105° C to remove excess moisture. The dried samples were homogenized using a mortar and pestle, and placed in pre-weighed crucibles. The difference between the weight of the sample plus crucible, and the crucible, gave the sample weight, DW105 (Heiri et al., 1999). For organic content, the samples were placed in a 550 °C furnace for one hour and then were immediately transferred to a dessicator for cooling. The samples were reweighed. The difference between the original sample weight, DW105, and the new sample weight, DW550, is an index of organic matter in the sediment.

For carbonate content, the samples were returned to the furnace and heated at 1100

°C for one hour. They were immediately removed and returned to the dessicator for cooling.

The samples were reweighed giving DWnoo- The difference between DW550 and DWnoo is an estimate of carbonate content. Since the Upper Iskut River watershed contains minimal carbonate bedrock, an additional test was conducted to determine the validity of the DWnoo values. Several drops of 10% HC1 were added to ten dried, homogenized samples of sediment. As expected, there was no visible reaction. Therefore, it is probable that the 53

anomalously high DWnoo values are due to absorption of humidity prior to weighing and

incomplete combustion of organic matter during the first heating event (Heiri et al., 1999).

A similar protocol was followed during processing of the 2004 Ekmans. However,

only two samples at the surface were used per core.

4.2.4 Grain Size Analysis

For the three vibra-cores, laser particle sizing was conducted using instruments from

two different laboratories. Twenty samples were taken per core, corresponding to

representative sediment structures. To eliminate organic content, samples were treated with

several drops of hydrogen peroxide and left for one week. Afterwards, samples were

homogenized using a mortar and pestle.

Samples taken from KIN 309, a distal core, were processed at Queen's University

using a Coulter Laser Sizer. Samples taken from KIN 311, the second distal core, and KIN

310, the proximal core, were processed at the University of Toronto using a Malvern Laser

Sizer. For the 2004 Ekmans, samples were processed in a Sedigraph.

Laser Sizers function on the standard that the size of a particle is inversely related to

the angle at which they diffract light. The sample is mixed with distilled water and ultrasonically disaggregated. Over a period of ten minutes, a beam of monochromatic light passes through the solution and a multi-element ring detector senses the angular distribution of scattered light (Singer et al., 1988). The proportion of sample within a number of grain

size categories is provided.

To compare the grain size results from the Coulter and Malvern Laser Sizers, a

Malvern test was run for three of the twenty samples analyzed on the Coulter system. These 54

results showed that there was not a significant amount of variation between the results

generated by the different instruments (Appendix B).

4.2.5 X-Ray Fluorescence [XRF]

Five samples of sediment from KIN 311 were analyzed using the X-Ray Fluorescence

Spectrometry Laboratory in the Department of Geology at the University of Toronto.

Samples exceeding 3.5 g were dried over night at 105° C before being homogenized with a

mortar and pestle (Table 4.2). At the XRF Laboratory, they were compressed into pellets and

then run through a Phillips 2404 Sequential XRF where they were bombarded with X-rays

from a high power X-ray tube. A semi-quantitative scan was used to measure the elements in

the sample.

Core Number Depth in Core Sample Mass (Waypoint) (cm) (S) Percent per Grain Size Class 311 15 4.821 70% Clay, 28.3% Silt, 1.7% Sand 311 77 6.76 25% Clay, 65% Silt; 20% Sand 311 162 5.077 60% Clay, 48% Silt, 2% Sand 311 285 5.325 50% Clay, 47% Silt, 3% Sand 311 362 5.389 63% Clay, 37% Silt, 0% Sand

Table 4.2 Characteristics of samples processed for XRF Chapter 5 - Results

5.1 Introduction

Past and modem processes define the distribution and composition of sediment across

Kinaskan Lake. Results here are derived from CTD measurements, aerial images, acoustic profiles, sedimentary structures, sediment composition and chronological measurements which are discussed in detail to highlight these processes.

5.2 Circulation of Kinaskan Lake

Two series of CTD measurements were collected for Kinaskan Lake. A series of CTD measurements taken to represent early season conditions was made on June 26, 2006, at sites

1.6, 5.1, 9.2 and 12.5 km fromth e delta (Figure 5.1).

Temperature (°C) Turbidity

34 56789 10 11 0 10 20 30 40 50 240 250 260 270 0 0 [, • \ \ 10 10 ""~~> \ 20- 20 (X V: K 30 30 • ts/•-" • > (a) (c) 40 40 - 1 50 „ 50 g 50 - \5

60 o 60 Q Q

70 70 • \

- 12,5 km 80 SO - - 9.2 km 5.1 km 1.6 km 90 90- -90

100 100 J - 100

June 26.2006 June 26.2006 110 110 -

T—>—I—'—I—•—I—*-T—P~1—>-T 3456709 10 1 10 20 30 40 50 240 250 260 270

Figure 5.1 CTD profiles of Kinaskan Lake from June 26, 2006. Turbidity was measured in NTU units and converted to mg/L using the calibration developed by Hodder et al. (2006) for this instrument.

55 56

These profiles show a distinct down-lake pattern of temperature, turbidity and conductivity. The proximal region appears to have active development of thermal structure, chemical structure and a turbid interflow. With distance away from the delta, thermal structure weakens; turbidity is limited to weak overflows and conductivity increases but is more uniform with depth.

The weak pattern of thermal stratification is related to lake circulation following turnover. The breakup of ice-cover at the end of May or start of June and the influx of melt- water disrupts the weak thermal stratification and facilitates wind-induced turnover. Because land is a better conductor of heat than water, the littoral and pro-delta region develops a warm cap of water (9-10°C). Therefore, proximal lake waters have a thicker epilimnion, a

steeply graded thermocline and a hypolimnion, while the other three sites have less developed stratification.

Turbidity follows a similar down-lake pattern. The epilimnion of proximal waters is higher in turbidity (50-30 mg/L), with an interflow between 20 and 25 m water depth. At

distal sites, turbidity is limited to weaker (25-10 mg/L) overflows.

Conductivity, a measure of dissolved solids in water, is lowest in the epilimnion of

proximal waters (250 uS/cm). Below the epilimnion, a chemocline marks the depth of

maximum rate of increase of conductivity (24 m; 251-263 uS/cm) and the hypolimnion

maintains a value of 266 uS/cm. Away from the delta, the overall conductivity increases to a

maximum of 270 uS/cm. However, chemical stratification disappears.

The second series of CTD measurements was taken on July 26, 2004 at eight different

down-lake sites (Figure 5.2). Warming air temperatures, greater wind and wave circulation 57 and increased precipitation produced lake-wide thermal stratification, a series of down-lake turbid interflows and a down-lake conductivity gradient with a distinct chemocline.

Temperature (°C) Turbidity (mg/L) Conductivity (uS/cm)

0 5 10 15 20 5 10 15 20 25 30 230 240 250 280 270 280 0 H 0 • (a) I v|| " (b) 10 - *Ji 10 • IjUTs. //pY^^2? 20 • P^ : 20 - K*^~" 30 • 1 30 • A 40 I 40- - 50 • 1 50- 1 - . S. 80 • \ % *° sz ' 7 5 °- • S 70- 80 - - 80- -

90 • - 90- -

100- - 100 - -

110 - - 110 - -

July 28, 2004 July 28.2004 120 - 120 • ' 130 - 130- —i—i—i—i—i—i—i—i—p— 5 10 15 20 5 10 15 20 25 30 230 240 250 260 270 280 Temperature Temperature Temperature

1.2 km 2.3 km 3.6 km 4.8 km 6.2 km 9.3 km 12.5 km 14.2 km

Figure 5.2 CTD profiles of Kinaskan Lake from July 26,2004 58

Down the lake, a similar pattern of thermal stratification is evident. By late July, the epilimnion is better developed throughout the lake with slight thinning down-lake (15 m in the pro-delta region to 11 m thick at 14.2 km from the delta). Greater activity of wind-waves in the distal lake could be responsible for reduced warming of surface water.

Turbidity, taking the form of interflows and one overflow, decreases in concentration away from the delta. The most proximal sites have interflows at 15-18 m depth, with turbidity in the range of 25-20 mg/L. The weakened turbidity concentration down-lake could be related to progressive settling from suspension with distance from the primary source.

There is an anomalously higher surface concentration of sediment at 12.5 km down-lake

(22.5 mg/L) whose source is unknown.

Finally, strong chemical stratification down-lake is shown in the conductivity profiles. Lake surface conductivity is at its lowest value along the water surface (245 uS/cm).

This vertical structure is constant down the lake. The pattern could be related to dilution of surface dissolved solids by greater precipitation in July than June, or the tendency of inflowing water and solutes to travel as an interflow rather than an overflow. With distance from the delta, this zone of low-conductivity thins. Under the lake surface, the zone is truncated by a chemocline which extends from 15 to 20 m depth in the pro-delta region and from 11 to 19 m depth at the most distal site. Below the chemocline, the entire lake shares a high conductivity of 264-265 uS/cm.

The conductivity in Kinaskan Lake is considerably higher than has been reported at

Lake Superior (97 uS/cm) but is less than Lake Independence, Michigan (316 uS/cm) and

Lake O'Hara, British Columbia [100 - 170 uS/cm (2004) and 90 - 190 uS/cm (2005)](Water on the web, 2004; Hood et al., 2006). Through a study of 76 Swedish lakes, Nilsson and 59

Hakanson (1992) attributed differences in conductivity (Mean Conductivity = 41 uS/cm; Min

= 16 uS/cm; Max =108 uS/cm) to watershed characteristics; the amount of open land, discharge and relief.

5.3 Attributes of the Upper Iskut River Watershed Inferred from Aerial Imagery

No excursions were made into the inaccessible upper watershed, therefore interpretation of surface conditions from aerial photographs taken in 1967 have been used to assess possible sediment sources. It is probable that most of the sediment delivered to

Kinaskan Lake originates from the Todagin Creek - Tsatia Creek sub-watershed. This sub- watershed is underlain by the highly erosive Bowser Lake Group (middle to upper Jurassic)- an amalgamation of siltstone, sandstone and conglomerate with minor amounts of limestone and rare coal. Outside of this southeast portion of the Upper Iskut River watershed, minor contributions of sediment come from the Spatsizi Formation and the Stuhini Group in the northwest.

A combination of colluvial action, alluvial processes, periglacial activity and glacial processes are responsible for mobilizing sediment throughout the watershed but especially in the Todagin Creek - Tsatia Creek sub-watershed. Mass-wasting processes range from rock- falls and rock-slides to debris flows. Evidence of these processes includes deeply dissected bedrock chutes (Figures 5.3 a & b).

Glacial melt and runoff provide sufficient discharge to maintain complex networks of channels. Steep debris and alluvial fans have formed at the confluence of minor streams and trunk-rivers (e.g. Tsatia Creek and Todagin Creek). Sediment from these fans has been subsequently eroded, along with (possibly late Pleistocene) valley fill sediments. Irregular erosion of the valley fill has caused channels to develop a terraced form in limited reaches of 60 tributaries. Most of the rivers appear to have at least some braiding, signifying their large sediment loads contributed from both glacial and non-glacial sources (Figures 5.3 a & b).

Periglacial processes, including accelerated weathering of bedrock and mobilization through creep or frost-heave,ar e active in the high plateaus of the Todagin Creek - Tsatia

Creek sub-watershed. Evidence of these processes is difficult to identify in aerial images, but their contributions of sediment to Kinaskan Lake should not be overlooked.

Finally, the importance of glacial activity has varied during the Pleistocene and

Holocene. However, evidence of past and present glacial activity is omnipresent in the entire

Upper Iskut River watershed. A series of cirques, in-filled with till, moraines and rock glaciers, dominates the highest elevations. Meanwhile, late Pleistocene till mantles hill- slopes (Figures 5.3 a & b).

