THE GLACIAL AND PERIGLACIAL'GEOMORPHOLOGY OF
THE FOURTH OF JULY CREEK VALLEY, ATLIN REGION,
CASSIAR DISTRICT ,‘ NORTHWESTERN BRITISH COLUMBIA

Thesis for the Degree of Ph. D‘
MICHIGAN STATE UNIVERSITY
ANN M. TALLMAN
1975

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This is to certify that the

thesis entitled

Thu GLACIAL AND FERIGLACIAL GEONORPHOLOGY

OF THE FOURTH OF JULY CREEK VALLEY, ATLIN REGION

CASSIAR DISTRICT, NORTHWESTERN BRITISH COLUIV'BIA
presented by

ANN M. TALLMAN

has been accepted towards fulfillment
of the requirements for

_Bh_-_D_¢___ degree 1n m

Date

Major professor
OC‘I'. 30 , I975

0-7639

ABSTRACT

THE GLACIAL AND PERIGLACIAL GEOMORPHOLOGY OF THE
FOURTH OFIJULY CREEK VALLEY, ATLIN REGION, CASSIAR
DISTRICT, NORTHWESTERN BRITISH COLUMBIA

By
Ann M. Tallman

The Atlin Region of northern British Columbia lies at
the northern end of the Wisconsinan Cordilleran Ice-sheet
in an area affected by Wisconsinan glaciation from distinctly
different provenances. The prime nourishment areas were the
Coast Mountains and the adjacent Yukon and Stikine Plateaus
in Northern British Columbia. In late stages of deglacia-
tion during warming climatic conditions local ice centers
contributed their share.

The study site, the Fourth oflJuly Creek Valley, is
situated in an interlobate region where ice from the Cassiar
Range and the Boundary Range in out-of—phase fluctuations
modified the landscape. Morpho-stratigraphic evidences re-
veal multiple glaciation sequences identifiable back to the
early Wisconsinan. Wisconsinan glaciation was very extensive
in the Atlin region and it's earliest phases covered all but
a few of the highest ridges and peaks in the region. Evidence

for pre-Wisconsinan glaciation is by inference from erosional

Ann M. Tallman

remnants on the high ridges of the Boundary Range and from
the presence of early drifts well outside of the study area.

A Wisconsinan chronology is presented in which an
early Wisconsinan phase and middle Wisconsinan (40,000 1
years B. P.) intraglacial are invoked, followed by a middle
to late Wisconsinan sequence of two main stages and three
lesser ones with intervening intraglacials. These are
termed Atlin I, II, III, IV, and V for ice from the Boundary
Range and Gladys I, II, III, and IV for ice from the Cassiar
source.

The evidence is based on glacial stage positions and an
array of related depositional sequences of terminal, lateral
and ablation moraines, crevasse fillings, eskers, kames,
outwahs trains, glacial-fluvial terraces and drainage pat-
terns resulting from the down-wasting of ice from both pro-
venances. The relative ages of weathering are discussed and
inferences attempted with respect to intraglacial conditions.
A tentative correlation is made with the glacial strati-
graphy and chronologies reported from adjoining regions.
Special attention is given to the chronology of the Juneau
Icefield and Taku District in Alaska which was the major
accumulation area for the Boundary Range ice. In addition,
a broader correlation is suggested with the well-developed
North.American mid-continent chron010gy.

The evidence presented suggests that global climatic
events which led to the major Pleistocene fluctuations,

though somewhat out-of-phase, are evidenced in the study

Ann M. Tallman

area. Where there are significant differences they are
considered relative to determining geographical and oro—
graphical factors. It is shown that minor fluctuation in
mid-continent glaciations do not always appear in this
Cordilleran chronology because the high latitude (sub-
Arctic) and mountainous terrain tends to produce glacia-
‘tions earlier than in mid-continental areas, and to retain
them for longer periods during intraglacial stages of

climatic amelioration.

THE GLACIAL AND PERIGLACIAL GEOMORPHOLOGY OF THE
FOURTH OF JULY CREEK VALLEY, ATLIN REGION, CASSIAR

DISTRICT, NORTHWESTERN BRITISH COLUMBIA

By

\
Ann MD‘Tanman

A THESIS

Submitted to
Michigan State University
in partial fulfillment of the requirements
for the degree of

Doctor of PhilOSOphy

Department of GeolOgy

1975

ACKNOWLEDGMENTS

I would like to express my deep appreciation to Dr.
Maynard Miller for his enthusiastic support throughout this
research project. His encouragement, arrangements for
logistic and financial support and scientific interest and
expertise in the problem made the study possible. To the
members of my committee, Dr. Jane Elliot, from whom I had
my first geolOgy course, Dr. Chilton Prouty, and Dr. E. P.
Whiteside, thank you for critically reading the manuscript
and making valuable suggestions. A special thanks also to
Dr. William Hinze and Dr. Sam Upchurch for their guidance in
the early stages of this endeavor.

Valuable guidance in the field was given by Dr. R. F.
B1adk, Dr. R. L. Nichols and many other Foundation for
Glacier and Environmental Research affiliates. Members of
this group also served as field assistants during much of
the field work. A special recognition is made to Christine
Dilts, Sandra Stewart, Jennie Tallman, and Vicki Pedone for
their help and company in the field. Mrs. Joan W. Miller
served not only as an invaluable lOgistition but a good

friend with a seemingly unlimited supply of moral support.

ii

Financial support was secured through Dr. Maynard
Miller and the Foundation for Glacier and Environmental
Research in grants from the National GeOgraphic Society,
the Army Research Office, the National Science Foundation
and the Readers Digest. Tables II and V and Figure 2 were
drafted by the National GeOgraphic Society.

A final thanks to Smith College and the students in

my classes for their patience as this thesis was being

completed.

Chapter

I.

II.

III.

IV.

VI.

TABLE OF CONTENTS

Page
LOCATION AND PHYSIOGRAPHICAL CHARACTER
OF THE STUDY AREA 0 O O O O O . O O C O O O C 1
IntrOdUCtion . C O O O O O O I O C O C O C 1
Geographical setting . . . . . . . . . . . 3
GEOLOGICAL] FRMEWORK O O O I I O O O C O O O O 7
Previous research and allied studies . . . 7
Main litholOgic and structural elements . . 8
surfiCial geOIOgY I I O I O C O I O O O O O 11
REGIONAL CLIMATIC PARAMETERS . . . . . . . . . 12
IntrOduction O O O O O I O O O C O I O O O 12
Available weather record . . . . . . . . . 12
Temperature and precipitation records . . . l3
Climatic variations and secular shifts
in regional storm paths . . . . . . . . . 15
SOILS AND VEGETATION RELATIONSHIPS . . . . . . 19
Introduction . . . . . . . . . . . . . . . 19
Variations in soil maturity . . . . . . . . 20
Regional comparisons . . . . . . . . . . . 20
Soils-vegetation characteristics with age . 22
PRE -WI SC ONSINAN GLACIATI ON 0 o o o o o o o o o 25
IntrOduCti on O O O O O O O O O O O O O O C 25
Morphoqenetic phrases . . . . . . . . . . . 25
Ancient ice centers and outer limits
Of g1 aCiation C O O O C C O O O I I O C O 26
WISCONSINAN GLACIATION . . . . . . . . . . . . 32
IntrOduCtion O O O O O O O O O O O O O O O 32
Gladys Ice moraines . . . . . . . . . . . . 39
Atlin Ice moraines . . . . . . . . . . . . 59
Local cirque activity and tandem cirque
levels 0 O O O O O O O O O O O O O I O O 71

iv

Chapter

VII.

VIII.

IX.

History of glacio-fluvial terraces in
the Porter Lake Valley . . . . . . .

Moraines and glacio-fluvial terraces in

Fourth of July Creek Valley . . . .
Esker-kame complex in the Fourth of
July Creek Valley . . . . . . . . .

Valley assymetry and abandoned channels

in the Gladys Lake depression . . .

PERIGLACIAL ENVIRONMENT . . . . . . . . .
Solifluction zones . . . . . . . . . .
Nivati on h01lows O I O O O I O O O O O
Palsas and related frost mounds . . .
Character of Atlin palsas . . . . .

Electrical resistivity assessment of
the Fourth of July Creek palsas . .
Details of the resistivity method and

computational technique . . . . . .
Interpretation of the geophysical data
on SUb-Surface ice 0 o o o o o o o 0

Significance of the Atlin palsas and
related considerations . . . . . . .

A WISCONSINAN CHRONOLOGY . . . . . . . . .
Keys to the interpretation . . . . . .
Pre-Atlin I glaciation . . . . . . . .
Boulder Creek Intraglacial . . . . . .
Atlin I - Gladys I glaciation . . . .
A suggested Pine Creek Intraglacial .
Atlin II - Gladys II glaciation . . .
A late Wisconsinan climatic

amelioration . . . . . . . . . . . .
Atlin III, IV, and V - Gladys II and
V . . . . . . . . . . . . . . . . .
HOLOCENE . . . . . . . . . . . . . . .

THE

The last 10,000 to 11,000 years . . .
Quaternary peat deposits in Boulder
Creek valley 0 O I C O C O O O O O O

Page

82
96
99

101

106

106

107

111

115

117

120

125

131

133

133

135

139

140

144

144

147

147

152

152

154

Chapter Page

X. SOME SUMMARIZING CONCLUSIONS . . . . . . . . . 159

Problems of teleconnection . . . . . . . . 159
Conceptual model for out-of-phase
Pleistocene glacial variations

of large magnitude . . . . . . . . . . . 163
XI. FURTHER RESEARCH PROSPECTS AND ADDITIONAL
STUDIES . . . . . . . . . . . . . . . . . . 166
APPENDIX
Glossary . . . . . . . . . . . . . . . . . 169
BIBLIOGRAPHY . . . . . . . . . . . . . . . . . 170

vi

LIST OF TABLES

Table Page

I. ClimatolOgical data summary for stations

in the Atlin region . . . . . . . . . . . . . 14
II. Holocene chronology for Atlin and Taku
districts 0 I O O I C O O O O C O O O O O I O 18
III. Comparison of cirque elevations in the Atlin 76 &
region and suggested regional correlation . . 77
IV. Fluvial terrace sequence in Porter Lake 86 &
Valley - A o o o o o o o o o o o o o o o o o 87
IV. Fluvial terrace sequence in Porter Lake 88 &
valley - B o o o o o o o o 0 o o o o o o o o 89
V. Tentative Wisconsinan chron010gy . . . . . . . In poCket

VI. Late-Pleistocene glacio-chronologic summary
for the Atlin region and Taku district. . . . 158

vii

Figure

1.

5a.

5b.

10.
11.
12.
13.
14.
15.
16.

17.

LIST OF FIGURES

Southeastern Alaska, northwestern British
Columbia and southwestern Yukon
Territory 0 O I O O I O O O O O O O O O O

Atlin region 0 O O O O O I I O O O O O O O

Atlin 104N 1:250,000 quadrangle . . . . .

Major divisions of Fourth of July Creek

valley . . . . . . . . . . . . . . . . .
Maximum extent of Gladys ice . . . . . . .
Maximum extent of Atlin ice . . . . . . . .
Mount Ewing . . . . . . . . . . . . . . . .

Fluted drift north of Black Mountain . . .

Erratics on Gladys Mountain . . . . . . .

Bedrock incisions on eastern flank of Mount
Barham O C O C O O O C U . C O C O O C .

Marble Dome . . . . . . . . . . . . . . . .
Intermediate Fourth of July Creek Valley .
Kettle . . . . . . . . . . . . . . . .i. .
Gladys Ice moraines near Gladys Mountain .
Underfit glacial run-off streams . . . . .
Atlin Ice and Gladys Ice run-off channel .
Atlin III moraine complex . . . . . . . . .

Aerial stereo pair of Fourth of July Creek
Valley with Porter Lake Valley . . . . .

viii

Page

In pocket

36
37
38
41
43
43

47
47
50
52
55
58
58
65

67

Figure Page

18. Mouth of the Fourth of July Creek . . . . . . . 73
19. Sketch map of upper Fourth of July Creek

Valley . . . . . . . . . . . . . . . . . . . 81
20. Porter Lake terraces . . ... . . . . . . . . . 84
21. Intermediate Fourth of July Creek transect . . 90
22. Lower-upper Fourth of July Creek transect . . . 98
23. Head of Gladys Lake . . . . . . . . . . . . . . 104
24. Nivation hollows on Mount Vaughn . . . . . . . 109

25. Stone circles in esker complex of Fourth of
July Creek Valley . . . . . . . . . . . . . . 109

26. Palsas in upper Fourth of July Creek Valley . . 114

27. Eskers and peat plateau in upper Fourth of

July Creek Valley . . . . . . . . . . . . . . 114
28. ~Wenner Array . . . . . . . . . . . . . . . . . 121
29. Lee Partition . . . . . . . . . . . . . . . . . 121
30. Map of resistivity profiles . . . . . . . . . . 123‘
31. Apparent resistivity, profiles 1 and 2 . . . . 124
32. Apparent resistivity, profiles 3 and 6 . . . . 124
33. Apparent resistivity, profiles 4 and 5 . . . . 126
34. Resistivity, profiles 1 and 6 . . . . . . . . . 126
35. Resistivity, profiles 1 and 3 . . . . . . . . . 129
36. Resistivity, profiles 4 and 5 . . . . . . . . . 129

37. Organic horizons in lower Boulder Creek
Valley I I I I I I I I I I I I I I I I I I I 155

ix

CHAPTER I

LOCATION AND PHYSIOGRAPHIC CHARACTER
OF THE STUDY AREA

W

Studies of the geomorphOIOgy and glacial history of
the northern Cordilleran region in British Columbia and the
Yukon have been made primarily in connection with the Geo-
IOgical Survey of Canada's natural resources and geology
research program (Cairnes, 1913; Kerr, 1934; 1936; Aitken,
1959). Other workers have availed themselves of the lOgistic
advantage of the Alaskan Highway (Fig. 1) to conduct recon-
naissance studies of the glacial history along geographically
limited sectors of this transect to the north and east of the
Atlin region (e.g. Denny, 1952).

The present dissertation concerns such an investigation
in a specific locale within this large region where only the
most general aSpects of Pleistocene glacial stratigraphy have
been outlined hJohnston, 1926; Watson and Mathews, 1944; Arm-
strong and Tipper, 1948; Denny, 1952; Wheeler, 1961; Mulligan,
1963; Miller, 1956, 1963, 1964a and b, 1973; Anderson, 1970;

Miller and Anderson, 1974).

 

 

 

 

 

   
   

    

 

 

 

     
  
 
  

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Figure 1. Southeastern Alaska, northwestern British
Columbia and southwestern Yukon Territory

3

GeOgraphical Setting

 

The study area is in the Atlin sector of the Cassier
District in northwestern British Columbia close to the Yukon
border (Fig. 2), an area included as part of a physiOgraphic
region termed the Canadian Cordilleran Province by Bostock
(1948). The major part of the Atlin area lies within the
Tagish Highland and the Teslin Plateau. TOgether these con-
stitute the south central portion of the Yukon Plateau of the
Interior System (Holland, 1964). The southwestern portion of
the Atlin map area reaches into the Western Physiographic
System and the Northern Boundary Range of the Coast Mountains,
north of the Taku River (Bostock, 0p. cit.). The Southern
Boundary Range, with certain geolOgic and physioqraphic sim-
ilarities, is considered to extend southward from the Taku
River to the area of the Stikine River. Both the Taku and
the Stikine Rivers follow antecedent drainage lines across
the Coast Mountains to the Gulf of Alaska (Fig. 1).