Figures 5.3 a & b (a) Aerial photo A20117-91 of Tsatia Creek and (b) Aerial Photo A20059-9 of Todagin Creek. 61

5.4 Acoustic Record

Sub-bottom acoustic survey lines which follow oblique across-lake transects and a long profile, were collected during the field season of 2004. The information contained within these images was used to create a bathymetric map and a map of sediment thickness for Kinaskan Lake.

5.4.1 Bathymetry, Sediment Thickness and Character

Lake-bottom depths from the acoustic imagery were used to map Kinaskan Lake's bathymetry (Figure 5.4; Appendix A). At the head of the lake, depth grades from less than twenty meters to eighty meters. Approximately 4 km down-lake from the delta, the lake bottom shoals to forty meters depth. Gilbert et al. (2006) identified similar sills in Llewellyn

Inlet and north of Copper Island in Atlin Lake. They attributed the contrast between greater proximal sedimentation and lower distal accumulation to the presence of these divides.

Down-lake from this sill, the lake deepens to 125 meters. The lake bottom only gradually rises over the final 4 km to the outlet. 62

Kinaskan Lake Bathymetry

57° N

130° W"

[Figure 5.4 Map showing the bathymetry of Kinaskan Lake, core sites and location of selected acoustic profiles shown in figures 5.5 to 5.7 (Canmap Water, 2006; Geogratis NTS 104G08, 2003; Geogratis NTS 104G09,

2003; Geobase NTS 104G08, n/a; Geobase NTS 104G09, n/a) 63

Sediments of Kinaskan Lake appear to be mostly fine (silty-clay) with isolated parabolic reflectors suggesting localized accumulations of coarse materials. Two patterns of sedimentation dominate the lake. First, the thickest accumulations of sediment are concentrated at four sites (see thickness contours in Figure 5.8). The first site is located ~1 km south of the Kinaskan Lake delta and covers -0.5 km of distance down-lake. In acoustic images, sediment reaches a maximum thickness of 30 m. It is likely that these thick deposits extend into the delta, however, we could not acquire acoustic imagery further north due to shoaling and sediment coarsening. In both the long profile and the cross-lake transects of this region, sediment is well-layered with multiple reflectors (Figure 5.5). Sediment is probably deposited from turbidity currents which spill off of the delta. The thickest deposits are conformable to raised hummocks on the periphery of the lake bed where energy is lower. In the deepest part of this site, a trench is in-filled with ten to fifteen meters of well-layered and horizontally-bedded sediment. Edges of these reflectors are not conformable with the trench walls suggesting infilling via higher energy turbidity currents. Gilbert et al (1997) noted similar structures in the sediments of Bowser Lake. Assuming that deposition began after local deglaciation (-11,500 BP) (Spooner et al., 2002), the thickest deposits of 30 m would require an average sedimentation rate of 2.6 mm/year (Spooner et al., 2002). 64

Figure 5.5 Acoustic image taken across-lake representing one of the thick accumulations of sediment in Site 1. The location is shown on Figures 5.4 and 5.8.

Between the first and second site of major sediment accumulations, sediment samples

KIN 310 and E4-06 were taken. Site 2 is located 2.5 km south of the sill and also covers -0.5 km of distance down-lake. Acoustic images show accumulations of sediment up to 40 m thick (Figure 5.8). From this site, Ekman 237 was collected. For 40 m of sediment to accumulate since local deglaciation, an average sedimentation rate of 3.48 mm/year would be necessary.

The third site begins at the distal end of the sill, -6.25 km from the Kinaskan Lake -

Tatogga Lake delta. Deposits appear well-layered and have a maximum thickness of 40 meters (Figure 5.6). They are thinner on the slopes and thicken towards the center of the lake.

Sediments are conformable to raised hummocks. In in-filled trenches, the deepest sediments are unconformable and are capped by conformable deposits. These deepest sediments were probably laid from high-energy under-flows. The conformable, well-layered sediments were probably deposited from sediments settling from suspension, suggesting a shift from under­ flows to inter-flows and over-flows as a dominant sedimentation process. Both E3-06 and

Ekman 240 were taken from this site. Assuming that deposition began after deglaciation, an 65 average accumulation rate of 3.48 mm/year would be necessary to generate 40 m of sediment

(Figure 5.8).

xj •w • 0[ B' • OS •0E

HOC • V

v <"* • » •jbw "w • N

^SB^^^HL •"

^^gnl

Figure 5.6 Acoustic image taken across-lake representing thick accumulations of sediment at Site 3. Here, both conformable sediments perched on a hummock, and unconformable sediments occupying a trench are evident. Location of image is shown in Figures 5.4 and 5.8.

The fourth site begins 12 km from the mouth of Kinaskan Lake. It spans a distance of

4.5 km down-lake and features a smooth and deep lake bottom (125 m). Thick accumulations of sediment are mostly confined to the deep center of the lake. Sediment is laminated and has weak, conformable acoustic reflectors. The sediment thins on the adjacent slopes (Figure 5.7;

Figure 5.8). Samples E2-06, KIN 309 and Ekman 258 were taken from the central part of this site. Meanwhile, samples El-06, KIN 311 and Ekman 247 were taken from the distal end of this site. Along the periphery, sediment thickness reaches 40 m on several occasions. This suggests an average accumulation rate of 3.48 mm/year since local deglaciation. 66

c C

i

Figure 5.7 Acoustic image taken across-lake representing the fourth site. The lake bottom is smooth and flat, and composed of conformable sediments which cap unconformable deposits. The unconformable deposits were probably deposited from underflows. The location of this image is evident in Figures 5.4 and 5.8.

All of these four sites appear to have been influenced either early on or continuously by turbidity currents. Areas with lower rates of sedimentation are fed by lower-energy interflows and overflows (Smith, 1981). Sedimentation on the western slope (right side) of the lake does not appear to be influenced by the Coriolis effect. 67 Kinaskan Lake Sediment Thickness

11111111 0 0.5 ! 2 Kilometers 130" W Figure 5.8 Map showing the spatial distribution of sediment thicknesses and the locations of transects shown in

Figures 5.5 to 5.7 (Canmap Water, 2006; Geogratis NTS 104G08, 2003; Geogratis NTS 104G09, 2003;

Geobase NTS 104G08, n/a; Geobase NTS 104G09, n/a). The dots represent spot depths and were used to develop the contours. 68

5.5 Sedimentology of Kinaskan Lake

The short (Ekman) and long (vibra) cores are used to characterize the sedimentology of Kinaskan Lake. Overall, the structure and composition of the sediment appears to have distinctive spatial and temporal patterns.

5.5.1 Sediment Structure

Most of the sediments in the three vibra-cores and the four Ekmans from 2006 are arranged in well-layered couplets. Each couplet consists of a coarse-grained, light-colored layer which is overlain by a fine-grained dark-colored layer. The contact between light and dark layers ranges from diffuse to abrupt. Periodically, couplets are interrupted by deposits of massive silts and disturbed beds. Among the vibra-cores, there is a weak trend of decreased

average couplet thickness with distance down-lake (from 4.2 mm at 1.5 km from the delta to

2.05 mm at 12.8 km from the delta). This is probably due to clastic-laden underflows which

cross the delta, losing momentum and creating deposits in the pro-delta region. Some variance is probably related to inputs from lake-shore tributaries and colluvial action (Table

5.1; Figure 5.9; Appendix C). The Ekman sediments show a similar couplet structure and

down-stream thinning to the vibra-cores. However, they were not compressed during

sampling (as were the vibra-cores) and their record is brief. Therefore, Ekman couplet

thicknesses should not be compared with thicknesses for the vibra-cores. 69

Core Mean couplet Type Core Number (Waypoint) Distance from Deha (km) thickness (mm) lo(mm) Vibra 310 1.5 4.42 3.37 Ekman E4-06 1.7 6.08 3.29 Ekman E3-06 6 5.26 2.36 Ekman E2-06 11.5 4.09 2.26 Vibra 309 11.52 2.85 2.04 Vibra 311 12.8 2.05 1.2 Ekman El-06 12.9 5.06 2.68 Table 5.1 Mean couplet thickness of sediments

• Vibra-cores • Ekman cores -Linear (Ekman cores) -Linear (Vibra-cores)

0 2 4 6 8 10 12 14 Distance (km) Figure 5.9 Mean couplet thicknesses relative to distance from the Kinaskan Lake - Tatogga Lake delta. Thinner couplets in vibra-cores are probably due to compression during sediment sampling.

Among the three vibra-cores, couplet thicknesses appear to have an inverse relationship with distance from the delta. Sediments from the proximal vibra-core site, KIN

310, have a mean couplet thickness of 4.42 mm (±3.37 mm). The two distal vibra-cores have a thinner average couplet thickness: KIN 309 = 2.85 mm (± 2.04 mm) and KIN 311 = 2.05 mm(± 1.2 mm). 70

A similar down-lake pattern was found for couplet thicknesses in the Ekman cores.

Couplets from proximal cores (E4-06 and E3-06) had an average thickness of 6.08 mm (±

3.29 mm) and 5.26 mm (± 2.36 mm). Down-lake, couplets thinned to 4.09 mm (± 2.26 mm) in E2-06 and 5.06 mm (± 2.68 mm) in El-06.

In addition to this weak relationship between the distance of the coring site from the delta and couplet thickness, thicknesses change throughout the core. This is particularly

significant for the three vibra-cores (Figures 5.10 a, b & c).

I I" •n

8 8 8 8 8 8 in •t (UIUI)OJOD LII ipdsfi Figure 5.10 a (a) Kin 309 couplet thickness versus depth in core. The base of the core contains massive silt

deposits and a series of anomalously thick couplets. 71

(turn) ojoa ui i[itlof | Figure 5.10 b (b) Kin 310 couplet thickness versus depth in core. Clastic-rich inflows crossing the delta leave these thick couplets.

(uiuOojODUUfldafl

Figure 5.10 c (c) KIN 311 Couplet Thickness versus depth in core. Couplet thicknesses are more variable than

those from KIN 309, but less variable than those from KIN 310. 72

KIN 310 is composed of couplets whose thicknesses range fromver y thick (29.8 mm) to very thin (0.4 mm). There do not appear to be any significant shifts in the long-term mean couplet thickness with depth (time), but rather a series of fluctuationsabou t the mean. These are probably related to discharges that are highly variable from year to year and between major climate periods (Figure 5.11; Figure 5.10 b).

Figure 5.11 The couplet (rhythmic) structure of sediments in KIN 310 Approximately 11 km down-lake from KIN 310, KIN 311 has a narrower range of values about a thinner mean. Couplet thickness rarely exceeds 10 mm (mean = 2.05 mm ±

1.2 mm), and thicker couplets appear to be clustered in the lower half of the core (Figure

5.12; Figure 5.10 c).

Figure 5.12 The couplet (rhythmic) structure of sediments in KIN 311 73

A similar deposition pattern to KIN 311, albeit more extreme, is found in KIN 309.

The mean couplet thickness is thinner (2 mm +/- 2.9 mm), with limited variability in the upper 2 meters of the core. Below this point, thin couplets are replaced by a series of extremely thick beds (40 mm thick), thin clay lamina, and a section of massive silt (Figures

5.13 a&b; Figure 5.10 a).

Figures 5.13 a & b Sedimentary structures found in KIN 309; (a) finely laminated couplets and (b) massive deposits of clayey-silt.

When smoothed with a 25 year filter moving average and plotted against couplet count, parallel peaks and troughs in couplet thicknesses of KIN 309 and KIN 311 (excluding

KIN 310 due to high sensitivity to seasonal or isolated discharge peaks and turbidity currents) are apparent. (Figure 5.10 d) 74

8 9 10 -i i i_—i t_

250 250 core 309 vs couplet # core 311 vs couplet # core 310 vs couplet #

7S0

1000 h 1000

f 1250 1250

1750

2000 2000

2250-^

2500 —i—>—i—•—i—i—i—•—i—i—i—i—i—•—i—<—i—•—i—) 2500 01 23456789 10 Couplet Thickness (mm)

Figure 5.10 d KIN 309, KIN 311 and KIN 310 Couplet thickness versus couplet count (smoothed with a 25 year filtermovin g average).

Relatively high values among the smoothed data (peaks) can signify a long-term elevation in temperature and/or precipitation. Deficits (troughs) suggest a long-term reduction in either or both variables. Disregarding thickness values below couplet number

1500, due to an apparent disturbance in KIN 309, a common pattern of peaks and troughs between the two distal cores is revealed. Between couplet 1500 and couplet 1300, deposits of 75 below average thicknesses occur (-1.5 mm). From couplet 1300 to couplet 1200, couplet thickness rises to 2.5 mm. Minor fluctuations in couplet thickness (-1.75 mm to -2.25 mm)

are shown in both cores from couplet 1200 to couplet 1000 with KIN 309 values lagging

slightly behind KIN 311. Between couplet 1000 and couplet 750, both cores show a large

deficit in couplet thickness (-1.5 mm) followed by a dramatic increase (-3 mm). Here, KIN

309 values appear to lead those of KIN 311. From couplet 750 to couplet 650, thicknesses

decrease slightly. From this depth until couplet 450, thicknesses are highly variable with a

slight overall increase. Finally, the couplet thickness pattern of the upper portion of these two

cores is reversed. In KIN 309, overall couplet thickness decreases to -1 mm by couplet 250.

Thickness slowly increases to -2.25 mm by couplet 100, from where values being to

decrease. In KIN 311, from couplet 450 to couplet 200, thickness rises to -3.25 mm. From

couplet 200 to couplet 50, thicknesses fluctuate about 3 mm. After couplet 50, couplet

thickness decreases to 2.25 mm and then slowly returns to 3 mm by the top of the core.

5.5.2 Organic Content

To estimate the organic content of sediment in the three vibra-cores, loss-on-ignition

was run on specimens that were taken at five centimeter intervals. For the four Ekman cores,

samples were taken at the surface, 3 cm and 6 cm depths. This sampling interval differs from

the protocol followed by a previous study of loss-on-ignition for the 2004 Ekman cores. In

2004, two samples were taken from the same depth in each core. Following 1 hr of heating at

550 °C, all samples lost a range of 6.6 to 10% of their mass (Table 5.2; Figure 5.14). 76

Mean 1

These values are low compared to those found at nearby Skinny Lake (a wetland lake surrounded by peats and other organics) where loss-on-ignition results ranged from < 10% to

60% (Spooner et al., 2002). However, they are high compared to those from glacier-fed Mud

Lake (< 1% to -5%) and they fall within the same range as those for Pyramid Lake, 1 to 9%

(Hodder et al., 2006; Mazzucchi et al., 2003). 77

• Vibra-cores • Ekman cores -Linear (Ekman cores) -Linear (Vibra-cores)

0 2 4 6 8 10 12 14 16 18 Distance (km) Figure 5.14 Mean Loss-on-ignition relative to distance from the Kinaskan Lake - Tatogga Lake delta.

Loss-on-ignition values do not have a strong down-lake pattern. However, the average LOI value from the proximal vibra-core (KIN 310) is lower than the average values from the two distal vibra-cores (KIN 309 and KIN 311). Lower values from the proximal cores are related to high clastic inputs and circulation patterns which maintain fine organic matter in suspension until it reaches calmer distal waters. Alternatively, distal sites receive greater water column contributions of organic matter and organic inputs from small local tributaries. Clastic inputs are also lower here. A similarly weak down-lake pattern was found for Ekman cores.

Among the three long cores, fluctuations in loss-on-ignition with depth suggest a temporal pattern of organic contributions. In the two distal cores (KIN 309 and KIN 311), organic content is high and variable from the base of the core to the 1.8 m and 2.4 m mark, 78 respectively. At this threshold, there is a sudden drop in organic content. In JON 309, the mean organic content below 1.8 m is 7.7% (± 0.88%). This corresponds to couplet (and approximate years BP) 1053. At the threshold, organic content falls to 6.2%. Above 1.8 m, the value stabilizes around a 6.5% mean with a much lower standard deviation of 0.2%

(Figure 5.15 a). In KIN 311, the mean organic content below 2.4 m is 7.1% (± 0.9%). This corresponds to couplet (and approximate years BP) 1123. At this depth, organic content suddenly drops to 6.2%. Above this point, organic content is highly variable (± 0.9%) but increases steadily to the present value of -8.5%. Above and below the threshold, the overall means are similar (above 2.4 m, Mean LOI = -7.2%; below 2.4 m, Mean LOI =

~7.1%)(Figure 5.15 c). In the proximal core, KIN 310, three different means are revealed.

The first mean occurs between the base of the core and the 3 m mark. Here, values fluctuate about -7%. Between 3 m and 2 m, organic content appears to decrease and there is a new mean of-6.4%. Above 2 m, there is a steady increase in organic content until a new mean of

-8% is established between 50 cm and the top of the core (Figure 5.15 b). 79

u

Cm9)ojODU|mdo(;i

Figure 5.15 a (a) KIN 309 Loss-on-ignition results versus depth in core (Losses from ignition at 1100 °C have not been shown as the watershed contains minimal carbonate).

(lM3)0JO3 III L[ld8(|

Figure 5.15 b (b) KIN 310 Loss-on-ignition results versus depth in core (Losses from ignition at 1100 °C have not been shown as the watershed contains minimal carbonate). 80

© vrt s

8 *

s s s § s *- (N r»l ^fr *n (ma) OJOO ui ipdaf | Figure 5.15 c (c) KIN 311 Loss-on-ignition results versus depth in core (Losses from ignition at 1100 °C have not been shown as the watershed contains minimal carbonate).

These values can also be presented as a standardized sedimentation index [S = (mean organic content of the core)/(organic content of the sample)]. This index, presented by Souch

(1994) in a study of clastic and organic sediment influx into glacial-fed lake systems, suggests that values greater than 1 indicate an above average rate of clastic sedimentation, values equal to 1 indicates average values, and values below 1 represent a below average rate

(Figures 5.16 a, b & c) 81

n %

n

(UIO)OJOO ui ifldafl

Figure 5.16 a (a) KIN 309 Loss-on-ignition results presented as a standardized sedimentation index versus depth in core

(UI9) 8103 III L{)dS(|

Figure 5.16 b (b) KIN 310 Loss-on-ignition results presented as a standardized sedimentation index versus depth in core 82

<~i •i > 8 •a .a

"g 6

<1 o

(llio) 8JO0 III LfldOfl

Figure 5.16 c (c) KIN 311 Loss-on-ignition results presented as a standardized sedimentation index versus depth in core

Unlike the long vibra-cores, the temporal record of organic matter from the Ekman cores is too short to be of any value. Furthermore, only a few samples were taken from each core therefore any short-term trends are only weakly supported.

5.5.3 Particle Size

Up to twenty samples were taken from representative layers in each dried vibra-core

(ntotai = 58) and were processed using a laser particle sizer (Table 5.3). Among the Ekman cores, only 237, 239, 240, 242 and 247 (of 2004) were tested for grain size. One sample was taken per core and was processed using a Sedigraph 5100 (Table 5.3). 83

MeanDso Core Core Number Distance from % % % grain size Iff Type (Waypoint) Delta (km) n Sand Silt Clay (1MB) Gun) Vibra 310 1.5 20 13.08 59 27.83 23.44 25.24 Ekman 237 2.29 1 1.514 26.19 72.3 1.54 Ektnan 239 4.48 1 1.71 20.1 78.22 1.32 Ekman 240 5.84 1 1.11 24.25 74.6 1.16 Ekman 242 8.29 1 0.13 18.75 81.12 0.94 Vibra 309 11.52 20 5.16 45.17 49.7 6.9 6.65 Ekman 247 12.45 1 0 13.16 86.84 0.67 Vibra 311 12.8 18 9.29 10.97 15.91 19.74 43.41 Table 5.3 Grain size results for Kinaskan Lake cores

Results show no significant relationship between distance of coring sites from the delta and grain size (Figure 5.17).

30 •

25 - •>

20 *> • Vibra-cores "i ^ • Ekman ceres .8 Linear (Ekman cores) Linear (Vibra-cores) <5 15 •

10-

<•

5 •

• -a— —^— 0 - 1 1 — a 6 8 Distance (km) Figure 5.17 Mean grain size with distance from the Kinaskan Lake - Tatogga Lake delta

Among the three vibra-cores, the greatest average D50 grain size and the highest percent of sand content was found at KIN 310 (Mean D50 grain size = 23.44 um ± 25.24 ^m, 84

% sand = 59). Here, sediment-laden inflow across the Kinaskan - Tatogga delta and from two minor tributaries on the eastern lake-shore becomes a density current which slowly deposits heavier grains as it loses momentum.

Conditions are calmer at distal sites. Therefore, KIN 309 and KIN 311 have a finer mean D50 grain size and lower percent of sand than KIN 310 (KIN 309 Mean D50 grain size =

6.9 ^im ± 6.65 urn, KIN 309 % sand = 5.16; KIN 311 Mean D50 grain size = 19.74 um ±

43.41 um, KIN 311 % sand = 9.29).

Ekman cores show a similar decrease in mean grain size and proportion of sand with distance fromth e delta. However, only one sample was processed per Ekman.

When considering the average D50 grain size it is important to note that some sites have experienced greater variability in delivery of particles of a certain size over time (Figure

5.18).

D^o Grain Size (um) 80 100 120 200

50 4**

• KIN 309 • KIN 310 AKJN311

Figure 5.18 KIN 309, KIN 310 and KIN 311 D50 Grain size versus depth in core 85

D50 values in KIN 310 fluctuate between 2.2 um and 108.7 um throughout the core.

These probably represent a combination of discharge inputs from the source delta that vary between high magnitude floods following spring snowmelt to late summer discharges tied to rainstorm and/or ice melt sources. Conversely, KIN 309 and KIN 311 have similar temporal patterns of median grain sizes. Both sites have a smaller range of median grain sizes with several clusters of outliers. For KIN 309, most values fall between 2.4 um to 15.17 um except for two samples of 23 um and two clusters of granules and pebbles found between 3.0 and 4.5 m depth. These clasts are sub-rounded and their length, width and height, volume and mass were determined (Figures 5.19 a & b; Table 5.4).

Figures 5.19 a & b Photos illustrating (a) the massive clayey-silt deposit in Kin 309 at 3.5 m depth and (b.) showing one of the large clasts. Dry Depth Length Width Height Mass Volume Core (m) (mm) (mm) (mm) (e) (ml) KIN 309 3.5 2 2 1 0.015 KIN 309 3.5 5 3 1 0.077 - KIN 309 3.5 6 4 2 0.145 0.25 KIN 309 3.5 7 6 3 0.235 0.5 KIN 309 3.5 9 5 6 0.69 1 KIN 309 3.7 30 25 17 18.45 18

Table 5.4 Location and characteristics of granules and pebbles found in KIN 309 86

In KIN 311, D50 values range from 1.2 um to 40.1 um except for three outliers; 121.1

um and 150.4 um between 3.0 and 4.5 m depth, and 40.0 um at 1 m depth. No granules or pebbles were present.