Several large lakes in the Atlin region play a signif-
icant role in the interpretation of the glacial history of
the area. Atlin Lake (elev. 2182 feet, 665 meters), the
largest natural lake in British Columbia, extends from the
southwestern portion of the Atlin map area and continues in
a northerly direction for 65 miles (104 kilometers) with its
northern 10 miles (6 kilometers) lying in the Yukon Territory
(Fig. 3, in pocket). The lake is positioned in a structural
as well as physiographic low and has been significantly

modified by continental-type glacial lobes which have

 

Figure 2. Atlin Region

5
originated in the Northern Boundary Range in the vicinity
of the present-day Juneau Icefield (Fig. 1). Teslin Lake
(elev. 2239 feet, 683 meters) and Gladys Lake (elev. 2915
feet, 888 meters) are part of this interior drainage system
trending in a north-northwest direction (Fig. 3). Each of
these linear depressions was the locus of Pleistocene ice
masses moving in a northerly direction from the Stikine
Plateau.

Of local significance to the geomorphic interpreta-
tions are several smaller water bodies, including Surprise
Lake, (elev. 2894 feet, 882 meters)(Fig. 3) the depressions
for which were largely scoured by Atlin and Teslin ice or
otherwise affected by glaciofluvial drainage from these ice
masses. The morpholOgy and character of these lakes also
reveal that their basins have suffered considerable modifica-
tion through the repeated and out-of-phase glaciations re-
lating to several local ice centers of alpine character.

Fourth of July Creek, which flows through the study
area, at present drains in a southwesterly direction with
its headwaters area located at 1330 20' West Longitude and
590 42' North Latitude. The watershed of this stream
initiates in a zone between 3000 and 6000 feet (915 to 1830
meters) elevation and extends to Atlin Lake at 133° 43' West
Longitude and 49° 42' North Latitude. This part of the study
area also includes the drainage basin of Consolation Creek,
which lies northeast of Fourth of.July Creek and includes
its outlet to Fish Lake 600 55' North Latitude and 133° 10'

West Longitude (Fig. 3).

6

The Fourth of July Creek Valley is flanked on the
southeast by an unnamed 20 mile (32 kilometers) long sub-
sidiary range, the upper ridges of which attain elevations
of 4500 to 6800 feet (1370 to 2075 meters). The major peaks
on the ridge line of this range include Mount Leonard (elev.
6000+ feet, 1829 meters), Mount Vaughan (elev. 6000+ feet,
1829 meters) and Mount Barham (elev. 6808 feet, 2075 meters).
Mount Ewing (elev. 5410 feet, 1649 meters) is the highest
crest on the valley's northwest flank and is part of a 12-
mile (19 kilometers) long ridge which separates Fourth of
July Creek Valley from the valley leading to Porter Lake
(Fig. 3).

CHAPTER II
GEOLOGICAL FRAMEWORK
Previous Research and Allied Studies

Maps of the Canadian National T0p0graphical series
cover the Atlin District at a scale of l:250,000. Only a
small part of the western portion, including some of the
Atlin Lake sector in the vicinity of the unincorporated
village of Atlin, has been photogrammetrically mapped at a
scale of 1:50,000. Over sixty years ago, a reconnaissance
report of the geology along portions of the Atlin Lake
sector was compiled by D. D. Cairnes (1913). Little else,
except local studies and brief geoloqical descriptions of
mining sites (e.g., Cockfield, W. E., 1925 and Gwillim, I. C.,
1901, 1902) was published on this area until the rather re-
cent geological map and report of Aitken (1955, 1959) the
field work for which was supported by the Geological Survey
of Canada. Part of Aitken's research has also been pre-
sented in his Ph.D. thesis (1953). Other recent maps and
geological descriptions are available for the Lake Bennett
sector, west of Tagish Lake (Christie, 1957), for the White-
horse area in the adjoining Yukon Territory (Wheeler, 1961).
the Teslin area to the north (Mulligan, 1963), Wolf Lake, to

the northeast (Poole, 1955) Jennings River to the West

7

8
(Gabrielse, 1969), and aSpects of vulcanological history of
the region to the south (Souther, 1968, 1972). A compila-
tion of these and other local studies is presented by
Douglas (1970). Also to be mentioned is the glacial
geological and regional geology report by Miller (1956,
1963) relative to the Taku River sector and theIJuneau Ice-
field south of the study area, and that of Gilkey (1951)
and Forbes (1959) covering the petrolOgy and structure of
the central icefield nunataks and key lithologies found on
a transect across the icefield from.Juneau to Devil's Paw

(8504 feet, 2590 meters).

Main Lithologic and Structural Elements

A review of these reports reveals that three main
rock groups dominate the Atlin map area, especially to the
south and east. These are the Cache Creek series which is
Pennsylvanian and/or Permian in age; the Coast Range
batholithic intrusions of Mesozoic and early Tertiary time
and the subsequent middle to late-Tertiary volcanic and
early-Quaternary volcanic sequences (Aitken, 1959). The
northeastern part of the map area contains much limestone
and metasediments in the Cache Creek group (Aitken, 1959;
Mulligan, 1963). These include cherts, argillites, lime-
stone, greywacke, sandstone and silt-stone. Over the rest
of the map area east of Atlin Lake there are predominant
exposures of the Cache Creek sequence with more recent

intrusives including some Holocene volcanics.

9

Quartz diorite and diorite intrusions in the Mount
MCMaster, Mount Llangorse, and Hayes Peak (Fig. 3) sectors
are considered by Aitken (1959) to be of the same general
age as the coarse-grained igneous stock which has been
termed the Fourth of July Creek body. All of these intru—
sions are considerably unrooted and by comparison of the
‘petrology as well as by structural similarities, they are
considered to be of the same age.

A later Mesozoic acidic intrusion came in adjacent to
the Fourth of.July body and has been referred to by Aitken
(1959) as an Alaskite stock which extends on the surface
from Surprise Lake to Trout Lake, just south of Gladys Lake.
West of Trout Lake is the Snowden Range (Fig. 3) where
Alaskite litholOgy occurs at a location beyond a geomorpho-
logically and lithologically identifiable fault zone found
to the east to Trout Lake (Aitken, 1959). A quartz
monzonite on Dawson Peak (Fig. 3) is assumed to be the same
age as the Alaskite noted above, the criteria being com-
position, texture and structure. Kerr (1948) has traced a
belt of similar intrusives from Surprise Lake to the Taku
River map area and from there to the Stikine River region
where the intrusion is dated as late-Lower or early-Upper
Cretaceous in age. This is quite comparable to the
structural and lithologic characteristics in the Stikine
area described by Buddington (1927) on the contimental flank
of the Southern Boundary Range. Thus revealed is the con-

siderable linear extent of this geological province.

10

'As connoted above, the most predominant intrusives
are the Mesozoic Coast Mountain granitic intrusions of
batholitic and stock prOportions, with associated lithol-
ogies affected by permeating metasomatism from these acidic
emplacements. By K/Ar dating these granodiorities are
known to be some 50 x 106 years B.P. (A. Ford, US
GeOIOgical Survey, personal communication). This granite
is particularly abundant in the southern half of the map
area adjacent to the higher Coast Mountains, with some out-
liers in the specific area of this study. One such outlier
is the Fourth of July Creek stock or batholithic window
which has been referred to as the Fourth of July Creek body.
This body is composed of granodiorite and quartz monzonite.
It occupies more than 300 square miles (777 square kilo-
meters) in the western sector of the study region. On the
southwestern edge of this part of the Coast Range batholith
and adjacent to Atlin Lake, this unit occurs as a pink
granite, quite porphyritic and with some unusually large
phenocrysts.

The southwestern corner of the Atlin map area contains
a thick group of predominantly volcanic rocks, the SlOko
group. This is an especially unique lithologic assemblage
as it includes pyroclastics (dacite and rhyolite) with sub-
ordinate andesite and basalt, all segments of which have been
well described by Aitken (1959) and Douglas (1970). Other
smaller intrusions of less areal importance are also dis-

cussed by Aitken (1959) and Souther (1972). Some of these,

11
including dolerites and trap dikes, have been reported in
abundance on the.Juneau Icefield by Miller (1956, 1963)

and Forbes (1959).

Surficial Geology

As for the surficial geology, only brief overviews are
suggested in many of the reports noted above, with prac-
tically nothing on the Quaternary deposits of the Atlin area
per se. It is clear, however, that unconsolidated Quaternary
drift covers more than one-half of the map area of this study
and thinly mantles the entire Atlin region. Chronosequences
in this drift will be of primary consideration in the glacial

geological discussions to follow.

CHAPTER III
REGIONAL CLIMATIC PARAMETERS
Introduction

The Atlin district has a continental type climate with
some maritime influence from the Alaskan Panhandle partly in
consequence of the presence of passes and valleys extending
well inland through large sections of the Coast Mountains.
The rain shadow effect of the Northern Boundary Range also
.produces low precipitation on the interior flank of the main
axis of the Boundary Range. It should be noted, however,
that orographic influences in the interior highlands also
cause local increases in precipitation. Though the winters
are generally long and cold, temperatures are not as extreme
as in the more interior areas of Yukon Territory and north-
eastern British Columbia. ‘The summers are short and rela-

tively warm.
Available Weather Records

The Canadian Department of TranSport standard synOptic
weather records are available for the village of Atlin from
1905 to 1946, with simplified daily observations carried
forward by the Government Agent of the B.C. Provincial

Government since 1947 (Department of Agriculture,.British

12

13

Columbian Provincial Government, 1975). Year round three-
hourly observations have been taken at the Sub-arctic Re-
search Station of the Foundation for Glacier and Environ-
mental Research in Atlin, B.C., beginning in the summer of
1968. At Carcross, 55 miles (90 kilometers) northwest of
Atlin, data have also been recorded from 1907 to 1946, as
they have since 1941 in Teslin 60 miles (96 kilometers)
northeast. Since 1943 continuous synOptic weather observa-
tions have been taken in Whitehorse on the Yukon River 120
miles (183 kilometers) north—northwest of Atlin. Anderson
(1970) has tabulated weather data for the Atlin region,

including Carcross, Whitehorse, Atlin and Teslin (Table 1).

Temperature and Precipitation Ranges

The mean annual temperature for the village of Atlin is
32°F (0°C) compared to 31°F (-6°C) at Whitehorse and 30°F
(-1°C) at Teslin. The range is —54°F to 87°F (-47°C to
31°C) with an average of 85 frost-free days. Total mean
annual precipitation averages 11 inches (28 centimeters),
with 53 inches (135 centimeters) of snowfall (Porsild, 1951;
Kendrew and Kerr, 1955). Seasonal observations by Aitken
(1959) and by the writer in the summers of 1969 through
1972, suggest that the climate of the low relief area around
the Atlin town site is drier than the surrounding highlands,
again reflecting the pronounced effect of tOpOgraphy on
precipitation. Temperatures are also lower in the surround-

ing mountain and higher valley areas. Frost has often been

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15
observed in July and August at elevations of 3500 feet
,(1067 meters) and above. Snowfall is also common from
June to August at and above 6000 feet (1829 meters). It
is apparent that much of the weather is strongly controlled
by the tOpOgraphy.
Climatic Variations angTSecular Shifts
. in Regional Storm Paths

Of regional significance is a long-term climate-glacier
study done by Miller (1963, 1967, 1972, 1973a) and his
colleagues on the Juneau Icefield Research Program (JIRP)
in the Northern Boundary Range. More recent JIRP reports
(Miller, 1975a) and Jones (1975) emphasizes the assessment
of NeOglacial to present climatic changes in the Atlin region
and relate this to Little Ice Age glacier fluctuations and
small cirque glaciers on the Cathedral massif at the head of
Atlin Lake. Through these investigations interpretations
have been made of recent glaciological data on theIJuneau
Icefield and its peripheral glaciers, as well as of secular
climatic records since the 1840's at a number of coastal
stations in Alaska. In particular, temperature trends and
periodicities have been correlated with solar energy varia-
tions. It has been found that over the past several centuries
during periods of high solar activity (maximum sunSpot num-
bers, associated with increased corpuscular radiation) the
anticyclonic continental high pressure cell in the Arctic
interior has intensified and expanded outward beyond the

Coast Range to meet the North Pacific maritime low near the

16
Alaska coast. During such times, the Atlin sector has
experienced generally warmer and drier conditions. With
decreased solar energy (reduced charged particle radiation)
the clockwise-turning continental air mass contracts and
moves inland, with‘the result that the maritime low moves
landward, resulting in increased maritimity. The conse-
quence is more precipitation and related cooling in the
Atlin region. The zone of interaction between the high and
low pressure cells is the locus of storm paths which are
considered to shift in generally northeasterly or south-
westerly directions as the major pressure systems are
modified by variations in solar energy flux in the atmos-
phere.

Through the 1950's and early 1960's, these storm paths
followed a westerly or coastal orientation, with the interior
climate being warmer and drier than it was at the turn of the
century when interior ice masses such as the Llewellyn Glacier
(Fig. 2), were at their Holocene maximum. Affected by an
apparently 90-year cycle which reached its peak in the late
1940's and 1950's, the storm paths at present are shifting.
into a more interior position. The ultimate interior
migration at the end of the current downward swinging hemi-
cycle should tend to shift back shortly after the beginning
of the let century. This concept has been involved in the
interpretation of palynological records from Atlin bOgS,
covering even larger cycles (Miller and Anderson, 1974). In

fact approximately 800 and 2400 year cycles of climatic change

17
have been recently described (Miller, 1973a) which appear
to be corroborated by the research of other groups (e.g.
Bray, 1971). During the maximum Pleistocene stages, storm
paths are considered to have been in a far more easterly
position via the mean inland shift of cyclonic centers*
(i.e., affecting the region to even stronger maritime in—
fluences). It is this which presumably gives rise to the
centers of major glaciation well to the east and northeast
of the present tOpOgraphic axis of the Juneau Icefield
(Miller, 1964; 1973a).

The position of late—Wisconsinan and Holocene storm
paths in the Atlin district and their associated climatic
conditions have been discussed in more detail by Anderson
(1970) and Miller and Anderson (1974) based on palynolo-
gical as well as geomorphological evidence relating to the

past 10,000 years (Holocene) in the Atlin Valley (Table II).

* It is noted (see glossary) that the term "storm
paths" refers to the interpretation of peripheral winds in
the interaction zone between the low—pressure maritime cells
and high-pressure continental cells along the north Pacific
Coast. This is distinct from the "cyclonic tracts" which
refers to the movement of the center of low-pressure systems

which is normal to the system movement on paths as here
defined.

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CHAPTER IV
SOILS AND VEGETATION RELATIONSHIPS
Introduction

In this region a relatively wide range of parent
material, primarily glacial drift, results in major soil
variations. Of particular importance in the Atlin region
is the presence of carbonates in glacial debris in re-
stricted areas near limestone outcrOps. Thick deposits
of outwash material (valley train deposits) are in general
well-drained. The less common till deposits, although often
quite thiCk, are much less permeable.

In addition to the effect that drift character and
t0p0graphy have on drainage, the moderate to high relief
of the area results in considerable microclimatic variation.
It must be remembered that 1,000 feet (305 meterS) of eleva-
tion is approximately comparable to a climatic shift of 300
miles (500 kilometers) north, or the equivalent of a climatic
change to be expected in about 4° of latitude change.

Micro-climatic conditions are also extremely important
in the mountainous areas of this terrain because temperature,
precipitation, wind direction and isolation all vary con-
siderably throughout the area. In locations above 3500

feet (1067 meters) it is common to find sporadic permafrost

19

20

of considerable thickness.
Variations in Soil Maturity

Because of the generally cold winter climate and low
rainfall, even in summer, the soil forming processes are
lew. This, along with a generally rapid recession of ice
from the major Atlin and Gladys Lake valleys makes the soil
formation process of less value than expected in the rela-
tive dating of glacial sequences of the last major Wiscon-
sinan advances preceding the Holocene. In fact, in the
study area the Atlin III and IV and Gladys III and IV
moraines are essentially azonal. This means that there is
not significant variation in soil maturity on any of the
moraines of upper Wisconsinan age. The moraines of Atlin
I and Gladys I age, however, are quite well oxidized and
can be well distinguished from later moraines by their
yellowish (limonitic) oxidation. This oxidized material is
particularly well diSplayed on the upper southfacing flank
of Caribou Ridge and at the t0p of the valley flanks of
Porter Lake Valley where slumped morainic debris is revealed
by yellowish soil taluses. Similarly, the lower (basal)
tills in Boulder Creek, Pine Creek, Spruce Creek and McKee
Creek show similar oxidation, quite comparable to oxidation
of upper Woodfordian moraines in the mid-continent sequence

of the interior Great Lakes region in Michigan and Indiana.