5.5.4 X-Ray Fluorescence

X-Ray Fluorescence measurements were taken on a small number of samples to see if there was temporal (down core) variation in composition which might indicate changing

sources. Five samples were taken from KIN 311 because it has the longest undisturbed

record. The highest mean proportion of sample attributed to a certain element or compound was 59.74% of Si02. This was followed by 16.81% of A1203, 9.71% of Fe203, 6.47% of

MgO, and trace amounts of 19 others (Table 5.5; Figure 5.20). 87

Depth in Core

20 cm 100 cm 200 cm 370 cm 470 cm average MgO g.HUrZM •Q39| MMIM A1Z03 17.189 16.852 16.777 16.374 16.836 16.8056

Si02 57.173 60.037 60.693 60.832 59.981 59.7432

P205 0.231 0.242 0.243 0.234 0.26 0.242

S03 0.069 0.071 0.072 0.073 0.067 0.0704

K20 2.425 2.303 2.377 2.278 2.382 2.353 CaO 1.009 0.988 0.936 0.915 0.884 0.9464

Ti02 0.811 0.895 0.884 0.911 0.925 0.8852 v2o5 0.044 0.042 0.052 0.041 0.038 0.0434

Element or Cr203 0.068 0.036 0.046 0.041 0.034 0.045 Compound MnO 0.162 0.143 0.119 0.142 0.134 0.14

Fc203 9.865 9.592 9.189 9.739 10.142 9.7054 Co 0.008 0.006 0.007 Ni 0.043 0.031 0.031 0.031 0.026 0.0324 Cu 0.016 0.012 0.014 0.012 0.0135 Zn 0.016 0.015 0.015 0.017 0.338 0.0802 Rb 0.008 0.008 0.007 0.008 0.008 0.0078 Sr 0.011 0.012 0.012 0.011 0.011 0.0114 Y 0.004 0.003 0.003 0.003 0.003 0.0032 Zr 0.014 0.013 0.014 0.015 0.015 0.0142 Nb 0.002 0.001 0.0015 Ba 0.102 0.12 0.093 0.112 0.104 0.1062 Pb 0.003 0.003 0.004 0.003333

Table 5.5 Percent of element or compound composing sediment samples taken from five depths in KIN 311

Relative to the depth in the core, the proportion of each element or compound does

not fluctuate dramatically. This suggests limited variability in the source of sediments.

However, the most recent (surface) sediment does appear to be slightly depleted in SiC^. 88

Percent of Sample 10 20 30 40 50 60 I 1 , I I \

100 \

150

200

U 250 .9

a 300

350

[.. 1 1 1 » <> 1 1 1 500 Figure 5.20 Geological make-up of KIN 311. Trace amounts of other elements and compounds were found in the sedimentary sample (see Table 5.5) but overlap on the figure.Therefore , they have not been included in the graph's legend. 5.6 Chronology

The chronology of sedimentary deposition was estimated using couplet counts and radiocarbon dating as well as careful comparisons of similar sedimentary structures between adjacent cores. Similarities between couplets from Kinaskan Lake and varves from nearby alpine lakes suggest that Kinaskan may be varved. Radiocarbon dates which were collected from the two distal Kinaskan vibra-cores support this suggestion.

5.6.1 Couplet Counts

Couplets in the three vibra-cores share similar characteristics to varves identified by

Gilbert (1975) in Lillooet Lake. These characteristics include: alternations of light and dark lamina, textural differences between winter and summer deposits, contacts between laminae, 89 and comparison among cores. Distinct light and dark lamina compose the couplets of

Kinaskan Lake, and textural differences between lamina were identified under a 40 x microscope (Table 5.6). Because of difficulties matching marker lamina, several radiocarbon dates were generated to affirm the suspected annual deposition. Cesium-137 or Lead-210 testing was not conducted because they only provide estimates of recent dates (i.e. Cesium provides dates in the 1950's to 1960's while Lead offers dates within the past century).

Core Core Number Distance from Delta Total Length Total couplet Type (Waypoint) (km) (*» counts Vibra 310 1.5 4.268 1146 Ekman E4-06 1.7 0.1095 18 Ekman E3-06 6 0.0895 17 Ekman E2-06 11.5 0.086 21 Vibra 309 11.52 4.193 2055 Vibra 311 12.8 4.606 2247 Ekman El-06 12.9 0.081 16 Table 5.6 Total couplet counts in each core 5.6.2 Radiocarbon Dates

Four isolated deposits of woody debris were recovered from KIN 309 and KIN 311

and processed by Isotrace Laboratory at the University of Toronto (Table 5.7; Table 5.8).

Core Number Depth Sample ID Core Type (Waypoint) in core (cm) Sample Type Sample Mass (g) A Vibra KIN 309 330 Woody debris 0.412 B* Vibra KIN 309 354 Woody debris 0.051 c* Vibra KIN311 269 Woody debris 0.197 D Vibra KTN311 457 Woody debris 0.213

Table 5.7 Radiocarbon sample details for Kinaskan Lake. 90 Accumu. Rate Couplet (based on Accunra.Rate count couplet Cat Years (based on Sample to depth count) Uncal. 14C BPp&S% Cat Years BP) ID in core (mm/yr) Years BP cJL) (mm/yr) A 1713 19.26 2450 ± 50 2782-2407 11.86-13.71 B* 1832 19.32 5030 ±60 5972-5697 5.93-6.21 c* 1269 21.20 3480 ±50 3942 - 3692 6.82-7.29 D 2235 20.45 2050 ±50 2202-1947 20.75-23.47

Table 5.8 Radiocarbon results for Kinaskan Lake (*these two samples provided anomalous radiocarbon dates and have been disregarded).

When examined alone, 14C dates are inconsistent. In KIN 309 the two samples were within 25 cm of each other but yield radiocarbon dates 2,500 years apart. In KIN 311 the chronology is inverted with an older date of 3,480 +/- 50 Uncal. 14C BP at a much shallower depth in the core. Anomalously older dates can occur for several reasons. Old carbon in the watershed, including rare coal deposits, could have been flushed into the lake. Avalanches and floods could introduce older buried organics from hill-slopes and proximal floodplain sequences. In KIN 309 the deeper sample was found in a highly disturbed section of

sediment while the shallower sample was located in a well-layered section. The disturbed

section has been interpreted to have resulted from mass wasting in the watershed. Therefore, the difference of ~2,500 radiocarbon years between the two samples may have resulted from

old carbon in the watershed being deposited in the lake.

Couplet counts appear to be correlated to the two accepted radiocarbon dates. The

KIN 309 radiocarbon sample was taken from couplet 1713 and provided a calibrated age of

2782 - 2407 years. The KIN 311 radiocarbon sample was taken from couplet 2235 and

provided a calibrated age of 2202 - 1947 years. Therefore, couplet counts have been

interpreted to represent annual (varved) deposits. Chapter 6 - Discussion

6.1 Introduction

Since the close of the Pleistocene, alpine glaciers in the Upper Iskut River watershed have influenced the hydrology and sedimentology of Kinaskan Lake. Above or below average inputs in melt-water and sediment input to Kinaskan Lake can be related to climate variability that has occurred on a seasonal and longer (i.e. millennial) time scale. To isolate long-term changes in sedimentary environments, both the results from Chapter 5 and the ensuing discussion will highlight: (a) the relative importance of contributing sub-watersheds

(e.g. glaciated sub-watersheds versus un-glaciated sub-watersheds), (b) the in-lake depositional patterns of sediment (e.g. the underflow/interflow/overflow components and the role of bathymetry) (c) typical structure and composition of sediment on a down-lake

(spatial) and on a temporal scale, and (d) typical seasonal patterns of hydrology and sedimentation. Finally, comparisons will be made between long-term patterns of sedimentation in Kinaskan Lake and in other nearby lakes to highlight major episodes of climate change.

6.2 Sedimentation in Kinaskan Lake

Among the sub-watersheds which feed the Upper Iskut River, contributions from

Todagin Creek - Tsatia Creek provide the greatest load of sediment and the highest volume of discharge to Kinaskan Lake. After the coarsest glacial sediments settle in Tatogga Lake, sediment-laden underflows create deposits on the Kinaskan Lake - Tatogga Lake delta fore- set while underflows (during deglaciation) and interflows and overflows (post-deglaciation) transport sediments down-lake. Deposits from interflows and overflows occur where sediment settles from suspension and are typically conformable to

91 92 underlying bedrock or de-glacial sediment reflectors (Gilbert et al., 1997). Sub-bottom acoustic images from Kinaskan Lake show that sediments composed of multiple, conformable reflectors cap the majority of profiled sites. In the central region of the lake, this cap is underlain by unconformable sediments which in-fill troughs. These sediments were deposited from underflows which dominated during deglaciation. Since deglaciation, sedimentation in the delta region has continued to be dominated by underflows. However, most sedimentation of distal regions has been through overflows and interflows which leave well-laminated, conformable deposits.

In aerial imagery, turbid plumes of inflowing water from Tatogga Lake into Kinaskan

Lake are deflected to the right (west) side of the lake. This could be related to the position of the main deltaic channel, which enters the lake towards the right. Alternatively, this persistent pattern could be related to the Coriolis effect, which has been found responsible for causing higher sedimentation on the right (west) side of Meziadin Lake and Atlin Lake,

British Columbia (Gilbert and Butler, 2004; Serink, 2004). In acoustic images from Kinaskan

Lake, sediment accumulations do not appear significantly greater on the right (west) side of the lake. Similarly, at Bowser Lake, British Columbia, no evidence of the Coriolis effect was found. This was attributed to the high velocity of turbidity currents which overcame the

Coriolis forces (Gilbert et al., 1997). Additional inputs of runoff and sediment in Kinaskan

Lake come from minor tributaries which feed directly into the lake or indirectly into lakes of the upper watershed. Occasionally, sub-aqueous and sub-aerial mass-wasting events along the Kinaskan shoreline create localized deposits of sediment on the lake-bed.

Differences in the structure and composition of sediment in Kinaskan Lake are related to (a) the type and volume of sediment coming from different sources and (b) bathymetric 93 controls. At proximal sites (KIN 310, E4-06 and 237) sediment is composed of roughly parallel couplets of coarse (i.e. sand and coarse silt), light-colored lamina and silty, dark- colored lamina. The greater average couplet thickness of 3.76 mm (~ 0.045 g/cm2/yr, assuming a density of 1.2 g/cm3), the range of couplet thickness (1 mm < x < 29.83 mm), the somewhat lower average organic content of 6.73% and the range of organic content (5.37% < x < 9.53%) in proximal sites, reflect variable inputs of clastic-rich sediment. Acoustic records suggest the occurrence of frequent but weak turbidity currents which spill down the delta and lose momentum at these sites. Alternatively, occasional stronger turbidity currents may rise up off the down-lake sill or from reflection of turbidity currents off the up-lake slope of the sill followed by a loss of momentum. The latter process was responsible for creating thick sediment deposits 13 km down-lake from the Bowser Lake delta (Gilbert et al., 2006).

Although insufficient material was available for radiocarbon dating these proximal cores, their sedimentary structure closely mirrors annual accumulations in Atlin Lake and

Green Lake, both in British Columbia (Gilbert et al., 2006; Menounos et al., 2006) (Figures

6.1a,b,c&d).

Figures 6.1 a, b, c & d Photos of sediment structure for (a) KIN 310 (depth 3.12 m - 3.18 m), (b) E4-06 (depth

0.05 m - 0.11 m), (c) Atlin Lake gravity core (note: photo in grey scale), showing varved deposits and (d)

Green Lake core (depth 1.34 m - 1.38 m, note: photo of thin section) showing annual laminae (Gilbert et al.,

2006; Menounos et al., 2006) 94

In each instance, annual deposits are composed of coarse material (often composed of multiple, sub-annual and super-imposed layers of sands and coarse silts) which fines upwards into thick layers of fine silts. The coarser components of Atlin Lake sediment originated at

Llewellyn Glacier and were transported by underflows and interflows to their site of deposition (Gilbert et al., 2006). Thick couplets from Green Lake are simlar in appearance to those from Atlin Lake.

The three proximal Kinaskan Lake cores are separated from distal cores by a sill which peaks at approximately 4 km down-lake from the delta. Distal sediments are arranged in couplets of light-colored fine sand and silt lamina and dark-colored silt and clay lamina. In

KIN 309, the average couplet thickness is 2.85 mm ± 2.04 mm. Down-lake, the average couplet thickness for KIN 311 is 2.05 mm ±1.2 mm. The finer overall grain size, the decrease in average thickness of distal couplets, and increase in average organic content of distal sediments relative to proximal values, reflects lower inputs of clastic sediment (for

KIN 309 = 0.0342 g/cm2/yr assuming a density of 1.2 g/cm3; for KIN 311 = 0.0246 g/cm2/yr assuming a density of 1.2 g/cm ) in a less energetic environment. More conformable acoustic reflectors suggest the greater dominance of settling from suspension and the reduced influence of turbidity currents. Sediment from the upper watershed via Todagin Creek travels as interflows and overflows to these remote sites (Section 5.2). Additional contributions of suspended sediment come from minor tributaries which enter along the lakeshore. For these

silt and clay grains to fall from suspension, conditions must be extremely calm. However, flocculation of particles, documented in Lillooet Lake (Hodder and Gilbert, 2007), can accelerate their settling. During movement through the water column, these floes can

scavenge additional particles. 95

A total of four radiocarbon dates were collected from the two distal vibra-cores. A date reversal in KIN 311 (269 cm depth = 3942 - 3692 cal. Years BP; 457 cm depth = 2202 -

1947 cal. Years BP) probably resulted from older carbon materials in the watershed being deposited higher in the core (Oswald et al., 2005; Bartlein et al., 1995). The younger radiocarbon date, lower in the core, is assumed to be representative of the time of deposition.

In KIN 309, both dates were taken from within 25 cm of each other but were ~2.5 ka apart in age (330 cm depth = 2782 - 2407 cal. Years BP; 354 cm depth = 5972 - 5697 cal. Years

BP). The older date was taken from a highly disturbed section of the core whereas the younger date was taken among finely laminated and undisturbed sediments. In KIN 311 and

KIN 309, radiocarbon dates suggest that older materials are being introduced from hillslope processes. However, younger dates (2202-1947 cal. Years BP and 2782-2407 cal. Years BP), which approach couplet count values, appear to be representative of the time of deposition.

The sedimentary structure of the distal cores, despite some variability, is similar to the structure of annual deposits from other low-energy glacio-lacustrine environments

(Figures 6.2 a, b, c & d).

Figures 6.2 a, b, c & d Photos of sediment structure for (a) KIN 309 (depth 0.9 m - 1.5 m), (b) KIN 311 (depth

0.5 m - 0.6 m), (c) Atlin Lake gravity core showing annual varves from a low sediment input period (note: photo in grey scale), and in (d) Green Lake core (varves from 1900 to 1944) (Gilbert et al., 2006; Menounos et al, 2006) 96

The fine varves from Atlin Lake were deposited after a pro-glacial lake formed at the toe of Llewellyn Glacier. This reservoir for melt-water and sediment reduced the volume of inputs into Atlin Lake. At Green Lake, the micro-laminated and slightly diffuse varves formed before high volumes of sediment became mobilized by the 1991 event. The similarity in structures with Kinaskan Lake sediments suggests a similar reduction in inputs commensurate with the distal location and/or lower overall inputs over time.

The greatest inputs of discharge and sediment in Kinaskan Lake occur between May to August and are related to nival and glacial melt. A second, smaller peak in discharge and sediment occurs between October and November. In this instance, runoff from autumn precipitation transports loose sediments into the lake (Chapter 3). No measurements of the concentration of suspended sediment were taken from the inflowing streams and multiple events (i.e. fine sand layers within a couplet) could not be identified between successive winter caps in the cores. Between episodes of high inputs (i.e. spring to summer) and low inputs (i.e. winter conditions), a transition period of moderate to low sedimentation is related to storage of water and sediment in frozen forms and/or exhaustion of loose sediments.

Similar timing of sedimentation has been identified elsewhere in Northern British Columbia

(Spooner et al., 2002; Mazzucchi et al., 2003; Gilbert et al., 1997; Gilbert et al., 2004).

Overall, since deglaciation of the watershed at -11,500 BP, up to 40 m of sediment have accumulated in Kinaskan Lake (Spooner et al., 2002; This study.). 97

6.3 Long-term Changes in Lacustrine Sedimentation and Climate Change in British

Columbia: Contextualizing the Kinaskan Lake Sedimentary Record

The longer-term, although lower resolution, sedimentary records from four alpine lakes in Northern British Columbia are compared here with results from Kinaskan Lake

(Figures 6.4 and 6.5). These four lakes were chosen because of their similarity in location, their elevation, and their drainage area relative to lake dimensions. Characteristics of these four lakes have been summarized (Table 6.1) and their locations can be seen in figure 6.3.

Results from loss-on-ignition and depositional thickness (in some instances sediment influx rate was used instead) from sediment samples taken from these lakes are compared with the record from Kinaskan Lake and within established episodes of climate change.

Drainage Length, Width Depth Elevation Area Site Location (km) (km) OB) (m a.s.I) (km2) Author Altin 59° 13'N, Lake 134° W 106 3-8 220 1500 6410 Serink, 2004 58°53'N, Pyramid 129° 50' Mazzucchi et al., Lake W 0.48 N/A 9 -1525 N/A 2003 57°35'N, Skinny 130° 10' Lake W 1 N/A 7 910 N/A Spooner et al., 2002 Kinaskan 57° N, Lake 130° W 17 1.5 135 1200 1257 This study. 56° 30' N, Bowser 129° 30' Lake W 23 1.5 119 368 1386 Gilbert etal., 1997

Table 6.1 Sites of the Northern British Columbia lakes chosen for comparison with Kinaskan Lake. 98

4- thin infill •fr thick deglacial infill * thick deglacial through Holocene infill

Figure 6.3 Map showing lakes whose sediment will be compared with results from Kinaskan Lake (adapted from Gilbert et al, 2006)

6.3.1 Selected British Columbia Lakes for Comparison of Sedimentary Records with

Kinaskan

Sedimentary records from Skinny Lake and Pyramid Lake were chosen for comparison with results from Kinaskan Lake because both lakes are located near (or within) the Upper Iskut River watershed. Skinny Lake is found within the Todagin Creek Sub- watershed. It is fed principally by a small tributary of Todagin Creek, which transects swampland before reaching the lake. Outflow from Skinny Lake is minimal. As a result, sedimentation rates are low for clastic material relative to organic matter. A series of 99 percussion cores were taken from Skinny Lake by Spooner et al. (2002) and composition was determined through pollen analysis, carbon/nitrogen ratios, magnetic susceptibility, grain size and loss-on-ignition. To understand changes in sediment structure, sediment influx rate values were calculated. Radiocarbon dates suggest a basal age of-12 ka BP for Skinny Lake sediment cores (Spooner et al., 2002).

Slightly northwest of Kinaskan Lake is a former tarn known as Pyramid Lake. Today, the main source of water to Pyramid Lake is from a ground-water fed stream which enters at the south end of the lake. Outflow passes through conduits in metamorphic and igneous bedrock (Mazzucchi et al., 2003). Three percussion cores of 2.7 to 3.5 m length were taken at approximately equal intervals down the center of the lake. Sediment was processed for pollen, dry bulk density, loss-on-ignition, magnetic susceptibility, and charcoal analysis.

Results show that Pyramid Lake sediment is predominately clay to silt sized and is deposited at a fairly constant sedimentation rate. The maximum age of sediment in the Pyramid core is approximately 11.8 ka (Mazzucchi et al., 2003).

Bowser Lake is found in the northern Coast Mountains of British Columbia, south­ east of Kinaskan. The contributing basin for Bowser contains significantly greater glacial cover (46%) than the other four lakes (Kinaskan Lake = 0.75%; Skinny Lake = 1.6%; Atlin

Lake = 9.7%; Pyramid Lake = N/A) (Gilbert et al., 1997; This paper.; Spooner et al., 2002;

Serink, 2004). However, sediment in Bowser Lake tends to be fine (e.g. clays and fine silts) because small lakes between the glaciers and Bowser Lake intercept coarse (e.g. coarse silts, sand and gravel) components.

Similar to Kinaskan Lake, distribution of sediment within Bowser Lake is closely tied to its bathymetry. Bowser Lake is divided into two basins by a sill. The sill is located 100 approximately 14 km from the delta, and shoals up to 40 m water depths (Gilbert et al,

1997). Sediment deposits in the proximal basin are thicker than in the distal basin. Like

Kinaskan Lake, sedimentation in the distal basin is dependent on overflows and interflows from the delta and from minor tributaries. Within the proximal basin, maximum lamina thicknesses occur at the sill. As turbidity currents travel down-lake from the delta, they are confronted by the sill. They rapidly lose momentum and leave pools of sediment up-lake of the sill. Multiple Percussion cores (0.7 - 2.7 m), gravity cores (0.46 - 0.67 m) and Ekman cores were taken from the proximal and distal basins of Bowser Lake. Sediments were tested for loss-on-ignition and grain size (wet sieve and Sedigraph). Additionally, measurements of lamina thickness were taken. The five cores from Bowser Lake record the sedimentation between 1992 to 1967 AD (Gilbert et al, 1997). Consequently, they will be compared with the upper portion of the Kinaskan Lake cores.

Among all five lakes, Atlin Lake has the greatest surface area, depth, and contributing drainage area. It is located father north than the other lakes and at a higher elevation. A sub- basin of Atlin Lake, Lewellyn Inlet, is fed by glacial melt from the Juneau Ice Field (Gilbert et al., 2006; Serink, 2004). Over the past half century, sedimentation in Lewellyn Inlet decreased due to formation of a proglacial lake at Lewellyn Glacier. Sediment in Atlin Lake has been verified as varved, and consists of fine silts and clays. A series of sediment samples;

Ekman cores (~6 cm) and percussion cores (20 cm - 124 cm), were taken from Lewellyn

Inlet and tested for loss-on-ignition and grain size. Additional measurements of varve thickness were also made. The two percussion cores being compared to Kinaskan Lake cores cover the time periods; 1965 to 1760 AD (P3) and 1950 to 1725 AD (P2) (Serink, 2004;

Gilbert etal., 2006). Skinny Lake Loss-on-ignition Pyramid Lake Loss-on-ignition Kinaskan Lake Loss-on-ignition Atlin Lake Loss-on-ignition Percent lost after Percent lost after 1 hr at Percent organic Percent organic 1 hr at 550 C Percent lost after 1 hr at 550 C 550Candlhrat950C matter (P2) matter (P3) 0 20 40 63 0 4 8 12 0 12 0 0.5 1.01.5 0.25 0.75 1.25 0 "lUncal. "C Yeats BP 0 ,4C Years BP

1939+/-46

^3000 o •3130+/-60

208 Varve Yrs BP KIN E1-06 KIN 309 -KIN E2-06 Serink, 2004 KIN 310 KIN E3-06 KIN 311 -KIN E4-06

Uncal. "C Years BP

;(KIN311 2.69 m) lll,433+/-97 ;3480 +/- 50

395 (KIN 309 3.3 m) Mazzucchi et al., 2003 2450+/- 50 "C Years BP Spooner et al., 2002 (KIN 309 3.54 m) 5030+/-60

Approximate isochron (KIN 311 4.57 m) 2050+/-50

9 TOe n"i Skinny Lake Kinaskan Lake Atlin Lake Bowser Lake 0\ Couplet Thickness Varve Thickness Laminae Thickness '•u Sediment Influx Rate g/cm2/yr 25 year filter(mm ) o .0.!.^.t?,6.7 8 9i0 :j rtf,t>1 .36 .06 C.I 0-12 Nofilter(mm) Uncal. "C Years BP 1 0 4 8 12 16 s. 1936+/-46 § O r>-t > o § 3t Gilbert et al., 1997

CO (KIN 309 3.3 m) 2450 +/- 50 cr O Spooner et al., 2002 (KIN 309 3.54 m) o 5030 +/-60

(KIN 311 4.57 m) 2050 +/- 50

a' Approximate isochron

S" OS

O 103

6.3.2 Long-term Trends in Sedimentation and Climate Conditions of British Columbia

Between approximately 11 ka BP and 8.2 ka BP, the contributing basins for Skinny

Lake and Pyramid Lake underwent deglaciation (Spooner et al., 2002; Mazzucchi et al.,

2003) Sediment samples from this period have low loss-on-ignition values (Skinny Lake

(325 cm - 225 cm depth in core) = -8%); Pyramid Lake (395 cm - 265 cm depth in core) =

~2%), related to a lack of established vegetation. Pollen counts from Pyramid Lake sediments, dating from 10.6 ka BP, are also low (920 grains per cm2/year). These low pollen counts could be related to dilution of higher concentrations of pollen by rapid clastic sedimentation, or related to low vegetation diversity and cover in the watershed (Spooner et al., 2002; Mazzucchi et al., 2003). Sediment samples from this period also indicate high rates of clastic sediment inputs related to the mobilization of sub-/supra-/en-glacial till. In Skinny

Lake, the sediment influx rate at this depth exceeds all other values throughout the core (0.06 to 0.12 g/cm2/year). Sediment composition ranges from coarse-grained lodgement till to thick deposits of glaciolacustrine sediment with patches of ice-rafted debris. In the Pyramid Lake core, the lowest facies is a massive diamicton composed of inorganic silt, clay, sand and unsorted cobbles. Mazzucchi et al. (2003) believe this diamicton formed at the toe of a retreating late Pleistocene glacier. Acoustic sub-bottom images from Kinaskan Lake show evidence (unconformable in-filled troughs) that clastic-rich density currents dominated the down-lake region during deglaciation. These images also show the maximum volume of sediment deposited in Kinaskan Lake since deglaciation. For sediment thicknesses of 30 m

(maximum in proximal regions) and 40 m (maximum in distal regions) to form, an annual accumulation rate of 2.6 mm/yr (proximal) and 3.5 mm/yr (distal) would be necessary. Over the past two millennia, proximal accumulation rates (4.42 mm/yr ± 3.37 mm/yr in KIN 310) 104 and distal accumulation rates (2.85 mm/yr ± 2.04 mm/yr in KIN 309 and 2.05 mm/yr ± 1.02 mm/yr in KIN 311) have been well below these rates. Therefore, at one time, sedimentation rates must have been considerably greater.