Regional Comparisons

With further respect to the comparison with mid-

21

continent soils, Lietzke and Whiteside (1972), in a paper
dealing with the development of spososoils on 5000-6000
foot (1524-1981 meter) ridge t0psoils of nunataks on the
Juneau Icefield to the south, have found interesting
similarities with the glacial soils of Michigan. They
report that these soils are characterized by the develOp-
ment of thin micrOprofiles in that more maritime region
which has been substantially wetter than the Atlin area.
Specifically they found two distinct breaks in soil
develOpment with respect to elevation and hence, presumably,
to age. In the central icefield area about 40 to 50 miles
(64 to 80 kilometers) inland from.Juneau and 80 to 90 miles
(128 to 144 kilometers) south of Atlin, the soils were
generally young at elevations below 4000 feet (1200 meters)
showing little profile develOpment except here and there
some leaching and podzolization (Spodosols). Above this
level, to 5000 feet (1500 meters), buried paleosols were
found, with the upper section about the same age as the well
weathered and "easily crumbled rock fragments and a strong
soil develOpment." Above this to 7000 feet (2100 meters),
the surface soils appeared even older. They conclude that in
this section "the soil processes may have commenced about the
same time as the paleosol formation or earlier."

Above 7000 feet (2100 meters), they found no develOped
soils, only azonal rubble produced by intense frost action
and high winds which have removed the fines. They further

note strong similarities between the high-level Coast

22
Mountain soils in this area of existing glaciation and the
typical SpOdOSOlS of northern Michigan produced by late-
Glacial and Holocene soil forming processes.

In this context, in the Atlin area the uppermost till
of ground moraine found in the lower Fourth of July Creek
Valley near the Atlin road (and also in many road cuts to
the south and north toward Atlin and Jake's Corner), dis-
play spodosol develOpment, possibly related in age to the
younger nunatak soils described above. Lietzke (1969) has
found further similarities between the surface soils of
Atlin and the upper Great Lakes region, suggesting similar

ages...i.e., Valders and Port Huron.
Soils-Vegetation Characteristics With Age

In a study of soils in the Atlin Region, Lietzke
(1969) also found soil-vegetation relationships to be quite
useful in providing a qualitative assessment of comparative
soil parent material. It was found, for example, that well-
drained soils with a Cca horizon, or a high pH throughout,
support aspen (POpulus tremuloids) and Lodgepole pine (Pippg
contorta) in prOportions which vary with soil texture. Lodge-
pole pine with a lichen mat (primarily Cladonia) and
occasionally aspen, are characteristic of coarse-textures
and hence well-drained outwash soils. In a few locations
well-drained soils with a thiCk organic mat support Douglas
fir and Alpine fir.

Also it is found that soil develOpment in this region

23
is greater in fir stands than in the areas of Lodgepole
Pine. This was borne out by an earlier reconnaissance
study of soils between Whitehorse, Y. T., Carcross, Y. T.
and Atlin, B. C., in which Lietzke (1969) also reported the
soils to be primarily Aridisols or Inceptisols with Spodosol
(podZol—like) soils occurring at some sites. Palynological
and radio-carbon evidence (Anderson, 1970) and from the
writer's work in Boulder Creek (discussed later) suggest
that all of the surface soils in the lower Atlin valley,
north of the Village of Atlin, are not much older than
10,000 years B. P. These spodosols are suggested to be no
older than that.

As for differentiating the relative ages of Holocene
surfaces in the study area, it is reiterated that the
aridity of the region has slowed weathering rates. This
is to such an extent that only the soil-vegetation approach
has validity, although it is rec0gnized that efforts have
been made in other interior maritime regions to use soils
alone for distinguishing Neoglacial variations (Mahaney,
1975).

Poorly drained bog areas are also common in the area
and in those bOgS soils have develOped in less than 10,000
years B. P. (Anderson, 1970). In fact, as will be shown, a
drained bOg in the Boulder Creek Valley has surface soils
dated by the C-14 method to be less than 1100 years B. P.
(see Ch. VIII, p.155 ). When such bogs are at relatively

high elevation, they are often sites of Sporadic permafrost

24
and unusual frost feature develOpment. This is especially
true at elevations above 3000 feet (1067 meters). Dwarf
birch and willow along with various grasses and reeds are
the predominant bog vegetation. Locales of Sporadic perma-
frost are generally raised and contain fewer grasses and
reeds and commonly display a thick mat of lichens. Each of
these aSpects will be discussed in more detail in later

chapters.

CHAPTER V
PRE-WI SCONSINAN GLACIATI ON
Introduction

There is abundant evidence that at least one pre-
Wisconsinan glaciation far more extensive than any relating
to the Wisconsinan advances covered northwestern British
Columbia and adjoining sections of Alaska and the Yukon
Territory. Rounded and smoothed highland tOpOgraphy, often
seen as remnants on partially serrated ridge crests of the
higher peaks in the Coast Range (Miller, 1964a) show the
effect of ancient transfluent ice overriding well above the
obvious limits of Wisconsinan glaciation. This has also
been reported at the highest elevations in the Cassiar
Mountains 100 miles (160 kilometers) east of the Atlin
region bJohnston, 1926). Considerable evidence is given by
Denny (1952) and others that old weathered tills exist north
of the Atlin District, well outside of the Wisconsinan

glacial limits.
Morphoqenetic Phases

The pre-Wisconsinan phases of glaciation is in general
referred to as relating to an Intermontane 'Icecap' Glacia-
tion by Miller (1956) and the Continental Ice Sheet Stage by

Kerr (1936). This maximum pre-Wisconsinan phase has also

25.

26
been reCOgnized in southwestern British Columbia by Davis
and Mathews (1944) who designated it as a Phase IV glacia-
tion. Attained only in pre-Wisconsinan time, this larger
sub-continental glaciation presumably covered all summits
in the Atlin map area with some radial flow from locally
high accumulation areas. The main outflow of this glacia-
tion moved toward the south and southeast and also toward
the northwest, with major spillovers into the Skeena River
and down the Skagway, Whiting and Taku Rivers in the Boundary
Ranges (Fig. 1). There was also significant diffluent dis—
persal of ice out of the Chilkat, Chilkoot, and White passes
into the Lynn Canal area near Haines and Skagway, Alaska
(Fig. 1). The pre-Wisconsinan Cordilleran ice sheet in
some sectors extended inland over 500 miles (800 kilometers)
and so in lateral extent was probably larger than the Green-
land ice sheet is today. At its maximum extent it Spread
eastward in contact with the Laurentide Ice Sheet (Flint,
1957, 1971). A great deal more work should be done on this
little known aspect of the glacial geolOgy in this larger
region where the Laurentide and Cordilleran ice-sheets were

confluent.

Ancient Ice Centers and Outer Limits of Glaciation

The central areas of pre-Wisconsinan accumulation
serving as the key ice centers which affected the Atlin
area are believed to be not only on the inland flank of

the Boundary Ranges but also in the Cassiar Mountains,

27
about 150 miles (240 kilometers) east of the Coast Mountains.
The broad region between is considered to have been a linear
zone of climatolOgically controlled accumulation of large
magnitude, as Opposed to the more local and orOgraphically
controlled Wisconsinan ice centers located within the con-
fines of the Boundary Range. Thus in the pre—Wisconsinan
glaciation ice radiated in all directions, including a strong
flow to the east into Alberta. To the west, its directional
movement was largely controlled by the Coast Mountains and
the trans-range valleys, particularly in the early waxing
and late waning phases. Some of the pre-Wisconsinan ice
coalesced with local ice from the high coast mountains, but
the main drainage was largely deflected by these mountains
in its later flow (Miller, 1964b).

In the Atlin region the local and subsidiary ice
center, which added its increment to the major pre-
Wisconsinan glaciation and probably to a lesser extent in
early Wisconsinan time, was presumably well southeast of the
study area of this dissertation. In fact its axis and
center in the lesser phases of these ancient glaciations
appears to have lain about 10 miles (16 kilometers) or more
east of the international border where it crosses the Juneau
Icefield and probably adjacent to but somewhat east of Mt.
Nelles (8400 feet, 2560 meters) and Devils Paw (8584 feet,
2615 meters). This is noted here because these are the two
highest massifs which center the backbone of the Coast

Mountains adjoining the map area of this study.

28

Very little more can be said about details of pre-
Wisconsinan glaciation except by reference to Pleistocene
stratigraphy investigated elsewhere in Alaska and the Yukon
(Karlstrom, 1961; Denton and Stuiver, 1965; Flint, 1971;
Péwé, 1975). As these areas are well outside of the locale
of the present investigation they are not discussed here,
except to note that there is evidence for late Miocene or
early Pliocene glaciation in western and southern Alaska
(Miller, D..J., 1957; also Flint, 199 gig). Also, it should
be mentioned that drift of Illinoian age is present in nearly
every glaciated area of Alaska (Péwé, 1975), and indeed even
possibly in the west flank of the Fairweather Range in S. E.
Alaska (Goldthwait, personal communication). In northern
Alaska such ages are based on stratigraphic and radiometric
evidence near Nome, while in other regions assignment is
made On basis of surficial expression and position in the
local drift sequence. Weathering profiles on Illinoian
drift are thicker and more conspicuous than on Wisconsinan
drift (Péwé, et a1, 1965). In the Yukon-Tanana Upland of
Interior Alaska, the Illinoian snowline was found to lie
approximately 1300 to 1500 feet (400—500 meters) below the
present snowline. In this interior Alaska region the maxi-
mum.Wisconsinan snowline was about 560 feet (170 meters)
below the present snowline. On the southern Alaska Coast,
on the basis of cirque distribution and elevations, (Miller,
1961) the early Wisconsinan snowlines were as much as 2500

feet (762 meters) below present mean late summer snowlines

29
or névé - lines. In the dry interior of east - central
Alaska, the Illinoian snowline was found to be about 4100
feet (1250 meters) (Péwé and Burbank, 1960). By extra-
polating this type of evidence to the Atlin region, it is
again mentioned that the maximum Illinoian and earlier (?)
glaciations probably overrode all of the major ranges,
including higher peaks in the area.

Tectonic activity in the Coast Range and the interior
mountain ranges of British Columbia as well as the air mass
circulation have played a significant role in the extent of
pre—Wisconsinan glaciation. Along the coast of the Gulf of
Alaska, between Cape St. Ellas and Yakutat Bay (Fig. 1), the
recent glaciation is nearly as extensive as Wisconsinan
maximum glaciation. D. J. Miller (1957) attribures this to
higher mountains in recent time than in the Wisconsinan
time. The former lower elevations would have permitted
more moist air to find its way into the interior ranges,
including the ice center locale in the Cassiar Mountains
where overriding is evidenced aS high as 10,000 feet
(Johnston, 1926). Thus it is presumed that pre-Wisconsinan
Coast Range elevations may have been sufficiently lower to
have allowed greater air mass transfer across them.

The foregoing may also explain the apparent difference
in vertical Spacing of cirques in the tandem sequences found
on the coast compared with those found inland (Ch. VI). A
further amplification of these concepts is given by the

recoqnition of Illinoian cirques, with this age designation

30
primarily suggested by weathering profiles in the Yukon-
Tanana Upland. Here cirques have been found to be far
more extensive on the north facing lepes. This is inter-
preted by some as evidence for air masses moving in from
the north rather than the south-southwest as today (Péwé
and Burbank, 1960). This could also be manifestation of a
very marginal climate which favored snow accumulation on
the north-facing lepes. Another concept discussed recently
by Miller (1973a) and also Budyko (1972) is that with colder
conditions during pre-Wisconsinan time (and, indeed, even
with reSpect to the Wisconsinan maxima) the Arctic sea ice
as well as terrestrial snow cover was much greater in
lateral extent. CorreSpondingly in mountain regions there
were lower climatolOgical snowlines. This would mean in-
creased areas where high albedo pertained, thus substantially
lowering the incidence of absorbed radiation. This cooling
would upset the heat balance of the northwestern sectors of
_the North American and Eurasian continents, and in turn
bring a more maritime influence farther inland.

It Should be mentioned in conclusion, that north of the
study area there is some possibility of pre-Wisconsinan soils
on isolated sectors of the high ridges at the north end of
Lake Atlin near.Jakes Corner (Fig. 2). This is suggested
by the develOpment of Gossan soils well above the moraine-
delineated Wisconsinan ice levels (characteristics of these
soils are yet to be investigated but they are strongly

stained with limonite - possibly, however, this is related

31
to the presence of ferruginuous limestones). For the most
part, Wisconsinan glaciation was so intense in the region
that it has erased or at least severely altered the pre-
Wisconsinan geomorpholOgy. Similarly, periglacial pro-
cesses of the Holocene have somewhat altered some of the

evidences of earlier glacial processes in the study area.

CHAPTER VI
WISCONSINAN GLACIATION
Introduction

Extensive evidence for Wisconsinan glaciation in the
study area indicates that ice covered all but a few of the
highest peaks in northwestern British Columbia and adjoining
areas of Alaska (Miller, 1956, 1964b; Swanston, 1967). The
most recent phases of glaciation are middle to late Wiscon-
sinan in age. In reconnaissance studies, based largely on
air photo interpretation of the Southwestern Yukon region,
Hughes et a1 (1969) have mapped all deposits directly north
and contiguous with the Atlin Map area as late Wisconsinan
in age. They refer to this as the McConnell glaciation.
Assuming continuity in such contiguous sectors, and further
assuming that their assessment is valid, this would permit
interpretation of the main glaciation which formed the pre-
sent landscape of the study area to be late-Wisconsinan in
age. Except for the presence of deep soils on well-weathered
till and associated erratics scattered at a few high eleva-
tions, it is tempting to extrapolate this broad assumption
to the Atlin region. As will be Shown, however, there are
inconsistencies in the direct transfer of this overview

chronology in the Yukon to the study area in the Atlin region.

32

33

An early Wisconsinan or pre—Wisconsinan glaciation
has been documented for much of the Cordilleran outside the
immediate study area. Rutter, (1975) identifies an exten—
sive early Wisconsinan or pre—Wisconsinan advance in the
north-central eastern sector of British Columbia. Denny
(1952) recognized an "Old" till in the Summit Lake region of
northeastern British Columbia. From a study of cirque and
berm levels, Miller (1956, 1961, 1975) interprets an early
Wisconsinan glaciation (early Juneau/Atlin stage) in the
Juneau Icefield region (Taku District) which far exceeded
any of the later Wisconsinan advances. Evidence for this is
the existence of low-level cirque basins near sea level and
-high-level berms, as well as high-elevation erratics on
peripheral ridges and relatively small amounts of weathering
on high-level tills relative to what Should be expected were
these upland surfaces mantled with pre-Wisconsinan ground

moraine.
Evidences of Early Wisconsinan Glaciation

Within the study area too there is evidence for a higher
glaciation than the more recent sequence discussed below.
Though more highly weathered than the mid to late Wisconsinan
deposits, these high-level tillS also do not show the degree
of weathering one would expect had they been exposed for the
entire Sangamonian interglacial.