Following 9 ka BP, loss-on-ignition values increased sharply (Skinny Lake (-8.9 ka

Uncal. 14C BP or -8.1 cal. BP) = 60%; Pyramid Lake (-8.2 ka BP) = 9%). Inputs of organic matter came from unvegetated soils on newly exposed slopes and high aquatic productivity

(Spooner et al., 2002; Mazzucchi et al., 2003). Warm and dry climate conditions are reflected in non-biogenic marl at the base of the Skinny Lake core. The first colonies of vegetation in the Pyramid Lake watershed: white spruce, birch, sedge, fern and alder, were eliminated by a series of forest fires related to the dry conditions. Marker deposits of charcoal in Pyramid

Lake sediment date to -8.2 ka BP and -6.8 ka BP (Mazzucchi et al., 2003). Concurrently, the sediment influx rate in Pyramid Lake and Skinny Lake dropped drastically (Pyramid Lake = from 0.06 to 0.02 g/cm2/year). This was due to decreased glacier cover and a resultant drop in production of clastic sediment (Spooner et al., 2002; Mazzucchi et al., 2003).

Sedimentation in Skinny Lake reached a core-long minimum between 7.3 ka BP and

7 ka BP. It is thought that a slight increase in precipitation, from 6.8 ka BP and 4.8 ka BP, caused clastic sedimentation to peak around 6.6 ka BP. Evidence of these moister conditions has also been found in sediment from Fiddler's Pond (White and Matthewes, 1982),

Kettlehole Pond (Cwynar, 1988) and Susie Lake (Spooner et al., 1997).

Despite favorable conditions for the productivity of vegetation and a mix of warm temperatures and frequent precipitation events characteristic of this Holocene Thermal

Maximum (Hypsithermal), organic content in sediments stabilized at 45% for Skinny Lake

and 4% for Pyramid Lake. Conversely, pollen counts from Pyramid Lake sediment 105

(dominated by serai species) increased (-17,630 grains per cm2/year). In particular, the pollen is composed of an assorted assemblage dominated by alder and birch. According to

Mazzucchi et al. (2003), this indicates the presence of an open forest. Unfortunately, these moist conditions were not sufficient to douse a series of forest fires in the Pyramid Lake watershed. Between 210 cm and 160 cm depth (9 to 7 ka BP) in the Pyramid Lake core -11 individual layers of charcoal were identified (Mazzucchi et al., 2003).

By 6 ka BP, as seen in the sedimentary record of Susie Lake (Spooner et al., 1997) and Fiddler's Pond (White and Mathewes, 1982), cool conditions resumed. Pollen data, the interpretation of tree-line dynamics (which did not change) and sediment influx rate (which fell to 0.02 g/cm2/year) from Skinny Lake, suggest that microclimatic moisture trends may have trumped wide-spread climate conditions (Spooner et al., 2002; Demarchi, 1996) around this time. At Pyramid Lake, productivity remained high with the highest concentration of conifer needles found between 5.5 ka BP and 4.7 ka BP, and the maximum concentration of pollen occurring in 4.9 ka BP (Mazzucchi et al., 2003).

An increase in moisture between 5.1 ka and 4.4 ka BP, characterizing the mesothermic interval, marked the close of the Holocene Thermal Maximum (Hypsithermal)

(Hebda, 1995). In Skinny Lake, organic content peaked between 100 cm and 75 cm depth

(average = 60%). At nearby Pyramid Lake, loss-on-ignition peaked at 140 cm depth (-8.5%),

130 cm depth (~ 8%) and 120 cm depth (~ 8.6%). (Spooner et al., 2002; Mazzucchi et al.,

2003). The increasingly moist conditions are probably responsible for reduced counts of

Artemisia and Ericaceae pollen in this section of the Pyramid Lake record (Mazzucchi et al.,

2003). 106

The response of clastic sedimentation to increased moisture differs between Pyramid and Skinny Lake. Four minerogenic layers, clustered between 120 cm and 140 cm depth (5.1 to 4.4 ka BP) in the Pyramid Lake core have been attributed to orographic rainstorms and hill-slope erosion. Although only one sediment core from Pyramid Lake was chosen for comparison with Kinaskan Lake, it is important to note that these same layers were found in all three of the Pyramid Lake percussion cores (Mazzucchi et al., 2003). In Skinny Lake, sedimentation rates appear unchanged at 0.02 g/cm2/year (Spooner et al., 2002).

Between 4 ka BP and 3 ka BP, the British Columbia climate cooled as it transitioned towards modern conditions. At Skinny Lake, loss-on-ignition values dropped to 50% where they stablized until 1.5 ka BP when they dropped again to 40% (Spooner et al., 2002). At

Pyramid Lake, changes in organic content lagged, reaching a major low of 1% around 2.5 ka

BP and then establishing a new mean (of 8%) by -1.5 ka BP (Mazzucchi et al., 2003).

Spooner et al. (2002) suggest that decreases in organic matter were related to falling productivity, tied to: extended ice-cover, decreased solar insolation and cooler water temperatures. Although sediment influx rate in Skinny Lake does not appear to have increased during these Neoglacial years (approximately 3 ka to 2 ka BP), magnetic susceptibility results suggest increasing ice-rafted debris. Meanwhile, heavier snowfall caused glaciers to advance in the Rockies and in the British Columbian Cordillera (e.g.

Tiedemann Glacier Advance; Dunn Peak Advance; Battle Mountain Advance; Peyto Glacier)

(Osborn and Luckman, 1988; Gardner and Jones, 1985; Alley, 1976; Lowdon et al., 1971).

The record in the Kinaskan Lake vibra-cores begins during the Neoglacial (~2 ka

BP). Elevated levels of organic matter, high proportions of sand, and thick lamina in the lower portion of distal cores (from ~2 —1 ka BP), are matched by a similar increase in loss- 107 on-ignition in Pyramid Lake and Skinny Lake, and greater sediment influx variability at

Skinny Lake.

Within this section of the KIN 309 core, several drop-stones were found in a massive silt matrix. Where similar deposits have been found in lake sediments, including those from

Pyramid Lake, B.C., Skinny Lake, B.C., and Saalian Glacial Lake, NW Germany, they have been attributed to several autochthonous and allochthonous processes (Mazzucchi et al.,

2003; Spooner et al., 2002; Winsemann et al., 2004). Autochthonous processes include seismically triggered sub-aqueous landslides which leave a massive deposit of silts and clays containing unusually large clasts (Blais et al., 1997; Mazzucchi et al., 2003). Allochthonous processes, most effective when the lake is ice-covered, include localized erosion (possibly related to greater precipitation), avalanches and debris flows (Gardner, 1983; Spooner et al.,

2002; Johnson, 1997). In each of these instances, sediment can run-out long distances on the ice (Luckman, 1975; Johnson, 1997). Due to the darker color of these materials, they melt through the lake ice to settle in situ or are rafted during break-up to other sites for deposition

(Gardner, 1983; Spooner et al., 2002; Johnson, 1997).

Around 1300 years BP, the onset of Little Ice Age conditions caused a shift in sediment composition and structure for many of these lakes. At Skinny and Pyramid Lakes, sharp drops in loss-on-ignition values were followed by a rapid recovery (Spooner et al.,

2002; Mazzucchi et al., 2003). In Kinaskan Lake, the average loss-on-ignition was 7.65% below 2.35 m (-1388 years BP) in KIN 309, and 6.99% below 2.55 m (-1200 years BP) in

KIN 311. At 2.35 m depth in KIN 309, loss-on-ignition dropped to -6.5%. Above this depth, organic content recovered rapidly, establishing a new mean of 6.5%. At 2.55 m depth in KIN

311, loss-on-ignition fell to -6.3%. Above this depth, organic content continued to increase 108 up to the core surface (new mean = 7.3%). These abrupt decreases in organic content are related to falling temperatures and/or decreased precipitation. The rapid recovery suggests renewed productivity in the watershed likely related to warmer and moister conditions. These moister conditions are probably responsible for increased sedimentation rates between ~1300 years BP and -1200 years BP and highly variable couplet thicknesses between -1200 years

BP and-1000 years BP.

As organic content in the upper half of KIN 309 declined and a new mean of organic content in the upper half of KIN 311 was established, lamina thicknesses decreased and the proportion of sand fell. In KIN 309 and KIN 311, decreased sedimentation between -1000 years BP and -650 years BP was probably due to falling temperatures and/or decreased precipitation. A slight increase in sedimentation rates (and variability) from -650 years BP to

-450 years BP implies higher and more variable temperatures and/or precipitation. From

-450 years BP to the present, sedimentation patterns in KIN 309 and KIN 311 are inverted.

In KIN 309, couplet thicknesses decreased between -450 years BP and -250 years BP.

Between -250 years BP and -100 years BP, sedimentation rates increased. From -100 years

BP to the present, sedimentation rates declined but organic content increased slightly.

Concomitantly, in KIN 311 couplet thicknesses increased between -450 years BP and -200 years BP but sand content decreased. Between 200 years BP and 50 years BP, couplet thicknesses fluctuated with a slight overall decrease. Since 50 years BP, sedimentation rates decreased slightly with a slight increase in organic content. A similar increase in organic content is reflected in the 2006 Ekman cores, suggesting warmer temperatures and/or increased precipitation. 109

The recent inversion of sedimentation rates in the distal cores may be related to a lag in the response of sub-watersheds to climate conditions and the relative importance of the sub-watersheds to sedimentation at each coring site. Alternatively, localized patterns of clastic sedimentation (including the release of stored sediments within these sub-watersheds or focusing of sediment from nearby lake slopes) could cause anomalies in couplet thicknesses at each site.

Sedimentation at Atlin Lake is described in detail by Serink (2004). Between 1725

AD and 1780 AD, sedimentation rates and varve thickness variability increased in both P2 and P3. Greater production and delivery of sediment were caused by shifts in the extent of alpine glaciers. From 1780 AD to 1800 AD, cool and dry conditions caused sediment accumulation and variability to decrease in P3. Slight advances (or retreats) of alpine glaciers provided thick inputs of sediment in P3 between 1800 AD and 1900 AD. During the later half of the 19th century, the thickest varves in P3 were produced. In P2 and P3, the start of the

20th century is characterized by a greater frequency of anomalously thick varves. However, the overall thickness of varves in P3 declined. A transition in the extent of ice cover from

1925 AD to 1950 AD caused greater varve thickness and maximum variability in P3. Varve thicknesses declined significantly following the formation of a pro-glacial lake at the toe of

Llewellyn Glacier in 1950 AD. Although both P2 and P3 tell similar stories of conditions in the Atlin watershed, slight variations are related to differences in sediment composition and deposition processes. P3 is composed of coarse silts and sands which are deposited as underflows, where as P2 is composed of clay and silt-sized, overflow sediments.

The Bowser Lake record, which shows lamina thickness from 1967 AD to 1992 AD, picks up where the Atlin record terminates (Gilbert et al., 1997). Two massive deposits in 110

Bowser core P2, dating to the start of the 20* century, are not presented in the diagram.

These deposits are related to an outburst flood from a glacier dammed lake (probably Tide

Lake) within the Bowser watershed. Following this catastrophic event, variability in lamina thickness decreased. Glacier dammed lakes, which were not already drained, appear to have stabilized. Minor fluctuations, including deposits of fine sand, are related to greater winter and summer precipitation and mass-wasting off of hill-slopes. Chapter 7 - Conclusion

7.1 Summary and Conclusions of Findings

This thesis focuses on a high-resolution and long-term investigation of Holocene environmental change in a remote region of Northwestern British Columbia. Results of this investigation show that different aspects of lake and/or watershed hydrology control sediment distribution in the proximal and distal regions of Kinaskan. Proximal sedimentation is driven by discharge events off of the delta with turbidity deposits most likely providing a primary record of nival melt and possibly a secondary record of autumn precipitation events. Distal sedimentation, which is dominated by interflows and overflows, is more closely conditioned by prolonged episodes of high suspended sediment flux from nival and glacier melt across the whole Upper Iskut River watershed.

However, the very small area (9.43 km ) and volume of contemporary glacier ice cover may not be sufficient to show a strong glacier-melt signal. Therefore, departures in sediment accumulation rate from the distal region may reflect persistent departures in temperature and precipitation affecting the whole of the upper Iskut River watershed.

This assumes sediment delivery in the watershed is "transport limited" which is likely for two reasons: a) the Bowser Lake Group rocks are easily eroded and there is ample evidence of temporary storage in the upper watershed and b) although small, glacier melt- runoff in the past may have provided large volumes of sediment to the trunk valleys.

7.1.1 Modern Patterns of Hydrology and Sedimentation

Modern lake hydrology shows a late spring - early summer heating pattern which is typical for similarly sized lakes of this region. Sedimentation is driven by turbid plumes, containing silt and clay sized grains, which travel down-lake as either interflow

111 112 or overflow currents. These grains settle from the water column under extremely calm conditions. Sedimentation is supplemented by underflows which are contained in the pro- delta region up-lake of a major sill. These currents are weakened by the storage and deposition of coarser sediment fractions in Tatogga Lake before entering Kinaskan Lake.

This sedimentation pattern differs from many glacier-fed lakes (i.e. Bowser Lake, Green

Lake and Atlin Lake) whose sedimentation is dominated by strong turbidity currents which travel for greater distances (Gilbert et al., 1997; Menounos et al., 2006; Gilbert et al., 2006).

7.1.2 Long-term Sedimentation

Sub-bottom images show smooth-surfaced, acoustically transparent materials deposited on the lake bottom in the pro-delta region. They were probably deposited from turbidity currents on the delta fore-slope. No long-term changes in sedimentation patterns can be discerned from acoustic images. Down-lake, the central region of the lake is occupied by sediments that in-fill trenches. The deeper sediments appear well-layered but unconformable to the lowest reflector. They were probably deposited from high-energy bottom currents during deglaciation. These are capped by well-layered, conformable sediments which probably settled mainly from suspension. This cap extends to the slopes of each shore with no apparent thickening due to Coriolis effects. The absence of sand

(graded or otherwise) sediments in this cap suggests that turbidity currents no longer play a strong role in distal sedimentation. In isolated regions, this cap is slightly hummocky.

These sediments were also probably deposited from suspension. Overall, distal sub- bottom acoustic results indicate a transition from energetic sedimentation of deglaciation 113 to calmer deposition and a decreasing importance of underflows relative to overflows and interflows.

7.1.3 Sediment Structure and Composition

Sediments are arranged in pairs of a coarse-grained light-colored layer which is overlain by a fine-grained dark-colored layer. The structure of these couplets is similar to the structure of varves observed in Atlin Lake and Green Lake (Serink, 2004; Menounos et al., 2006).

Down-lake, mean couplet thickness and variance decreases. Thicker and more variable couplets of proximal cores KIN 310 and KIN E4-06 are related to irregular discharges crossing the delta and thick deposits from turbidity currents. In one of the distal cores (KIN 309) a massive deposit of disturbed sediment was found. Overall, the timing and magnitude of long-term changes in couplet thicknesses and composition in the two distal vibra-cores (KIN 309 and KIN 311) are similar and correspond reasonably well to sedimentation peaks in nearby lakes (Spooner et al., 2002; Mazzucchi et al., 2003;

Gilbert et al., 1997; Gilbert et al., 2006; Serink, 2004).

Grain size, loss-on-ignition and mineralogy, were used to identify down-lake and down-core variations. Grain sizes decrease from the delta to distal regions, save for several drop-stones which were found at the base of KIN 309. Larger grain sizes in the pro-delta region are related to higher energy turbidity currents crossing the delta. At distal sites, weak over-flows and inter-flows deposit fine-grained sediments. However, the mean grain size in KIN 311 is greater than KIN 309 (disregarding the drop-stones).

Down-core, grain sizes fluctuate greatly in KIN 310. However, there is no discernible pattern. In KIN 311, grain sizes are slightly coarser deeper in the core and fine upwards. 114

In nearby KIN 309, grain sizes are fine and show minimal variance down-core. These differences in mean and variance of grain size between the two distal cores complicate reconstruction of climate.

Loss-on-ignition results from both cores (1 hr at 550 °C) show a distinct down- lake and down-core pattern with some variance. The average loss-on-ignition for KIN

310 and KIN E4-06 is lower than the average in distal cores. This is because discharge crossing the Tatogga Lake - Kinaskan Lake delta is clastic rich and high-energy currents maintain fine organic matter in suspension. Down-lake, higher organic content also comes from greater water column contributions and inputs from tributaries. Down-core,

KIN 310 shows high variability in loss-on-ignition with a slight increase in organic content from couplet count 211 to the top of the core. In KIN 309 and KIN 311, organic content is higher and less variable than KIN 310. In the lower section of both distal vibra- cores, organic content is high and either stable or increasing at a fairly constant rate. At the 2.35 m mark in KIN 309 (-1388 years BP) and 2.55 m mark in KIN 311 (-1200 years

BP), organic content drops suddenly in both cores. Above this depth in KIN 309, loss-on- ignition stabilizes about a new, lower mean. Above this depth in KIN 311, organic content drops but then returns to pre-2.55 m levels by the top of the core. Differences between the loss-on-ignition results for the two distal sites are probably due to site- specific sedimentation patterns or events. However, these differences do not seem more significant than common trends between the two cores.

The final test of sediment composition involved XRF. Five samples were taken from KIN 311 and tested for elemental (oxide) composition. In each sample, SiC>2

(followed by AI2O3) was the major component. With depth in the core, composition only 115 varied slightly (i.e. minor decrease in SiC^) suggesting that there has not been major changes in the primary mix of sediments sources over the last 2 to 2.5 ka.

For ease of comparison between the Kinaskan Lake sedimentary record and records from nearby lakes, couplet counts and four radiocarbon dates were used to create a chronology of deposition. The three vibra-cores (KIN 310, KIN 309 and KIN 311) contain approximately 2,000 years of sediment history. At most, Ekman records represent several decades. Sediments from Skinny Lake, B.C., Pyramid Lake, B.C., Bowser Lake,

B.C., and Atlin Lake, B.C., contain longer-term records of sedimentation at a lower resolution than Kinaskan Lake (Spooner et al., 2002; Mazzucchi et al., 2003; Gilbert et al., 1997; Gilbert et al., 2006; Serink, 2004). However, the records of these cores provide a context to place Kinaskan Lake results. Furthermore, the section of these cores which corresponds with the Kinaskan Lake core (the late Holocene) shows a similar pattern of cooling and/or drying related to the onset of the Little Ice Age (~ last 1 ka) and recent warming and/or increase in precipitation.

7.1.4 Findings of this Study

Combining the results of this study; a record of past and modern hydrology and sedimentation in Kinaskan Lake, with published findings from nearby lakes, several important conclusions can be established.

(1) Clastic sedimentation has varied on spatial and temporal scales over the

Holocene. During the late Pleistocene and early Holocene, deglaciation of Northern

British Columbia watersheds caused mobilization of large volumes of coarse clastic sediments. Accumulation and sedimentation rates (and the relative importance and extent of underflows) increased in both the proximal and distal regions of glacier-fed lakes of 116

Northern British Columbia. Since the early Holocene, clastic sedimentation has decreased significantly, with relatively minor peaks and troughs related to short-term departures in climate conditions (temperature and/or precipitation). The greater distance between contributing glaciers and glacier-fed lakes has favored storage of coarser sediments in the watershed (i.e. in braided rivers and in up-stream lakes). Of the clastic sediments that do reach the lake, the largest volumes of sediment (and coarsest factions) are deposited in the proximal region whereas the smaller amounts (and finer components such as silts and clays) are flushed down-lake through over-flows and inter-flows.

(2) Inputs of clastic sediments from geomorphic events can mask fluxes in climate conditions. This type of event was responsible for creating a massive deposit of silt and drop-stones at the base of KIN 309. However, the composite sedimentary record (i.e. the sedimentation or accumulation rate + sediment structure + grain size distribution + organic content + geochemical make-up + records from other sites) can reveal climate conditions. Analysis and interpretation of distal sediments from Kinaskan Lake suggest a regional increase in temperature and/or precipitation ~2 ka BP (Neoglacial), a possible decrease in temperature and increase in precipitation between 1300 years BP and 1200 years BP (transition into the Little Ice Age), falling temperatures and decreased precipitation between 1200 years BP and 1000 years BP, an increase in temperature and precipitation from 650 years BP to 450 years BP, fluctuations in temperature and precipitation from 450 years BP to 250 years BP, and elevated temperatures and increased precipitation from 100 years BP to the present. In many instances, evidence of these episodes of climate change have been identified in the sedimentary record of nearby 117 lakes: Skinny Lake (B.C.), Pyramid Lake (B.C.), Atlin Lake (B.C.) and Bowser Lake

(B.C.), show several episodes of regional climate change.

(3) Lakes with similar characteristics to Kinaskan Lake appear to be "sensitive" to climate change because of the small amount of glacier ice cover and availability of easily eroded rocks. Alternatively, watersheds containing no ice-cover and highly resistant rock types risk exhaustion of their sediment supply. As the regional climate continues to warm it is expected that Kinaskan Lake will see more variable and increasing rates of sediment accumulation with higher organic accumulation. Repeat sampling in the future may help resolve any debate on using Kinaskan Lake as a gauge of climate change.