The only direct evidence for an early Wisconsinan

Maximum is the well-weathered soil mantle on peaks of White

34

Mountain (Fig. 1) 50 miles (80 kilometers) north of the
study area. There, at elevations above 4500 feet (1370
meters), gossan soil develOpment is apparent which may be
related to a pre—Wisconsinan surface. The carbonate com-
position of White Mountain must also be considered, as
weathering would be more rapid in this kind of litholOgy
than in the granitic peaks most commonly found within the
study area. Further field work in this region is needed
for a firm interpretation of an early Wisconsinan maximum
throughout the larger northern B. C. - southern Yukon region.

In the Atlin district ice coming out of the Cassiar
Mountains covered areas to the north, east and south.
During the Wisconsinan glaciations this ice was largely
diffluent, i.e. channeled by tOpOgraphic lowS as it flowed
in a distributary fashion onto the Stikine, Teslin, Taku and
Yukon Plateaus. The part of the Cassiar Lobe which most
affected the study area passed northward across the Teslin
Plateau and into the Gladys Lake depreSSion (Fig. 2 & 3).
In the present report this portion of the Cassiar Lobe is
referred to as representing Gladys Ice. During Wisconsinan
time the other significant glacial mass affecting the study
area was the Coast Range ice-sheet which extended into the
Atlin Lake depression from the south. This glacial mass is

referred to as representing Atlin Ice.

Evidences for a Middle to Late—Wisconsinan Seguence

The following chronolOgy illustrates mid to late

35
Wisconsinan glacial activity within the area of investigation.
The first stage of both Atlin and Gladys ice is that stage
for which there is no field evidence which can clearly
delineate the extent of ice within the area of immediate
concern. Again, though it Should be kept in mind that there
was at least one earlier and more extensive Wisconsinan
glaciation preceding the glaciations discussed below.

The Fourth ofIJuly Creek Valley is divided into three
major glacial-morphological sections (Fig. 4). The ngg;
Fourth of.July Creek Valley is from the juncture of the
Fourth of July Creek with the Porter Lake Valley, (the
location of the Ruffner* Juncture moraine) to its mouth at
Atlin Lake. This region contains thick till deposits and
is the section which contains McDonald Lake.

The Intermediate Fourth of Jplv Creek Vallgy extends
from the juncture moraine to the confluence of Volcanic Creek
and is characterized by glacio-fluvial terraces and by
glacio-fluvial erosion in sectors where much drift has been
'removed.

The area above Volcanic Creek to the drainage divide,
is referred to as the Upper Fourth of July Creek Vallev and
includes the relatively flat plateau at the drainage divide
of Fourth of July Creek and Consolation Creek. This region
contains a complex esker system, many kettles and some
Sporadic occurrences of frost mounds.

In Figure 5 the maximum extent of each of the following

* Named for Ruffner mine near moraine complex.

36

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MT . LEONARD

Figure 5b. Maximum extent of
Atlin ice

39

stages is illustrated.

Gladys Ice Moraines

The glacial map of the southern Yukon Territory
(Hughes et a1, 1969) shows that ice from the Teslin Valley
extended north through the Teslin depression where it was
joined by ice flowing directly west from high centers in the
Cassiar Mountains. The Gladys portion flowed through the
northern part of the Atlin map region in the direction North

450 to 700 west (Fig. 3).

Gladys I
Ice of the Gladys I stage completely overrode the

summits of Mt. Ewing (5410 feet, 1649 meters) and Mt. Boofus
(5060 feet, 1542 meters - as Shown in Figure 3). Good
evidence for this is shown in Figure 6 where the peak is
smoothed with fluted and drumlinoid t0p0graphy on its flanks.
Mass wasting, local cirque develOpment in the waning phases,
and subsequent effects of relatively intense glacio—fluvial
activity have made it impossible to delineate the true extent
of this phase at such type localities as Black Mountain to
the northwest (Fig. 7). However, the existence of relatively
unweathered high-level till and erratics (Fig. 8) on the tOp
of the mountain between Chehalis and Davenport Creeks (i.e.
on Gladys Mountain for the purposes of this study) indicate
that Wisconsinan ice of Gladys I Phase did indeed override
this peak at an elevation greater than 5500 feet (1676

meters). This state is assigned a middlg Wisconsinan age on

40

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42

Figure 7. Fluted drift north of Black Mountain

The drumlinoid fluted drift of Gladys Ice cut
by run-off channels of Atlin Ice and Gladys
Ice. Note truncation by kame terraces of
Atlin Ice. Black Mountain is in the lower
part of the photograph.

Canadian Energy, Mines and Resources
A11371-335, 1948

Figure 8. Erratics on Gladys Mountain

Looking southeast across the Gladys depres-
sion from summit of Gladys Mountain.

43

 

IV.
a. o

A
\ 4t

0.»

44
the basis of sequential levels and the great thiCkness of ice
in the Gladys depression which would be required to achieve
this level. Geomorphic reasons for this interpretation are
considered below.

A more highly weathered pre-Atlin I till above 4500
feet (1372 meters) on Mt. Leonard and Mt. Vaughn is evidence
that on Gladys Mountain this high ice phase did not make its
way down the Fourth of July Creek Valley at this elevation
on the valley walls. Serving as obstructions to this Gladys
I ice were the Mt. Leonard, Vaughn, Barham and Edmund massifs
(Fig. 3). These channeled diffluent lobes into the Surprise
Lake Valley and down the present Fourth.of July Creek Valley.
Any evidences of terminal positions of these lobate ice
tongues in these valleys has been erased by the overriding
effects of later ice.

Also any sharply defined morainic evidence of these
limits has apparently been destroyed by the later develOp-
ment of local high-elevation valley glaciers out of cirques
at the head of the upper Consolation Creek Valley and on the
flanks of Mt. Barham and Mt. Edmund. Thick till deposits,
however, are present to an elevation of greater than 5000
feet (1524 meters) in Spite of the fact that in this sector
no clear upper limit of Gladys I ice has been delineated.

On the flank of Mt. Barham, due west of Consolation
Creek, at an elevation of slightly more than 4500 feet (1371
meters) evidence for overriding during this and earlier

glaciations is shown by deep incisions into a bedrock ridge

45

resulting from sub—glacial streams (Fig. 9). Though
modified by Gladys I stage ice, these features were carved
by-precedinq high-level glaciation as well.

The maximum western extent of the Gladys I stage in
the northern part of the Atlin map area is also not dis-
cernible as more recent kame terraces from later Atlin ice
have abruptly truncated the fluted and drumlinoid t0po-
graphy of Gladys I ice and, indeed even those of the Gladys
II stage (Fig. 7). Although this truncation has destroyed
evidence of the maximum position of Gladys Ice it has
actually demonstrated the areal insignificance of Atlin
Ice in this region, and its relatively restricted nature
after the last major stage of Gladys ICe.

From the degree of surface weathering and the forma-
tion of tors on the upper ridges of Mount Ewing (Fig. 6)
the Gladys I stage is designated as middle Wisconsinan.
This general phase of glaciation correSpondS to the Greater

Mountain Ice-sheet Phase of Miller (1964b).

Gladys II Stage

Prominent lateral moraines on Gladys Mountain repre-
sent a lower phase of Gladys Ice. Here the multiple moraine
system at 5200i feet (1585 meters) flanks the eastern side
of the mountain and can be traced laterally for several
thousand feet. During this phase ice again made its way up
the lower Consolation Creek Valley, and as well into the
Surprise Lake sector and the upper and intermediate Fourth

of July Creek valleys to the west. At this time the plateau

46

Figure 9. Bedrock incision on the eastern
flank of Mount Barham

Figure 10. Moraines on Marble Dome

Canadian EnergY. Mines and Resources
A11390-275

47

 

 

’F‘.

It!

(I:

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48

area northwest of upper Fourth of July Creek was filled with
ice which in areal extent equalled the coverage experienced
in Gladys I time. This possibly suggests a more temperate
thermo-physical character than the Gladys I stage...i.e.
more mobile ice. No lateral moraines, however, are found

on Mt. Ewing. These were presumably destroyed by subsequent
mass wasting. The lower phase Gladys II ice was also pre-
sumably fed by prominent cirque sources in the highlands
near and above 5000 feet (1524 meters). Documenting this
phase is a pronounced moraine complex on Marble Dome (Fig.
10), at less than 5500 feet (1676 meters).

Further moraines can be traced to quite a low eleva-
tion (i.e. 3000 feet, or 900 meters) in the Fourth of July
Creek Valley to a moraine complex above Spear Point Terrace
(Fig. 11). From the geometry of these moraines the ice
gradient from Gladys Mountain to Spear Point Terrace is
interpreted aS about 100 feet (30.5 meters) per mile. The
moraine is waterworked and densely kettled (Fig. 12).
Proximal to the valley wall large fragments of talus cover
the moraine. (This moraine complex was also affected by
subsequent local ice from high level cirques as discussed
later).

There are no lateral moraines in the high—level plateau
area northwest of Gladys Lake. This again makes it impossible
to do more than speculate the farthest extent of ice into
this area during Gladys II time. Ice gradients elsewhere

indicate that a glacier sheet covered most of the low-level

plateau.

49

Figure 11. Stereo pair of the intermediate
Fourth of July Creek Valley

Volcanic Creek juncture with the Fourth of
July Creek is Shown in lower part of photo-
graph. Note esker development and fluvial
and glacio-fluvial terraces on valley wall.
The Spear Point moraine is just below the
esker develOpment.

Canadian Energy, Mines and Resources
A11381-328 & 329, 1948

 

51

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52

 

53
Still a Mountain Ice Sheet, the Gladys II phase re-
presented a major complex of still—stands in the general
downwasting of Wisconsinan ice. As such it is considered

to relate to the beginning of late-Wisconsinan time.

Gladvs III Staqe

Although not always apparent on strictly morphostrati-
graphic grounds, lateral moraine positions are clearly
identified for this stage by the distinctiveness of drainage
controls on the flanks of Gladys Mountain. Here deeply
incised stream channels of strikingly limited extent are
found to terminate where the water flowed onto the ice in
the Gladys Lake depression (Fig. 13). Morphologically-
controlled by the ice position these channels delineate
three major levels of downwasting - i.e. at about the 4900
foot (1493 meters) level (designated as Gladys III-A), at
4500 feet (1371 meters) (noted as Gladys III—B), and at 4000
feet (1219 meters) (noted as Gladys III-C). These separate
levels are all considered to be parts of one major stage or
general ice-level. This is because of the Similarity of
erosion features associated with each and because of the
high gradient and common angle of lepe of the ice associated
with the three events. In this connection it is worth
mentioning that even still-stands can be oscillating to the
extent that there are intervals of greater and lesser abla-
tion in consequence of short-term climatic variations.

No moraines for this series are present in the upper

Fourth of July Creek or in lower Consolation Creek valleys.

54

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co mmcflmuoz 00H mhpmao .mH mummmw

 

56

The steep gradient which has been extrapolated suggests that
this ice could not have made its way very far up the valley
of Consolation Creek and that it was probably confined to
the broad flattish region lying northwest of Gladys Lake.
In that sector there are three Significant underfit streams
still draining the northern boundary of the Atlin Map two of
which are Shown in Figure 14. The broader valleys in which
these streams now lie as underfit drainage ways served as
major runoff channels in the Gladys III stage. A fourth
underfit stream (Fig. 15) lying beyond an intervening 3500
to 4000 foot (1066 to 1219 meter) ridge also served as an
important runoff channel for the contiguous Atlin ice and
as well for the downwasting of Gladys ice in the Gladys II
stage.

The upper Fourth of July Creek and middle Consolation
Creek valleys were apparently filled with ice at this time.
But local accumulation to the volume of remnant ice left from
earlier Gladys phases. The significance of this local pro-
venance of ice is discussed in a latter section.

The extent of the Gladys III glaciation is considered
to correspond to Miller's (1964b) Lesser Mountain Ice-sheet
Glaciation in the Boundary Range and is presumed as lapg

Wisconsinan in age.

Gladys IV Stage

Also late Wisconsinan in age is a lesser glaciation

comparable to the Extended Icefield Glaciation of Miller

57

Figure 14. Underfit glacial runoff streams
Note the drumlinoid t0pography and small
eskers in eastern portion of photo.

Canadian Energy, Mines and Resources
A11371-329

Figure 15. Atlin Ice and Ice Runoff Channel
Photo northeast of Black Mountain on Yukon
border.

Canadian EnergY. Mines and Resources
A11371-334

 

59
(1964b). In this glaciation the ice downwasted and stagnated
in the Gladys Lake Valley. This is evidenced by a deeply-
pitted surface at an elevation of 3000 feet (914 meters)
(Fig. 12). Associated deposits of massive prOportionS are
reflections of an intensive fluvial deposition on tOp of
slowly downwasting dead ice. A complex pitted region at
.Airplane Lake just northwest of Gladys Lake marks the

northernmost extent of this important dead-ice zone (Fig. 3).
Atlin Ice Moraines

In middle to late Wisconsinan time, a lobe of Boundary
Range ice also filled the Atlin valley. During its down-
wasting phases it served to create much of the landform
configuration in the vicinity of the village of Atlin and
on the adjoining flanks of the main Atlin valley. This
glacial mass sent distributory fingers up into all valleys
which today are tributaries of the main Atlin valley. The
limit of Atlin ice in these peripheral valleys was, of
course, controlled by gradients and thicknesses of the main
lobe, and as well by thermal prOperties of the ice, by bed-
rock lepe of the valley and in some cases by the addition
of ice flowing out from local accumulation areas. Except
for weathered high-elevation tills on Mt. Leonard, nearly
all evidence of a maximum extent of this earlier advance of
easterly-flowing Atlin ice into the Fourth of.July Creek
Valley was completely removed or covered by deposits from
the later Gladys ice or by ice relating to Extended Icefield

phases originating in the local accumulation centers.

60

Atlin I Stage

Following what is presumed to have been a pre-classical
Wisconsinan glaciation, evidence for the next most extensive
advance of Atlin ice is given by a steeply dipping and well-
weathered yellow-colored moraine on the main southeastern
ridge of Mt. Ewing, referred to in this text as Caribou
Ridge (Fig. 3). This moraine extends downward from an
elevation of 4300 feet (1310 meters) to 3800 feet (1219
meters) at which elevation it is truncated by a quite un-
weathered younger moraine of late Atlin ice. An Atlin I
moraine at 4000 feet (1525 meters) can be traced for 2 miles
on the eastern Side of the mountain south of Mt. McIntosh
(presently referred to as Mt. McDonald). A moraine lying
at 4500 feet (1372 meters) on the west side of Mt. McDonald
further delineates the extent of ice from the main Atlin
valley in this early phase.

The steep gradient of the Caribou Ridge moraine
indicates that though ice was thicker in the lower Fourth
of.July Creek Valley than it was in later Wisconsinan stages,
it was characterized by a steeper snout. This means it did
not make its way into the intermediate Fourth of July valley
as far as it did later on - i.e. when it was thinner and
hence at lower levels. Once again it is suggested that
this stage may have been thermOphysically more polar (again
the ice being not as mobile) in middle Wisconsinan time. AS
such it would correspond to the Greater Mountain Ice-Sheet

Phase in the Boundary Range MorphOgenetic Sequence (Miller,

1964b).

61

The Atlin I ice passing up into Porter Lake valley
(Fig. 3) did not have as steep a surface gradient as sug—
gested by the fact that its lateral moraine does not lepe
as steeply. The total distance which ice flowed up this
valley at that time is not known, because its northern
limit is obscured by outwash. Also there is no evidence
of a Porter Lake lateral moraine on Mt. McIntosh where
local cirque activity and mass wastage have been particular-
ly active. During this phase it is clear, however, that
main valley ice did flow through the valley between Mt.
McIntosh and Mt. McDonald and pass well up through the

valley of Tel-Cabin Creek (Figs. 2 and 3).

Atlin II Stage

A much younger ice advance is documented by a lower-
level and less weathered moraine on the valley walls in the
intermediate Fourth of July Creek valley. Near the juncture
of Crater Creek valley this moraine lies at about 3700 feet
(1128 meters) elevation and can be traced down to 3000 feet
(914 meters) at a point where it is truncated by terraces
cut by meltwater from the upper Fourth of July Valley. Its
terminal position lies beyond the entry of Volcanic Creek
(Fig. 3). The Atlin II stage is the Intermediate Mountain
Ice—sheet Phase (Miller, 1964b) and is considered to be of
early late-Wisconsinan age. From comparable weathering it
is suggested that the bold lateral Wisconsinan moraine at
5300 feet (1615 meters) elevation near the Camp 29 research

station in the Cathedral Glacier Valley (Figs. 2 and 3) re-

62
presents the same phase of glaciation.(Jones, 1974, 1975).
From that location the ice would have had a surface gradient
of about 50 feet (15 meterS) per mile.

Ice channeled through the Porter Lake valley reached
a point about 1% miles (2.4 kilometers) beyond Porter Lake
(Fig. 3) where a moraine complex truncates the fluted t0po-
graphy produced in Gladys II time. This is one of few
localities where there is evidence that there was contem-
poraneous and contiguous Atlin Ice and Gladys Ice. This
total glacier mass included Atlin ice which.made its way up
the Tel-Cabin Creek Valley and through the valley south of
Mt. McIntosh.

Other major intrusions of Atlin ice in this stage were
between Halcro Peak (Mt. Hitchcock) and Mt. Carter to a
position three miles (4.8 kilometers) north of Indian Creek
(northwestern corner of map in Fig. 3). Here it is again
indicated that Atlin II Stage ice met Gladys II Stage ice
in an area where molded drumlinoid tOpOgraphy produced by
Atlin Ice merges with the landforms produced by the most
westerly limit of Gladys Ice. During the downwasting of
these ice masses, meltwater drained north across the area
and resulted in fluvial dissection of this drumlinoid and
fluted terrain which had initially been produced by ice
masses of the Gladys I and II stages (Fig. 14).

Although the originating ice-sheet of the Atlin II
stage can be considered morphogenetically as relating to a

Greater Mountain Ice-sheet Phase in the source regions of

63
the Boundary Range, the confinement of this ice to major
valleys with some diffluence through minor higher-level
valleys in the Atlin region leaves room for misinterpreta-
tion of the magnitude of the glaciation unless the total

regional picture is kept in mind.

Atlin III Stage

In the waning stage of Wisconsinan ice in the inter-
mediate and lower Fourth of July Creek valleys three closely
associated moraines near and just below the juncture of the
Porter Lake valley are grouped into the Atlin III Stage.
These are referred to as the Ruffner moraine complex.
Morphogenetically they are bold terminal and lateral moraine
remnants. They are also well displayed on the flanks of Mt.
Leonard.

The first phase, Atlin IIIA, is represented by a sandy
moraine (Figs. 16 and 17) with its highest section in this
locale being at 3000 feet (914 meters). This moraine is
closely associated with what is designated as Atlin IIIB,
which is the juncture moraine complex (Fig. 18). In this
situation Atlin IIIA ice again moved up the Fourth ofIJuly
Creek valley to the mouth of Volcanic Creek, but this time
at a much lower elevation. A pitted ground moraine repre-
senting this stage is Shown in Figure 17.

The juncture moraine complex played a Significant
role in the late-Wisconsinan drainage of the Fourth of July
Creek. During this stage the ice front continued to block

off the lower Fourth of.July Creek valley from its inter-

64

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66

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68

mediate sector until the Porter Lake valley Opened to re-
verse its pre-glacial drainage and permit channeling of
meltwater into the Porter Lake depression. For some time
ice-cored moraines remained here, veneered with outwash
deposited by streams from melting ice in the upper Fourth
of July Creek valley. There was some deposition too from
feeder glaciers in re-activated high-level cirque systems,
as discussed later.

As the ice receded from the juncture zone the massive
kame-moraine complex continued to deflect meltwater through
Porter Lake valley resulting in deposition of thick mor-
ainic material southeast of McDonald Lake. This is designated
as the Atlin IIIC moraine, much of which was well-washed by
meltwater. Low—level lateral moraines on the southeastern
side of McDonald Mountain represent this final phase of
glaciation in the juncture area, which we are reminded
separates the lower Fourth of July Creek Valley segment from
the intermediate valley segment.

The juncture moraine is composed of water-worked till
and contains many gently leping yet still relatively smooth
areas dramatically‘punctuated by kettles and pits revealing
that during deposition there was a large volume of meltwater
coming out of the upper Fourth ofIJuly Creek valley. Some
of this outwash, coming from different provenances, buried
dead ice zones and so also became incorporated in the junc-
ture moraine. Though no identifiable lake deposits are

found today in the intermediate Fourth ofIJuly Creek valley

69

above the juncture moraine, an ice-dammed lake is presumed
to have existed there until the glacier receding below the
juncture Opened up flow into Porter Lake valley. A series
of fluvial terraces to the southwest and in the Porter Lake
valley itself reflect variations in fluvial activity
associated with the final stages of deglaciation in the
intermediate and upper valley and in the adjoining highlands.

The gradient from upper Fourth ofIJuly Creek to the
juncture moraine today is about 50 feet (15.25 meters) per"
mile. This gradient is sufficient to have flushed out
drift in the western part of the intermediate Fourth of
July Creek Valley. The present gradient from the juncture
moraine to Porter Lake is at most 15 feet (4.6 meters) per
mile, which of course resulted in extensive valley train
deposits being laid down in the Porter Lake Valley. The
higher-level fluvial terraces of the Porter Lake Valley are
also densely pitted indicating the presence of much dead ice
in the valley during the period when glacio-fluvial deposi-
tion began.

Correlation of these terraces is discussed in a later
, section. Overall, this phase of glaciation is considered to
correspond to a waning phase which altered with the down-.
wasting of an Intermediate Mountain Ice-sheet (Miller,

1964b)in the Boundary Range.

Atlin IV Stage
A large recessional moraine in lower Fourth of July

Creek Valley lies just below Lower McDonald Lake (Fig. 3).

70
This is evidence for the last significant pulsation of Atlin
ice in the Fourth of July Creek Valley. Deposition of this
moraine impounded water and formed Glacial Lake McDonald,
(Fig. 5) a body of water which iS contiguous to the large
kame moraine at the juncture of Porter Lake Valley and the
intermediate Fourth of July Creek Valley. This recessional
moraine and the glacial still-stand it represented was
followed by broad deposition of ground moraine during
rather continuous and rapid downwasting of Atlin valley
ice in late-Glacial time.

During this thinning and retreat drainage Of Glacial
Lake McDonald was initially through the juncture moraine
(Ruffner) which is diagrammatically illustrated on page
As ice downwasted and retreated Glacial Lake McDonald
drained out via the Fourth ofIJuly Creek to a level com-
parable to the high water elevation of present Atlin Lake
and hence was probably reduced to near its present day
extent. The reversal of this drainage, which was Iikely
both supra and sub-glacial, was assuredly representative of
what was happening to all distributaries in the Atlin glacia-
tion system during the intervals of final downwasting of
Atlin Valley ice.

Glacial Lake McDonald was also produced by wasting ice
blocks in morainic material near the valley juncture. Ground-
water produced from ice in these deposits had to drain to the
Southwest, while contemporaneous surficial drainage from the

upper Fourth of July Creek valley was diverted through Porter

71

Lake valley.

Atlin V Staqe

Evidence for the final phase of Atlin ice in the over-
all Fourth of July Creek valley is shown by pitted deposits
on the gentle floor of the main Atlin valley near the pre-
sent-day mouth of the Fourth of July Creek (Fig. 3). These
deposits lie at an elevation of about 2250 feet (785 meters).
That the final retreat of main Atlin valley ice in this
region involved systematic downwasting and frontal reces—
sion is shown by the presence of drumlinoid hills below the
confluence of Two.John Creek and Fourth of July Creek (Fig.
18). It is also revealed by the lack of dead-ice features
above 2250 feet (785 meters) and by the existence of fluted
tOpOgraphy in places down to an elevation of 2500 feet (762
meters). Other than very low-level kettles and associated
pits in glacio-fluvial mantles and a few eskers, the area

lacks many striking evidences of this ice stagnation.

Local Cirgue Activity

One of the most striking features of the Fourth of
July Creek valley is the greater thickness of glacial drift
on the southeast Side of the valley and on the flanks of Mt.
Leonard, Mt. Vaughn, Mt. Barham, and on the Mt. Edmund
massif. This is evidence that during all of the glacial
phases described glaciers originating within nearby high-
elevation cirque basins were active and indeed as cirque-

headed alpine glaciers played an ultimate role in the volume

72

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74

of highland ice involved, and consequently also in the
provenance of some of the glacial drift in the lower
valleys.

NO terminal moraines of high-level cirques were found
in the main Fourth of July Creek or Consolation Creek val-
leys, only ground moraine. Thus, it is concluded that the
maximum develOpment of cirque glaciers took place during
phases Of Mountain Ice-sheet GlaciatiOn when Atlin and/or
Gladys Ice filled these valleys. In the waning phases of
the glaciation which filled_these deep valleys with ice,
glaciers from the cirque sources thickened, re-advanced and
coalesced, eventually to downwaste again leaving classic
dead-ice features behind.

That retracted glaciers occupied the cirques even
after all ice had melted from the main lower Atlin and
Gladys Lake valleys is evidenced by glaciO-fluvial erosion
of the youngest lateral moraines in these main valleys, and
as well by alluvial fan deposition related to much greater
water flow than at any time during the Holocene or at pre-
sent. This is particularly well illustrated where Consola-
tion Creek reaches the lower gradient of the water divide
of the Fourth of July and Consolation Creeks. Here a large
alluvial fan greater than one-quarter square mile in area
(0.4 kilometers squared) is densely dissected by abandoned
channels. The compOsite of erosiOnal and depositional
evidences which has been discussed above confirms that there

was intensive cirque activity in the waning phases of

75
Wisconsinan glaciation. It is also probable that there was
re-occupation of some of the high level cirques in Ander-
son's cool—dry zone II (Table II) 750 to 2500 years B. P.,
equated to the early NeOglacial, and in the late-NeOglacial

Little Ice Age of 300 to 500 years B. P. (Matthes, 1949).

A Seguence of Tandem Cirques

A detailed inventory of cirque levels and the tandem
arrangements where they occur in the study area is rendered
difficult by the 500 foot (152 meter) contour interval of
the Atlin map.* A cirque distribution is obtained, however,
by integrating map and air-photo interpretations, but again
this is made imprecise by the large contour interval which
exists on available maps. Helping the interpretations,
however, is the availability of larger-scale (1:50,000)
maps in the Atlin Village to Teresa Island sector (Canada
Dept. of Mines and Technical Surveys, 1956) where, although
out of the study area, comparable cirques can be measured
at loo—foot contour intervals. In general, therefore, four
distinct cirque levels can be identified. Their average
elevations are at 6000 feet (1829 meters), 5500 feet (1676
meters), 4900 feet (1494 meters) and 4100 feet (1249 meters).

In Table III the cirque levels are correlated with
levels of coastal cirques in southeastern Alaska as deter-
mined in previous studies by Miller (1960) and Swanston

(1967). Numbers are assigned on the basis of Miller's

* Canadian Department of Mines and Technical Surveys,
1954; also shown as Figure 3 in pocket.

76

TABLE III

COMPARISON OF CIRQUE ELEVATIONS IN THE ATLIN REGION
AND SUGGESTED REGIONAL CORRELATION

Cirque
Level

C 1

Southeastern Alaska

Juneau Icefield

apd the Taku District '
Miller (1956)

300 ft (90 m)

1000 ft (305 m)

1800 ft (550 m)
King salmon-Port Huron?

2500 ft (760 m)
Tulsequah-Early Valders?

3200 ft (975 m)
Sittakany-Late Valders?

3900 ft (1190 m)
Early Holocene and
NeOglacial

4600 ft (1400 m)
Thermal Maximum

Prince of Wales

Islapd
Swanston (1967)

0.500 ft (0-150 m)

650-950 ft (200-
290 m)

1050-1350 ft (320-
410 m)

1450-1950 ft (440-
590 m)

2050-2650 ft.(625-
805 m)

77

TABLE III

COMPARISON OF CIRQUE ELEVATIONS IN THE ATLIN REGION
AND SUGGESTED REGIONAL CORRELATION

Atlin Map Area

 

Cirque Fourth of July Qaphgggaliflymmg
Level Creek Region Miller (1975)
Jones (1975)
C 1
c 2
c 3 4100 ft (1249 m)
c 4 4900 ft (1494 m) 4500 ft (1370 m)
c 5 5500 ft (1676 m) 5100 ft (1550 m)
C 6 6000 ft (1829 m) 5800 ft (1770 m)

c 7 6500 ft (1980 m)

78
(op. cit.) original designations. AS to be expected with
this inland positioning the series in the study area
averages 2275 feet (693 meters) above corresponding levels
on the Juneau Icefield, which, interestingly enough, is
close to the difference in elevation today between the mean
nevé-lines on the coastal versus the continental Sides of
the Juneau Icefield (Miller, 1975a).

Precipitation patterns in the study area were oro-
graphically controlled in the waning stages of Wisconsinan
glaciation. The cirque level at the mean elevation of
maximum snowfall during the mid-Holocene (Thermal Maximum)
is not represented as there are no peaks in the immediate
study area above the projected elevation of 6875 feet (2095
meters). In the study area one semi-permanent snow-field
was found at an elevation about 6200 feet (1890 meters)
which, allied with ice-filled cirques at the same elevation
on the Cathedral Massif (Jones, 1975), supports the concept
of occupation of cirques in this region at the projected
elevation of 6875 feet (2095 meters) during the mid-

Holocene.*

Effects of Late-Glacial Climatic Ameliorations and Rising

FreeZlng Levels

 

The plateau area, which is the drainage divide for the
Fourth ofIJuly Creek and Consolation Creek, lies at an eleva-

tion of 3200 to 3400 feet (975 to 1036 meters). During the

* Of interest is that this elevation (5th interior
level) does indeed support a set of higher cirque basins in
the vicinity of Camp 26, on the continental flank of the
Juneau Icefield - i.e. some 15 miles southwest of Llewellyn
Inlet at the head of Atlin Lake (Fig. 2).

79

Atlin I and Gladys I stages this sector served as a major
accumulation zone adding substantially to ice masses from
both source areas. In the waning phases ice from the main
nourishment zones of these two glaciations thinned, as
described in a previous section, but a large volume of
later-stage ice from Mt. Barham and Mt. Edmund continued
to feed the plateau.

During each climatic amelioration as freezing levels
(see glossary) rose and zones of snowfall migrated to higher
and higher cirque levels the plateau ice stagnated and down-
wasted forming esker complexes and ground moraine kettles,
that is on the plateau till plain. In a final phase
associated with this deglaciation a rather extensive ice-
dammed lake formed (Fig. 5). In several localities subse-
quent stream cutting during the Holocene exposed silt-clay
lacustrine deposits up to at least 15 feet (4.5 meters)
thick. NO true varved structures were found in these lake
sediments. At one exposure lacustrine deposits drape an
esker, truncated by post-Glacial erosion. Also these dis-
sected lake deposits lie conformably on and beside sandy,
kettled moraines with numerous eskers and crevasse-filling
features present.

It is of interest that the ice-dammed lake in this
plateau sector formed while ice was still present in most
of this high area. Final drainage of the lake followed

lowering and recession of Gladys Lake ice to the Gladys III

position.

80

As noted earlier, the thickest drift deposits are con-
fined as an assymetrical wedge on the south Side of the upper
and intermediate sections of the Fourth ofIIuly Creek Valley.
This drift was largely deposited by ice from the high-level
cirque-headed glaciers to the south. The extent of this
cirque ice is impossible to determine except to note that
it materially added to the relict Atlin and Gladys Ice
resting in the upper Fourth of July Creek Valley. In the
final deglaciation all of this ice downwasted together some
10,000 years B. P. (see later discussion with respect to
dating methods). It was during this time that the en-
glacial eskers develOped, the complex esker system of which
is shown in the Sketch map of Figure 19.