7.2 Future Directions

The outcome of this study demonstrates the valuable information which can be found in the sediments of glacier-fed lakes in British Columbia Cordillera. Studies of clastic sediment accumulation chronologies can provide high resolution evidence of changing contributions from large watersheds. To understand the changes presently occurring in the climate of mountainous regions of northwestern North American

(specifically northwestern British Columbia), more high-resolution and long-term studies of glacier-fed lake sediments should be conducted. It will be particularly important to expand the range of climates and geologies that these lakes occupy to get a full regional picture of climate change impacts. 118

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Winters, H.A. 1961. Landforms associated with stagnant glacial ice. The Professional Geographer 13 (4), 19-23. Start: 57° 40.688' N, 130° 05.8398' W Kinaskan Lake Acoustic Long Profile Images begin at proximal (North) end of lake

O O i At the delta, there is a thin layer of acoustically transparent material. This gives way to a fifteen meter, well-layered and conformable cap of mud on a hummocky substrate, with up to ten identifiable acoustic reflectors. a £ o S £ CL

Start: 57° 39.5259' N, 130° 07.1631'W

The mud cap thins to ten meters with five acoustic reflectors before being interrupted by an in-filled basin. Afterwards, the cap expands to twenty meters with ten acoustic reflectors. Start: 57° 38.1138' N, 130° 07.1553'

Kinaskan Lake Acoustic Long Profile Conftnued

The acoustic signal is muffled suggesting turbid water along the lake bed. When the signal clears, an eighteen meter thick cap of stratified mud is evident. This cap expands where it overlays a shallow basin which has been filledwit h acoustically transparent, unconformable mud.

Start: 57° 36.7671' N, 130° 07.1523'

7SI . i "> * s I s 1 « * a Lake bottom mounds culminate in a sill which is capped by five meters of well-layered mud. There are five visible acoustic reflectors. Further South, the lake bottom is irregular and leads to an in-filled basin. Start: 57° 35.5058' N, 130° 08.5127' Kinaskan Lake Acoustic Long Profile Confirmed

:_--nm, . .. -y- "•"— ». »• ™ "• , 30 "" „, " j • - - -w - ». „ «. **• ,. ». - — 4 "" «*->" ' •»4- ir I Si - 4- » 1 m • - - •„-,- • _ "5 1- —5- — 5- - • * - « 5 ™ir —5— ]•_"„-.•-

I • ~% i i "I "r"' i 1 S - f ' i - I i This section contains two in-filled basins and a thinly capped miniature sill. The basins contain well-layered, unconformable mud. The sill has five meters of stratified mud with at least seven faintly visible acoustic reflectors.

Start: 57° 34.4609' N, 130° 09.6494'

...} s-*j 4* - '

A twelve meter thick cap of stratified, conformable mud covers a series of acoustic parabolas. The mud has at least six acoustic reflectors. Further South, the number of parabolas increases and the mud layer thins. Start: 57° 33.4380' N, 130° 10.6660' W Kinaskan Lake Acoustic Long Profile Images terminate at distal (South) end of lake

—l.f~- i ? 3- I _____ 1 120 - 3 *S q ? -t : ? "J """""—9 ™s """:• .-y.... . -3- - _ j^_ A series of hummocks composed of eight meter thick, well-stratified mud with at least six acoustic reflectors, are underlain by faint parabolic signals. This suggests that below the mud are multiple boulders.

Start: 57° 32.3672' N, 130° 11.7021' W

Towards the South end of the lake, the water depth shallows and the cap of mud thins to several meters. The mud is still stratified, but is dominated by fewer acoustic reflectors.

U> Kinaskan Lake Acoustic Profiles

Transect 1: Farthest distal transect of lake, West

Section A: Mud cap is smooth at the topo f the slope. Two strong reflectors dominate the upper mud surface and two dominate the lower boundary. Section B: Hummocks composed of well-layered, conformable mud compose the middle of the basin. Mounds have a mud cap of three meters, composed of four reflectors. Troughs have five meters of mud with six reflectors. Section C: Six acoustic reflectors, with a combined thickness of six meters, dominate. They are well-layered and conformable. There is a small trough at the top of the east slope. Transect 2: West East

Section A: A cap of five meter thick, well-layered and conformable mud with two acoustic reflectors dominating the upper and lower surfaces. Section B: At the nadir of the trough the mud layer becomes increasingly hummocky. A series of acoustic parabolas suggest the presence of boulders. Transect 3: West East

Section A: The west slope is smooth and capped by a thin layer of acoustically transparent mud. Hummocks, parabolic reflectors, and a trough composed of unconformable, well-layered mud are evident in the middle of the basin. Section B: At the slope crest, five acoustic reflectors dominate the signal for four meters of mud.

Transect 4: West East

Section A: Slope is covered by a fivemete r thick layer of mud which is well-layered with eight reflectors. Section B: The basin is hummocky with several in-filled troughs. Parabolic reflectors at the base of the eastern shore suggests the presence of boulders. Mud thickness ranges from eight to twenty meters. Section C: The mud layer thins from three to one meter towards the slope crest. The number of acoustic reflectors also decreases from five to two. Transect 5: West East

I -\~ZA \ ,M: j J 4 _. I B _J_ Section A: The slope is capped by three meters of acoustically transparent mud. Section B: The basin has a smooth upper surface which masks several in-filled troughs. Mud thickness ranges from two meters on the slope, with two acoustic reflectors, to twenty meters in the troughs, with four to nine reflectors. Mud in the troughs is not conformable to the lowest reflectors.

Transect 6: West East

i - - - »i

»- » „_„ -* •*k^ """'- ~- „ _— *. ^\ ;• " .. - -^^ : \- - — „. „ „. ^~ X .. - ... — ^&tifr* ,. 1 s. .. - . "-,-^M™ ' \ i "_ •° • ~jftfPr t •>- T »*". — r " -t .. ~ __JHBBP"* ~^ % t ' i. -».»-- t_ . Mj|Hp~; % \

Section A: Up-slope there is a thin layer of smooth mud, this gives way to hummocks of mud three to five meters thick. Section B: Two troughs are in-filled with approximately twenty meters of unconformable, well-layered mud. Between, there is an area of raised lake-bed with parabolic signals suggesting the presence of boulders. Section C: The lower portion of the slope is covered in five meters of mud, this thins to one meter at the top.

Transect 8 : West

••* •• c Section A: The slope is capped by five to eight meters of well-layered mud. Section B: A trough is in-filled by fifteen to twenty meters of unconformable, well-layered mud. Section C: The lower portion of the slope is has a five meter thick cap of mud. Towards the top of the slope, this thins to about two meters. Transect 9 : West East

^,1

• \ - „m-vgG^^^0* * : ; . \ «riM^ - t * - ^1^-^i * 1 •ilif ^^m^.". -. ;•--"-*-• . ^ " A • • • : I • - ; *B- : Section A: A five meter layer of well-stratified mud tops the slope. At the base of the slope, there is an in-filled basin with fifteen meters of unconformable, well-layered mud. Section B: The slope is capped by five meters of hummocky, well-stratified mud.

Transect 10: West

Section A: The steep slope is covered in five meters of mud with weak reflectors. Section B: The base of the slope grades into an in-filled basin with up to thirty meters of well-layered, unconformable mud. The mud thins to five meters on the eastern slope.

as Transect 11: West East

Section A: West slope is characterized by a series of hummocks, capped by a five-meter thick layer of weakly stratified mud. At the base of the slope is an in-filled basin with twenty meters of unconformable mud. Section B: The slope is smooth and capped by five to two meters of layered mud.

Transect 12: West East

Section A: The slope is capped by a layer of mud with varying thickness. At the base of the slope, there is an in-filled basin. After the basin, a series of hummocks are covered in a thick layer of stratified mud. Section B: The second basin is also in-filled with unconformable, well-layered mud. The mud is approximately twenty-five meters thick. On the lower part of the slope there is approximately ten meters of well-layered mud. This thins to five meters at the slope crest. Transect 13: West East

Section A: A series of hummocks composed of conformable, well-layered mud of varying thickness. Section B: Abasin in-filled with ten to twenty meters of unconformable, well-layered mud with up to seven reflectors. Section C: A two to five meter mud cap with five acoustically weak reflectors positioned atop the east slope.

Transect 14: West East

m~~ r—

'1 iy^mmfmmpjsp*iqig0£ffh?.

'" — - "i -

Section A: Abasin in-filled with fifteen meters of unconformable, acoustically transparent mud. Section B: Hummocks consisting of five to fifteen meter thick, conformable and well-layered mud with up to ten reflectors. Transect 15: West East .. • ; tf* J* *••**- - - - — -

k "• A • •'- B • Section A: Athin layer of poorly layered mud which gives way to a series of hummocks composed of twenty meter thick, conformable mud with at least nine reflectors. Section B: At the bottom of the slope, mud is acoustically transparent. The upper half is capped by a five meter thick layer of well stratified and conformable mud with at least four acoustic reflectors.

Transect 16: West East

Section A: Atwenty meter thick cap of well-layered and conformable mud, with at least nine reflectors, covers the west slope. Section B: In the basin, only the upper five meters of in-filling has visible structure. The slope is capped with a four meter layer of stratified mud with acoustically weak reflectors. Transect 17: West East

SectionA: Athick cap of well-layered and conformable covers the slope, thinning towards the base. SectionB: Aseries of parabolic reflectors dominates the surface of the eastern slope. This suggests local armoring by boulders. 141

Appendix B Graph Comparing Malvem and Coulter Particle Size Results

Comparing Malvern and Coulter Particle Size Data

-KIN309D16- Coulter Trial 1 -KIN309D16- Coulter Trial 2 -KIN309D16- Coulter Trial 3 -KIN309D19- Coulter Trial 1 -KIN309D19- Coulter Trial 2 -KIN309D19- Coulter Trial 3 -KIN309D20- Coulter Trial 1 -KIN309D20- Coulter Trial 2 -KIN309D20- Coulter Trial 3 -KIN309D16- Malvern Trial 1 -KTN309D16- Malvern Trial 2 KIN309D16- Malvem Trial 3 -KIN309D19- Malvern Trial 1 KIN309D19- Malvern Trial 2 -KIN309D19- Malvern Trial 3 K1N309D20- Malvern Trial 1 -KIN309D20- Malvern Trial 2 - KIN309D20 • Malvern Trial 3

& ^ ^ f / £ tf '

Grain Size Classes (urn) Appendix C KIN 309, KIN 310, KIN 311 and Ekman Stratigraphic Logs and Sediment Structure and Composition Graphs

" Kinaskan 309 Stratigraphic Log Loss-on-ignition Grain Size Distribution Laminae or Bed Thickness % Lost after 1 hr at 550 C % per grain class (mm) 0 j H ? IP 1,2 0 2040 60 80 100 _g 22 44. 0

V \

U \

\ \

Clast

I Massive ! Disturbed

Fine Parallel Beds Thick Parallel Beds •ti- ToTi 0 20 40 60 80 100 0 22 44 143

M| Kinaskan 310 Stratigraphic Log Loss-On-Ignition Grain Size Distribution Laminae or Bed Thickness % Lost after 1 hr at 550 C % per grain class (mm) ,0 2 4 6 8 10 12 0 2040 60 80 100 J) 22 _44 1 i i y i—l V " T T .*==•*

V "">

^ i Fine Parallel Beds Thick Parallel Beds

I i !•.,[,. 0 2 4 6 8 10 12 0 20 40 60 80 100 0 22 44 144

kinaskan 311 Stratigraphic Log Loss-On-[gnition Grain Size Distribution Laminae or Bed Thickness % Lost after Ihr at 550 C % per grain class (mm) 0 20 40 60 80100 0 22 44,

I

iFine Parallel Beds •.••.I I Thick Parallel Beds

0 2 4 6 8 10 12 6% 40 60 80 100 6~~ ~~TT ~M 145

Loss-On-Ignition, Percent Lost Kinaskan 2 4 6 8 10 12 2006 Ekmans

El-06 g 4 -E1.06(Loss»fteil h.at550C) 6. -El-06(Lossaft«l hiatllOOC) E2-06(Lossafteil hi at 550 C) Q E2-06 (Loss aftei 1 hi at 1100 C) - E3-0C (Loss after 1 hi at 550 C) 30^ 50 -E3-06 (Lots after! hi at 1100 C) -E4-06 (Loss aftei 1 hiatSSOC) 0 -E4-06 (Loss aftei I hiatllOOC)

E2-06 60 •

Laminae or Bed Thickness (mm) 2 4 6 8 10 12 14 16 50 0

S 40 E3-06

•3 p.

D 50 80 0

E4-06

•3 a.

50 ] Fine Parallel Beds

1 Thick Parallel Beds