Two Opposing tributary patterns are notable in the
esker system and indicate two directions of flow. From this
we see that the subglacial drainage divide was some 3 miles
(4.8 kilometers) south of today's water divide. The explana-
tion lies in the englacial (ice contact) rather than sub-
glacial nature of some of these drainage ways.

The Fourth of July plateau was the most continuously
glaciated region in the whole study area. During the onset
of major Wisconsinan glaciation as the névé-line rose high-
level cirques increasingly fed the plateau. This was
followed by invasions of the main ice-sheet during Atlin and
Gladys Ice times. The final phase of glaciation also was
strongly influenced by nourishment from the local cirque

centers. In the plateau area, however, the extent of

81

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82
glaciation between early and middle—Wisconsinan time is not
known because the evidence was so completely erased by the

intensity of late-Wisconsinan glaciation.

Radiocarbon Dates Relevant to Deglaciation

In the upper Fourth of July Creek area alder twigs
at the base of a thick peat layer in a peat plateau at about
the 3500 foot (1970 meter) level have been collected and
dated by radio-carbon technique. A 9315 i 540 C-14 years
B. P. age (Geochron, 1972) on these samples gives a minimum
date for the plateau's deglaciation, i.e. that of the upper
Fourth of July Creek Valley. A C-14 date was also Obtained
from basal organic material in a bOg along Fourth of July
Creek 4 miles (6.4 kilometers) farther down valley and at
an elevation of 2900 feet (880 meters). This date is 8050 I
430 C-14 years B. P. (Geochron, 1973). In this case, how-
ever, the lowest contact between organic and inorganic
sediments was not reached, suggesting that the beginning of
peat formation at this locality was even greater than 8050
C-14 years B. P. In view of the first sample noted above
the onset of vegetation following deglaciation is suggested

to have taken place somewhere between 9000 and 10,000 years

B. P.

History of Glacio-Fluvial Terraces
in the Porter Lake Valley

The late-Glacial and Holocene glacio-fluvial history

of the Porter Lake Valley is well recorded in an unusually

83
complete set of paired and non-paired fluvial terraces near
the junction of the lower Fourth of July Creek Valley and
Porter Lake valley (Fig. 17). On foot traverses across
this zone a detailed tOpOgraphic and textural transect was
produced. The observations were based on aneroid altimeter
readings referenced to the Atlin meteorological station on
the shore of Atlin Lake. Details of this transect are given
on line AB which is the cross-sectional plot in Figure 20,
Shown also in plan on Figures 3 and 5b. Differential
elevations are plotted as absolutes, and by use of a Wallace
and Tiernan surveying altimeter the elevations of each ter-
race On this transect are indicated to within 25 feet (7.6
meters) of actual elevation above lake level, corrected to
a base datum at mean sea level.

During the Atlin III stage, the terminus of Atlin ice
was at the juncture of Fourth of July Creek Valley where, for
some time, it remained in contact with the narrow drift-
filled valley of Porter Lake and the intermediate Fourth Of
July Creek Valley segment. Evidence for this is given by
the massive dimensions of the terminal moraine and the
abundance of mixed drift the concentration of which has
been described for that sector. It is presumed that ice-
cored moraines filled even more of the valley following
downwastage of the preceding advance of Atlin II ice. Not
much ground moraine is visible today, however, several seg-
ments remain and are veneered with outwash from the inter-

mediate Fourth of July Creek valley. These remnants are seen

84

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85

as a pitted surface on the floor of Porter Lake valley about
2 miles (3 kilometers) northeast of the juncture moraines.
The main part of this remnant shows up well in the air
photo of Figure 17. This is a boldly pitted terrace of
outwaSh-veneered ground moraine and is graded to the highest
fluvial terrace in the proximal zone. It is considered to
be Atlin III in age, because the Atlin IIIA glacier tongue
contributed previously to the building of the juncture
moraine complex. The apparent sequence of subsequent events
is shown in Table IV. This table relates to the following
discussion, which depicts the total chrono-sequence which
was involved in the fluvial terraces shown in the transect
sketch of Figure 20 and the allied transect in Figure 21.

The highest and first terrace, Atlin Fluvial I (AF-I)
is about 55 feet (16.8 meters) below the moraine laid down
on McDonald Mountain and Caribou Ridge during Atlin IIIA
time. The only remnant of this oldest level as indicated
in Table IV and the cross sectional plot, lies on the eastern
flank of McDonald Mountain. This first level was produced
by a combination of runoff from Atlin ice and upper Fourth
of July ice plus drainage from the local cirque glaciers
at higher levels, in consequence of cirque build-ups when
freezing levels rose substantially in Atlin IV time.

The next terrace was Atlin Fluvial II (AF-II in Table
IVA) which was mapped 20 feet (6 meters) lower. This re-
presents a stage of significant runoff and down-cutting as

the Atlin ice experienced a still-stand at the valley

86

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89

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91

juncture. This resulted in final construction of the massive
morainic-outwash complex in the Ruffner system. This re-
presents the Atlin IIIB glacial stage, and is correlated
with the lower pitted terrace two miles (3 kilometers)
farther down the Porter Lake valley (Figs. 3 and 20). This
terrace is well shown in the aerial photo view of Figure 17.

As downwasting progressed, assuredly a large volume of
water was derived both from melting ice in the upper Fourth
of July Creek valley and from that in the retreating Atlin
ice in the lower Fourth of July Creek Valley. For a short
time, the meltwaters which were dammed in a rather restricted
early Glacial Lake McDonald (Fig. 17) succeeded in draining
to the north. This outflow from the melt-back of Atlin III
ice cut Atlin Fluvial terrace III (AF-III), which we see in
Figure 20 on the west side of the valley. As the Atlin ice
receded farther and Glacial Lake McDonald further enlarged,
the elevation of the lake apparently fell enough that the
AF-III terrace level became abandoned. This resulted in a
drainage reversal in the immediate area, and formed the re-
versed Atlin drainage terrace (RAD). This sequence of
events was accomplished by headward erosion and slumping
as ice blocks which had been buried by the outwash and ice—
cored moraines melted out. Thus, this channel (RAD in Fig.
20) was occupied only during theitime the local ice source
Survived. Evidence for this reversal of drainage is given
bY'the fact that the gradient of the now abandoned channel

153 directed downward toward the southwest.

92

with respect to the foregoing, water draining south-
westward out of the intermediate Fourth of July Creek
Valley (i.e. from remaining Gladys and local plateau ice)
was sufficiently deflected by the juncture moraine that it
curved back to the northeast. As it flowed around the bend,
it formed a meander and cut a new channel, designated in
the transect as Fourth of July Fluvial I (4JF-1). This, in
turn, left an unpaired terrace remnant noted as AF-Z. The
two distinct streams, AF-3 and 4JF-1, then joined and flowed
into the Porter Lake valley. Water levels were by then
quite low and the stream very well entrenched. This means
it cut its way deeply through the west side of the now
densely pitted AF-l and AF-2 terrace seen to the north of
the juncture (Fig. 17).

There are four closely spaced terraces on the east side
of the cross section which are designated as 4JF-1 A, B, C,
and D. These are all associated with 4JF-l on the west side
and the highest, A, is considered a paired terrace of that
level.

Unequivocally, the stages of 4JF-1 represent fluctua-
tions in runoff from what might be termed local ice sources
in the upper Fourth of July Creek valley. It was during
this time, corresponding to the Atlin III C Stage, that
vast amounts of runoff came from the intermediate and upper
Fourth of July Creek Valley. But the thick deposits of ice-
filled drift which still lay in the Fourth of July Creek

Valley, much restricted this flow along the southeast side

93

of the valley and produced terrace remnants on the sides of
Caribou Ridge in the intermediate Fourth of July Creek
Valley (Fig. 21).

Entrenchment of the drainage in GLM (Glacial Lake
McDonald) time is considered a direct result of the severe
downwasting and ultimate collapse of the ice-cored moraine
complex at the juncture. This was then followed immediately
by the drainage of Glacial Lake McDonald (Fig. 5) which
emptied into Porter Lake Valley. The bold and decisive
terrace remnants seen today suggest that this was a short
and catastrOphic episode and one which, in fact, initiated
the present day drainage of Fourth of.July Creek in quite
the Opposite direction. The final channel level in Mc-
Donald Fluvial II time (GLM-II) represents another short-
lived still-stand impounding Glacial Lake McDonald, during
which time it again drained north into Porter Lake Valley.
These neatly paired terrace levels (Fig. 20) correspond to
the glacial condition in Atlin IV time when the then steep
and receding terminus of Atlin ice not only produced the
meltwater for this pro-glacial lake but as well dammed its
runoff thus forming Glacial Lake McDonald. As the ice in
this terminus retreated farther to the Atlin V position,
which was well down-valley and at a much lower elevation
toward the southwest, we may assume that Glacial Lake Mc-
Donald decreased in depth and ultimately drained southwest-
ward into the present lower Fourth of July Creek Valley.

Today's channel in the Fourth of.July Creek thus was

94

Opened. This episode marked the end of all late-Wisconsinan
glacial drainage through Porter Lake Valley. The reason for
assigning a late-Wisconsinan time table to these events is
fully discussed in the following sections on the glacial
chronology.

As for modification of the geomorphology of the Porter
Lake Valley, the terraces and channels described have suf-
fered little change in Holocene time except via the eventual
melting out of ice cores and the final develOpment of a
vegetative cover. The relatively small amount of precipita-
tion in this region, plus the coarse sedimentological texture
which produced good drainage quality in the surface deposits
have prevented any further channeled runoff in these valleys,
except for the presence of an intermittent stream just north
of McDonald Mountain between it and Mt. McIntosh (Fig. 3).

During the sequence of events described in the fore-
going it is apparent that the drainage of Porter Lake was
essentially controlled by the presence of a natural outlet
valley seen on the map (Figs. 2 and 3) as the valley of Tel-
Cabin Creek. It is also clear that early high strand lines
along Porter Lake extended north and south of the shore
limits of the present lake.

Much of the early drainage was also through the lower
and broader outlet valley of Indian Creek, from which in all
probability the main meltwater regressed directly on to the
Atlin Ice in the main Atlin valley, or at least was impounded

against it, forming huge side-valley moats north and south of

95
Halcro (Hitchcock) Peak and Black Mountain (Fig. 3).

During these high ice levels, it is also possible that
some of the meltwater drainage was through the valley between
Mt. McIntosh and McDonald Mountain as well as from Tel-
Cabin Creek, both of which are the next east-west trending
valleys to the north. There is, in fact, geomorphic
evidence that the McIntosh valley outflow drained across
the present water divide. At lower levels, however, the
drainage reversed when ice-cores in the drift of the valley
melted out, resulting in a lower elevation than the present
bedrock divide between Porter Lake and Atlin Lake.

By the time that terraces AF-3 and 43F were develOped
the water had reached a sufficiently low elevation that all
further drainage was westward via Tel-Cabin Creek. It is
suggested that the multiple terraces of 4JF-l time may well
reflect local base-level changes because Tel-Cabin Creek,
via the melting of ice, assumed the main role of drainage
in this sector and so lowered the level of Porter Lake
almost to its present stage. i

The lower Fourth of July Creek lacks post-glacial
notching and intrenchment from fluvial drainage. This is
interpreted as further evidence for diversion of early melt-
water through the Porter Lake Valley followed by the
temporary base-level control of Glacial Lake McDonald.

Thus during the time of most ice melting in the upper Fourth
of July Creek Valley there was no channeled run-off in the

lower valley. Post—Glacial drainage in the time span since

96

has been relatively minor in this arid region.

Moraines and Glacio-Fluvial Terraces in the
Intermediate Fourth of.July Creek Valley

 

The glacial history in the intermediate and upper
Fourth of July Creek Valley is recorded by the presence of
lateral moraines and glacio-fluvial terraces in both the
upper and intermediate valleys. Two composite schematic
transects are presented along lines AB and CD in Figures
21 and 22. The general location of these transects is shown
on Figure 5b (page 38) and also on Figure 3.

The two most prominent moraines in the lower valley are
those of Atlin II and Atlin III ice, each composed of till
and veneered with glacio-fluvial sediments. The higher
lateral moraine of Atlin II age has a distinct channel
between the moraine crest and the valley side of Caribou
Ridge. Subsequent to Atlin II time, there was lateral drain-
age on the north side of the valley as evidenced by pro-
nounced glacio-fluvial terraces on the Atlin III-A lateral
moraine. The final downwasting of Atlin III-A ice is re-
presented by a pitted terrace on the south edge of the valley
floor (Fig. 17). An isolated remnant of this terrace re-
mains in the center of the valley indicating that this part
of the intermediate Fourth of July Creek Valley was once
filled with a much greater volume of drift which was later
flushed out by the vigorous glacial streams cascading as
’torrents out of the intermediate valley.

The Atlin III-A terraces are related to the earliest

97
terraces in the Porter Lake Valley (i.e. AF-l and II and
4JF-l in Fig. 20). Though at least four such terraces can
be traced for several miles, there are other terrace
remnants probably relating to the individual stages of
4JF-l in Porter Lake. Because of slumping and poor eleva-
tion control it is impossible to correlate these terraces
as individual units.

It should be noted too that there are small fluvial
terraces of Holocene age in the present Fourth of.July Creek
Valley. Figure 22 illustrates the drift deposits in the
lower part of the upper Fourth of.July Creek Valley. Here
eskers become prominent as well as the more massive till
deposits associated with the Gladys I and II phases and with
the downwasting of local plateau ice. The south wall of the
valley contains thickly concentrated pitted outwash mantling
a kettle-holed moraine which has been covered with post-
Glacial talus off the flank of the unnamed massif to the
southeast. The northwest side of the valley also contains a
moraine remnant of Atlin II ice which is thought to have been
in contact with the late-Glacial ice-cap glaciation formed
from combined plateau and Gladys Ice. Evidence for this is
a strong linear moraine continuing up valley for several
miles.

The glacio-fluvial terraces of Atlin III-A ice first
appear just above Volcanic Creek. It is mentioned here that
these may be from remaining Atlin II ice combined with

plateau ice when Atlin III ice was in the lower valley.

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There is no evidence on the south valley wall that Atlin
III-A ice made its way this far up valley.
Finally, it is noted that lesser (younger) terraces
in the Fourth of July Creek sequence are also present in

the lower section of the upper Fourth of.July Creek valley.

Esker-Kame Complex in Upper Fourth of.July Creek Valley
The upper Fourth of July Creek Valley is filled with

a large and well-preserved esker complex extending just
above the Spear Point out-wash mantled moraine system for
about 12 miles (19 kilometers) into the Consolation Creek
Valley. The terminal zone of this esker complex is shown
in Figure 11. The esker system forms a reticulated network
of narrow and steep-sided ridges of the classical esker
form, none exceeding a length of 2 miles (3.2 kilometers)
(photo in Fig. 27). In places other and more non-descript
deposits of irregular form cross-cut the eskers, some of the
cross-cutting features being elongated and others not, thus
making a unique group of overlapping water-deposited or water-
worked features. This kind of esker complex has been
described by Armstrong and Tipper (1948) as compound eskers
and later by Tipper (1971) as esker complexes. In this

area they have sometimes been referred to as esker-kame
"swarms" (Miller, 1975a). The axes of the ridges and en-
closed depressions follow the elongate form of the major
eskers in the system and are sub-parallel to the present
valley drainage.

The esker "swarm" was laid down in meltwater channels

100
in and under the downwasting ice in the upper Fourth of
July Creek Valley. The major continuous eskers, diagram-
matically shown in Figure 19 reveal two directions of
tributary flow with a small non—esker zone between. This
delineates a subglacial drainage divide some distance to
the west of and below the present day water divide.

The morphologies of the ridges in the esker complex
are divided into three main kinds. The first represents the
classic englacial or subglacial esker and in this region I
these are less than 25 feet (7.5 meters) high and not more
than 20 feet (6 meters) wide. They are composed of very
coarse and sub-stratified sand to boulders. Almost all of
the clastic material shows the strong effects of fluvial
sorting, water-scouring and erosion.

A second type is a larger esker-like ridge, 30 to 50
feet (9 to 15 meters) high, containing a higher percentage
of sand and gravel-sized sediments, hence with much less
coarse material. It is thought that these particular
"eskers" were probably formed in Open channels between ice
walls or as water-laid deposits in wide crevasses during
the waning or dead-ice phase of the plateau glaciation.
Technically these should be considered as kames.

The third type of morphology involves smaller ridges
which are usually not parallel to the main esker axes in the
system. These are interpreted as crevasse fillings. They
have textures more like the larger Open channel ”eskers" or

kames. The probability is that these are not totally

101
stratified features, but formed as ice-contact deposits
which suffered much slumping and redistribution during the
lowering of the bottom of crevasses in the final stage of
melting.

The area north of Gladys Lake also contains a series
of esker complexes similar to those in the upper Fourth Of
July Creek Valley. Here, however, the esker "swarms" are
quite linear with fewer overlapping non-descript forms such
as the irregular kames and crevasse fills found in the
plateau section. These, however, are related to the main
meltwater channels of Gladys IV ice during final down-
wasting and decay of the ice.

Since esker complexes form in downwasting ice they
cannot be associated with vigorous movement. Therefore,
this interpretation is in order, i.e. when we find such an
array of deposits it indeed connotes wasting ice, changing
into a dead-ice phase. For this reason the tributary rela-
tionship, in the case of the upper Fourth of.July Creek
Valley eskers, has been helpful not only in determining
englacial and subglacial drainage but, indeed, has served
to delineate the final downwasting of plateau glaciation in

Gladys IV and Atlin IV and V time.

Valley Asymmetry and Abandoned Channels
in the Gladys Lake Depression

 

Asymmetrical valleys are common in the study area,
particularly on Gladys Mountain. As such they represent a

unique study in themselves, for the asymmetry appears to

102

controvert the law of divides. The asymmetry concerns
angle of repose, relative erosion and slope stability. In
the Gladys Lake depression, the abrupt beginning of some
tributary streams On the north lepe suggests that a thin
mass of ice remained on these lepes for some time after
the main glacier surface had lowered (e.g. see upper Davené
port Creek, Fig. 23). From this, meltwater runoff cut
through thick till deposits, resulting in deeply incised
channels. The nature of the drainage indicates interest—
ingly enough that many of the gulley streams washed out onto
the downwasting surface of the Gladys IV stage glacier in the
Gladys depression. The evidence for this is the sharp trun-
cation of Gladys IV ice by the pitted moraine i.e. the strong
presence of pit degression surrounded by stratified material.

It is also of significance that the south-facing lepes
of these small streams are steeper than the north-facing
lepes. Being close to the angle of repose, these south-
facing lepes do not support much vegetation. Surprisingly,
they have not been susceptible to extensive delevelling by
mass wastage. The reason may be the permeable nature of
the drift which is comprised mainly of sands and gravels
through which saturating ground water readily moves. Also
because of differences in insolation, the north-facing
lepes are much drier and less affected by cryoturbation
and other forms of mass wasting. Even where the outside of
a meander has a north-facing lepe, it retains a much

gentler gradient than the south-facing cliff above the in-

103

Figure 23. Head of Gladys Lake

Note Consolation Creek flowing into Fish
Lake and Davenport Creek flowing into
Gladys Lake.

Canadian Energy, Mines and Resources
A12106-90, 1949

 

105

side Of meanders. Low annual precipitation of this region
abets the lack of vegetation on the south-facing lepes
where, of course, available ground water is slight and
summer evaporation great. This has resulted in relatively
denser vegetation on the north-facing lepes which have
now stabilized their gradients to relatively gentle relief.

Another manifestation of marginal precipitation allied
to lepe gradient is found on many of the larger eskers,
and on the crevasse fillings and non-descript kames in the
upper Fourth of July Creek Valley. Here again, steep south-
facing lepes are less vegetated and north-facing slopes
gentle. All of these features too are composed of highly

permeable sand and gravels.

CHAPTER VII
PERIGLACIAL ENVIRONMENT

Periglacial processes have played a large role in the
evolution of Holocene geomorphic features in the area of
consideration. Some processes are active today while
others are revealed by relict forms that show no evidence
of present day or even Little Ice Age activity, that is

for at least the past four to five centuries.
Solifluction Zones

Frost action and mass wastage have greatly modified
valley flanks in the Fourth of July Creek area. Although
frost solifluction is still an active process in the
tundra zone above timberline (c 4000 ft., 1212 meters) and
at higher elevations, talus on steep lepes is common. At
lower levels the talus is heavily lichen-covered, indicating
relative inactivity and stabilization today. Anderson's
study (1970) of Holocene environments in the Atlin region
(Table II) defines a relatively cool/dry period 1000-2500
years B. P., the early Neoglacial. This is the classical

condition of a periglacial climate.* Thus it is presumed

* The dry/cold periglacial condition of the Atlin re-
gion is the antithesis of the wet/cold glacial condition
typifying most of the Juneau and Stikine Icefields and the
Alaskan Coast.

106

107
that frost action and associated mass wasting processes, as
well as other periglacial processes, were intensified during
this time.
Today no evidence of solifluction is found at eleva-
tions below timberline. Above this zone (4000 feet, 1212
meters) numerous solifluction terraces and associated flow-

lobes are present.
Nivation Hollows

NO semi-permanent snowfields or firn patches are
found in the drainage basins of Fourth of July Creek or
Consolation Creek valleys. One such local firn patch,
however, occurs on Ruby Mountain at an elevation 6000 feet
(1830 meters). This pocket of residual firn, lies in the
lee of a col and its persistence through recent years makes
it, in effect, a glacieret. An abandoned addit of a tungsten
scheelite prospect about 400 feet (120 meters) below this
glacieret contains large stalactites of ice up to 2 feet
(0.6 meters) long and 1 foot (0.3 meters) in diameter. Some
16 feet (5 meters) within the addit, ice crystals extend up
to 2 feet (0.6 meters) from the walls and ceiling. Large
hexagonal crystals of sublimation ice reach 6 inches (15
centimeters) in diameter. Clearly this is in the perma-
frost zone, i.e., above 5500 feet (1575 meters).

As we have seen in the foregoing discussion of climatic
parameters, at the higher elevations of this region the mean

annual temperature is well below freezing, thus producing

108

apprOpriate conditions for permafrost develOpment below
the surface and for nivation processes at the surface.
In fact, at the permafrost seep level, about 5500 feet
(1670 meters), the mean annual temperature, extrapolated
from a prevailing lapse rate of 3.30F/1000 feet, is 21°F
(-9°C). Therefore, in this region a minor climatic
deterioration can readily produce a re—occupation of high—
level nivation hollows and cirques.

within the study area, a number of nivation hollows
are present at elevations down to 5000 feet (1525 meters)...
i.e., on Mt. Vaughan (Fig. 24) just above the juncture of
the Fourth of July Creek Valley and the valley to Porter
Lake. These hollows are strikingly apparent because of
the lack of lichen on rocks formerly covered by snow and/Or
ice. In view of the very fresh surfaces in the bare rock
areas and rubble zones eXposed by de-nivation it is apparent
that the removal of firn from these nivation hollows does
not relate in time to Anderson's Zone II (Table II) but
rather to a very recent warm period following a general
cooling. As this condition is cyclic and the climate is
now returning to a colder phase (Miller, 1972; Miller and
Anderson, 1974a), these sensitive firn patches and glacierets
should be monitored. They will certainly reSpond to the
continuing cooling, and in fact already there is re-deveIOp-
ment of nivation patches in many of the recently-bare depres-
sions at and above timberline, a trend currently observed on

the Alaskan Coast as well (JIRP and also the U. S. Forest

109

 

Figure 25. Relict stone circles in upper Fourth of July
Creek valley. Photo by M. Miller

110
Service, communication from the Forestry Sciences Labora-
tory in Juneau).

It is of regional significance, therefore, that even
in the last five summers an increasing number of snow
patches have been observed on the hills east of Atlin, with
the late—summer firn being retained as a base for further
accumulation in the on-coming autumn and winter seasons.

A number of sorted and unsorted stone circles, stone
stripes and stone polygons are found in the study area.
These are both as relict forms and as ones active today.

The best develOped relict stone circles are found in the
intermediate Fourth of July Creek Valley at an elevation

of about 3200 feet (975 meters) (Fig. 25). These circles,
about 3 feet (1 meter) in diameter, are now often under
water during the summer and are positioned well inside of

the swarm of eskers shown in Figure 19. They are considered
to be relicts from the earliest Holocene (ca. 7000-9000 years
B. P.) but they also may have been affected by cooler, drier
climate of the early Neoglacial 2500 years to 1000 years

B. P. The rationale for an early Holocene age assignment

is the fact that similar relict stone circles have been
found at 2500 feet (760 meters) On the Alaskan coast and have
been with some confidence assessed as very early Holocene or
late-Glacial in age (Miller, 1975a). Regardless of vintage,
it is clear that increased temperatures and higher water
tables succeeded in terminating the growth of these forms.

At elevations above 5000 feet (1524 meters) quite

111

active and well—defined circles are found. The circles at
or above the 5000 feet elevation, are seen to contain frost-
heaved particles no larger than pebble size. Most of these
circles average a half meter in width. Those at 6000 feet
and above are 4 to 5 (1.2-1.5 meters) across and have been
seen in abundance from the air. Details of the texture and
morpholOgy of these features where they occur in the study
area are not yet known, however, some studies have been
carried out on comparable features at 5000 to 7000 feet
(1524 - 2121 meters) elevation on Teresa Island near Atlin
(Buttrick, 1974) and at Camp 29 on the Cathedral Massif west

of Torres Channel (Jones, 1975)
Palsas and Related Frost Moungp

The Atlin district lies directly south of the dis-
continuous permafrost limit (Brown, 1970). In spite of
this, perennially frozen ground has been found to be common
in poorly-drained bOgs at elevations considerably below the
permafrost seep level. For example, in the study area this
is well evidenced by the presence of circular to elongate
mounds of peat and/Or mineral sediments containing a
partially frozen core. It is of special interest that these
mounds are found as much as 2000 to 2500 feet (610 - 670
meters) below the local permafrost level (seep level). The
common occurrence of frost mounds puts the Atlin district in
the sporadic permafrost zone.

Frost mounds of the type common to the study area have

112
been investigated extensively in Scandinavian countries and
are given the Finnish name palpa (Maarleveld, 1965; Seppala,
1972a and 1972b; and others). Larger areas of frost mounds
showing raised relief are referred to as palsapplateaus or
plate palsas (Kujansuu, 1969). Several Of these are found
in both the intermediate and upper Fourth of.July Creek

valleys (Figs. 26 and 27).
Origin of Palsas

The origin of palsas and palsa plateaus is dependent
on temperature, water supply and amount of snow cover as
well as on physical properties Of the sediments involved.
The process involves a peat bOg which completely freezes
over in the winter and forms domes from ice lenses and
frost action at random locations below the surface of the
bOg.- Once the bOg water freezes this freezing extends
into the saturated peat and underlying silt. The raised
peat mound then grows and becomes drier. Containing many
air pockets in the vegetation it also becomes a better
insulator to atmospheric heat in the summer months.

The initial exposure of peat and its subsequent drying
seems to be the most critical step in palsa formation.
Ideally increased precipitation in the fall and decreased
evapO-transpiration leads to increased moisture in the
raised mound in the winter. Because of wind scour, snow is
less likely to accumulate on the raised palsa mounds than

the surrounding bog, thus decreasing its insulation in the

113

Figure 26. Palsas in the upper Fourth of
July Creek Valley '

Figure 27. Eskers and peat plateau in the
Fourth of July Creek Valley

114

 

115
winter months and actually causing an increased cold
penetration, with consequent enlargement of the extent of
frozen ground and its allied ice inclusions. Resulting too
is a consequent increase in palsa height. As the palsa
grows, the peat and/or mineral soil breaks away from the
sides and eventually forms large cracks across the mound
sometimes exposing the frozen core. The height a palsa
reaches and the length of time it can exist are functions
of the peat thickness and the consistency of the peat or
mineral sediments exposed, as well as of the climatic
conditions.

Palsas commonly occur in bOgs where the mean annual
temperature is less than 33.6°F (1°C). They are recorded in
northern Sweden where the mean annual temperature remains
below 32°F (0°C) during more than 200 days per year and are
present only where precipitation during the period of November
to April is less than 14 inches (31 centimeters) (Lundquist,

1962).

Character of the Atlin Palsas

 

Three prototypical localities lie in the study area
and there are many others in the Atlin region e.g. at the
head of O'Donnel River (Fig. 3). The first locality dealt
with in this investigation is the bog sector at the drainage
divide of Fourth of July Creek and Consolation Creek. Here
there are two well-develOped and active palsas and two

sectors of raised peat plateaus (Fig. 26 and 27). The

116
extent and depth of sub-surface ice in these frost mounds
were studied using the electrical resistivity technique.
Specifics of this geOphysical method are described later.
Elevation of this particular palsa area extends from 3300
to 3400 feet (1005 to 1035 meters).

Another sector was investigated one mile (1.6 kilo-
meters) southwest of the highland zone and at about 50 feet
(15 meters) lower elevation. Here several small frost
blisters were studied (Fig. 19) which appeared to represent
the beginning of palsa formation. Here the ground was
frozen 8 to 12 inches (25 to 30 centimeters) below the sur-
face for a thickness of 10 to 12 inches (25 to 30 centi-
meters). These mounds are not considered to be remnants
of deflated palsas as there is no evidence of the charac—
teristic vegetation of the well-drained and mature palsa
surface. Observations over a three-year period have shown
only minor fluctuations in the depth of permafrost in these
palsas and in their heights above water table. One year
when the bOg water level was unusually high the frozen
material was much less in both aerial and vertical extent.
This points out the critical affect the water table has on
palsa formation. As freezing takes place at or near the
water table, a rise of the water table can easily destroy
small palsas.

The third palsa locality is still within the esker
complex but about 1% miles (2.4 kilometers) farther down

the valley, and another 50 feet (15 meters) lower in

117

elevation than the frost blisters. Here a very large and
well—develOped palsa was found to be raised about 10 feet
(3 meters) above the bog surface. Its size was 150 feet
(45 meters) by 40 feet (12 meters) (Fig. 19). The depth to
permafrost was variable but averaged 15 inches (38 centi-
meters). Active sloughing was seen to be occurring in
places on the sides of the palsa but much of the flanking
area was well vegetated indicating that vigorous active
growth is not at present taking place. The fact that all
of these palsas lie well inside of the late-Glacial esker
swarm connotes an age of less than 10,000 years. Confirma-
tion of this is provided by C-14 samples discussed later.

The electrical resistivity method was used to deter-
mine the depth and thickness of permafrost in the highest
palsas and peat plateaus of the water divide sector described
above.

Electrical Resistivity Assessment
of the Fourth of July Creek Palsas

The freezing of interstitial water has little or no
effect on the density, magnetism, and radioactivity of rocks
or soil, but this can result in large changes in the velo-
cities of compressional seismic waves and in the measured
electrical conductivity characteristics (Timur, 1968). As
seismic and electrical geOphysical methods have been used
extensively in permafrost delineation (Barnes, 1963), the
electrical method was applied in this interpretation.

Before considering the geOphysical data some basic

118
considerations are reviewed. First, the resistivity or
inverse conductivity of sediments is a function of the
amount and electrolytic nature of water present in the
interstices. It decreases as the water content increases.
Conversely when water freezes in the sediments, there is a
marked increase in resistivity. Also the resistivity of
pure ice increases with a decrease in temperature. Keller
(1966) has shown that the resistivity of a porous sandstone
increased by 184 times when the temperature is decreased
from 20°C to —12°C.

In frozen sediments, the salinity of the water also
has a complex effect on resistivity. Salts lower the freezing
point and when freezing begins salts migrate to unfrozen
water, thus increasing the salinity there and decreasing
the resistivity of the liquid water present (Keller and
Frischknecht, 1966).

Resistivity techniques have been used to map upper
permafrost surfaces (MacKay, 1970). The depth to perma-
frost can be determined quite accurately because the primary
control of resistivity is the change of water to ice at the
permafrost-active layer interface. Frischknecht and Stanley
(1970) were able to get good depth soundings in areas not
seriously affected by lateral variations and the thickneSs
of permafrost was determined where the current completely
penetrated the frozen ground. Because of the multiple con—
trolling variables within the permafrost (i.e. relative

amounts of ice, the electrolytic nature of water and

119

temperature) resistivity measurements cannot be used for
positive identification of sediment type at depth. Only
with ample direct Observational data can the real nature of
sediment character be inferred from resistivity measurements.

In areas of relatively thin and Sporadic permafrost,
the resistivity method offers a quick and inexpensive re-
connaissance method not only for mapping permafrost surfaces
but also for determining thickness of the sub-surface ice.
It is particularly useful in palsa investigations (Tallman,
1973).

within the highest swampy area of the upper Fourth of
July Creek Valley, the most active palsas have a height of
about 10 feet (3 meters) above the bOg surface (Fig. 26).
Each palsa surface is covered with about 24 inches (60 centi-
meters) of peat and the depth to permafrost varies from 12
to 30 inches (30 to 75 centimeters) where frozen peat or
silt and clay is encountered. At about 4 feet (1.2 meters)
below the surface, lenses of ice 2 centimeters thick-are
common and the quantity and thickness of ice lenses increases
with depth. Excavation was only made to a depth of 6 feet
(1.8 meters) from the palsa surface and did not reach the
water table.

Also present is a larger and well-defined peat plateau
(Fig. 27). This had a broad flat surface 6 feet (1.8 meters)
above the general level of the bOg. This plateau was found
to have a thinner layer of peat ranging from 6 to 12 inches

(15 to 30 centimeters). Here the depth to permafrost was

120
12 to 18 inches (30 to 45 centimeters). In the upper four
feet (1.2 meters) it lacked the isolated lenses of ice found
in the active palsa.

Other slightly raised permafrost areas, with less
well-defined limits and more undulating surfaces, are also
present in the investigation area. These areas are pro-
minent because of the distinctive vegetation on the well-

drained permafrost sites.

Details of the Resistivity Method and Computational Technigpe

The resistivity measurements were made using an ex-
panding Wenner configuration (Fig. 28) with four equally-
spaced electrodes. The potential difference was measured
between two points on the ground, B and C. A center elec-
trode (Lee partition) was added to measure the potential
difference between B and E and C and E, or the right and
(left half of the array (Fig. 29). These additional readings
permitted comparison of the apparent resistivity on either
side and also served as a check on accuracy of the measure-
ments. In general, little variation was found in the left
and right half of the configuration, and only the apparent
resistivity in the full circuit is plotted for analysis.

A direct current instrument was used which gave the
value of ¥ where V is the voltage across the potential
electrodes, and I is the current flowing through the ground.
Both forward and reverse current readings were taken, and the

average value was used to calculate the apparent resistivity

121

 

 

WENNER ARRAY

 

 

 

 

77

l-—->as—-l——->o<—-l——>o<———l

 

A B C 0
Current Potential Potential Current
electrode electrode electrode electrode

Figure 28. Nenner Array

L EE PARTITION

 

 

 

 

I V Val

A " a E c D

 

Figure 29. Lee Partition

 

 

122

from the formula:

A=2na¥

where: A apparent resistivity

a = "a” spacing of Wenner array

= 3.1416

and: X = Eggggglglfiglifigggggg’ read directly from
I cur e the meter.

The "a" Spacing is assumed to be the depth of current penetra-
tion.

In all, six profiles of apparent resistivity were
plotted against "a" spacing. Their general location is
shown on Figure 30. Profiles 1 and 2 (Fig. 31) are in
locations where there is no surficial evidence of permafrost.
Profile 1 is across a well-drained area of coarse sand and
gravel. Profile 2 is across the bOg with a surficial layer
of damp to saturated peat. The apparent resistivity was
greater in the well-drained profile but, with increasing
depth the two profiles were very similar. It must be kept
in mind that apparent resistivity at depth is an average of
the resistivity of the entire sub-surface from ground to
depth of current penetration. There is no indication of a
hurried high resistance layer in either profile.

Profile 3 (Fig. 32) was made across an active palsa
mound about 120 feet (36 meters) long. The Wenner array was
expanded from a 2 to 40 feet (0.6 to a 12 meter) "a" spacing.
By stOpping at the edge of the palsa, complications due to

change in relative elevation and the effect of electrodes in

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125
the surrounding bOg water was eliminated. The sharp de-
crease in apparent resistivity in profile 3 indicates that
the limited "a" spacing was sufficient to penetrate com-
pletely the layer of high resistance.

Profiles 4 and 5 (Fig. 33) were both made diagonally
across the highest peat plateau in this sector. Again, it
was possible to expand the array to the limits of the mound
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high resistivity layer. It should be noted that apparent
resistivity was greater in profile 5 indicating a higher
resistivity at depth and/or a greater thickness. A high
resistivity could be due to differences in temperature as
well as the quantity of ice at depth.

Finally, profile 6 (Fig. 32) was across a slightly
raised hummocky area. Test-pits dug here showed a wide
variation in the thickness of the active layer and also

considerable lateral variation in sediment texture.

Interpretation of the Geophysical Data on Sub-surface Ice

Profiles 3 to 6 all showed high resistance layers at
depth. The maximum apparent resistivity in each profile
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the actual resistivity at depth but it is also dependent on
the ratio of the thickness of the frozen layer. As succes-
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127
greater change in apparent resistivity than it would
deeper in the substrate.

The apparent resistivity values were plotted against
"a" Spacings on lOg paper and compared to Mooney and
Wetzel's (1956) theoretical curves for 1, 2, 3, and 4
layers. Because of the gradational variation of resis-
tivity with depth in the frozen layer, it was not possible
to match the curves accurately.

The method employed for interpretation was the Barnes
layer method (Barnes, 1954), devised to minimize the masking
effect of the overlying material. Assuming an "a" spacing
of 3 feet (1 meter) the first increment measures the resis-
tivity of the volume of material to a depth of 3 feet (1
meter). The 6 feet (2 meters) increment measures resistivity
to a depth of 6 feet (2 meters) and includes the previous 3
feet (1 meter) increment plus an additional 3 feet (1 meter)
layer. This can be considered as two resistors in a parallel
circuit where the conductance of one resistor is known (the
tOp 3 foot (1 meter) increment) and the total resistance is
known (the 6 feet (2 meter) increment). It is thus possible
to solve for conductivity of the bottom 1 m layer by the

following formula:

....1...=l. -l
Rn Rn Rn_1
where: %“ = layer conductance of a given increment,
n in mhos
% = total conductance between ground surface

n and bottom of given increment, in mhos

128

and: l = total conductance between ground surface
Rn—l and the bottom of the increment just above
given increment, in mhos

Layer resistivities were calculated for all six pro-
files using the smallest "a" spacing available. These are
plotted in Figures 34 to 36 at the mid-point of each layer.

The layer resistivities calculated for profiles 1 and
2 differed very little. Profile 1 (Fig. 34) showed a
slightly higher resistivity value near the surface due to
drier surficial material. Only profile 1 is plotted and
used to aid interpretation in profiles where permafrost is
present. The sharp decrease in resistivity in the 2 to 4
feet (0.6 to 1.2 meters) layer is interpreted as delineating
the water table in this high bog sector.

Profile 3 (Fig. 35), which transects the most active
palsa is plotted 9 feet (3 meters) above the bOg surface.
Here the maximum resistivity is above the bog surface and
falls off to a value comparable to that of unfrozen sedi-
ments in this area at a depth less than 6 feet (2 meters)
below the bOg surface. This suggests that freezing is
currently taking place at or slightly below the water table.

The layer resistivities of the peat plateau (profiles
4 and 5) are shown in Figure 36. Again, the resistivities
increase rapidly below the first 2 feet (0.6 meter) layer
and remain within the frozen ground range to a depth of 55
to 70 feet (17 to 21 meters). The wide variation of
resistivities within the permafrost could be a functiOn

of differences in porosity, temperature variation, pockets

 

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130
of unfrozen ground, and/or isolated ice masses. The sub-
stantially greater thickness of permafrost in the peat
(plateau sample suggests a different genesis than that of
the round palsa denoted in Figure 35.

The final profile (6 in Fig. 34) again shows a marked
increase in resistivity but at the 4 to 6 feet (1.2 to 1.8
meters) layer. Here the resistivity at depth was not as
great as in the other permafrost profiles but was still far
greater than in non-permafrost areas, and well within the
range of resistivities of standard frozen ground. The wide
variation in resistivities within the permafrost range is
considered to reflect the uneven distribution of frozen
sediments found in the pits dug. The smooth decline of
resistivities, however, makes it more difficult to determine
the lower permafrost boundary. Comparison with the non-
permafrost profile suggests the boundary to be between 30
to 40 feet (10 and 13 meters) below the main bog surface.

In general it is known that resistivity ranges within
permafrost vary greatly. It has not been the purpose Of
this study to try to correlate resistivity ranges with the
specific sediment type with which we are here concerned,
but rather to delineate the depth and extent of the sub-
surface freezing in this sector at considerably lower
elevation than the regional permafrost level. In a geomor-
pholOgical sense, one usually has a knowledge of the under—
lying sediments, as we do here, so that this helps the in-

terpretation.

131

It is now clear that resistivity readings taken in
non-permafrost sectors can add useful information about
the substrate and be valuable to the interpretation of
depths of localized permafrost profiles.‘ With this in-
formation, we now have a clearer understanding of the
delineation of sporadic permafrost pockets in this dis-
continuous permafrost zone at intermediate elevations in

the Atlin region.

Significance of the Atlin Palsas and Related Considerations

The plate palsas or palsa plateaus described and dis-
cussed in this study are considered to be examples of relict
permafrost. Their presence in this region suggests that
they are likely the last remains of what must have been a
far greater extent of permafrost. In the same vein, with
the present short-term cooling trends to the end of this
century (Miller, 1972), they appear to reflect a re-expan-
sion of the permafrost zone.

That actively growing palsas are forming at this
elevation and latitude under present climatic conditions
is indeed significant, as no palsas have heretofore been
reported at or below this_parallel in the Canadian sub-
Arctic. Also made clear by this study is the fact that the
initial exposure of peat by differential upwarp is a critical
step in palsa formation. Furthermore, these observations
suggest that the level of bOg water tables in palsa areas
is extremely impOrtant to the genesis of such features.

For example high water tables may prevent sufficient

132
drainage in the upwarped peat and result in rapid melting
of subsurface ice in summer. If the water tables are too
low there is not sufficient water transport through the
sediments for the thin interstitial ice layers and thicker
ice lenses to form.

Textural variations within such bOgs may also exert
an important control over the transport of capillary water
and hence over the initial differential upwarp. In this
sense that part of the study carried out to determine the
textural variations in unfrozen bog sediments and palsa
sediments were somewhat inconclusive. In a follow-up
prOgram it is anticipated that more concentrated research
aimed at specific characteristics can be accomplished on

this interesting aspect of palsa formation.

CHAPTER VI I I
A WISCONSINAN CHRONOLOGY
Keys to the Interpretation

These investigations in the Atlin area reveal three
major glacial phases followed by three lesser phases in the
interior region of the Alaska-Canada Coast Range. Each re—
presents a significant and identifiable glacio-climatic
event which took place during Wisconsinan time at the
northern end of the Cordilleran Ice Sheet. undoubtedly,
there are others that have not been identified, but the
main lines of the sequence seem clear. It should also be
noted that these glaciations are the result of generally
contemporaneous accumulations in two distinctly separate
but contiguous regions, one in the Cassiar Range and its
flanking plateaux* to the east Of the study area and the
other in the Boundary Range to the south and southwest.
DiscussiOn of the reCOgnized phases of glaciation and their
designations into stages follows. The key to understanding
the complex interrelationships and to the interpretations
which follow is the reCOgnition and appreciation of the
different provenances of the ice involved. A conceptual

model for this situation is considered in the final chapter

* The Teslin, Kawdy and Taku Plateaux. (Bostock, 1948).
As a part of the greater interior plateau.

133

134

based on certain conclusions derived from the long-term
glaciological investigations which have been carried out
on the Juneau Icefield by JIRP personnel.*

The provisional chart of the Wisconsinan and Holocene
stratigraphic sequence and glacial chronolOgy in the North-
ern Boundary Range, Alaska-Canada is presented in Table V
in the pocket of this volume. It should be appreciated by
now, that in the study area many details and evidences of the
actual chronology involved have been eradicated by the inten-
sity of later events. The outline does, however, provide a
helpful guide to understanding the complex glacio-geomorpho-
lOgies of the three segments of the Fourth of July Creek
Valley with which we have been particularly concerned, and
which are so well distinguished by different glacial and
fluvial limits. The chart is also most useful in delineat-
ing the Pleistocene chronology of the Atlin region at large.
Conversely, study of the geomorpholOgy of both the smaller
and larger areas has abetted develOpment of the regional
chronology.

For some key interpretations, evidences from outside the
immediate study area and even beyond the Atlin region are in-
voked. Thus an effort is made in Table V to compare the
chrono-sequence with those found in the Taku District
immediately to the south and west and in other adjacent
sections of British Columbia, the Yukon and Alaska. Lastly,

some teleconnectional inferences are drawn through comparison

* The Juneau Icefield Research PrOgram.

135
with the Wisconsinan chronolOgy of the Rocky Mountain region,
the Puget Sound region of Washington State, and the classic
mid-continent region of the Great Lakes. The Puget Sound
chronolOgy is perhaps of more direct significance in that
it is to the southern end of the Cordilleran Ice-sheet what
this study and its suggested chronolOgy are to the northern
end of the Cordilleran Ice-sheet. Allied, Of course, with
this is the Taku District chronolOgy, including the inter-
pretations from the Tulsequah and southern Atlin region

(Miller, 1956).
Pre-Atlin I Glaciation

An early Wisconsinan glaciation is documented by well
weathered till and relatively deep glacial soils on the
higher summits within the Fourth of July Creek study areas.
In addition, a peat horizon has been dated at I 31,000 C-l4
years B. P. in the valley of Boulder Creek on the south
flank of Mount Leonard (Fig. 3) (Radiocarbon, 1976; also
Institute of Marine Sciences, University of Alaska Radio-
carbon Dating Lab, sample AU-59, 1975). In this connection
Miller (1975a) has reported a C-14 date Oftt36,000 on wood
fragments at a bedrock-till interface in McKee Creek 10
miles south of Atlin. These organic fragments are beneath
the lowest weathered till in the Vesnover Mine and probably
can be correlated with the Boulder Creek peat noted above
(see Radiocarbon, 1976; and University of Alaska, 1975,

sample AU-114).

136

Beneath the above-noted peat horizon is well-
weathered glacio-fluvial and till material which is pre-
sumed to correlate with tills found as ground moraine on
the higher elevation lepes and ridges in the immediate study
area. Also erratics have been found on summits high above
the substantiated upper glacial limits of subsequent ad-
vances. Added to this is the strong develOpment of tors
on high ridges in the Atlin region, such as on Sentinel
Mountain at 6300 feet (1920 meters) elevation in the adja-
cent area south of the Atlin townsite (Fig. 3). All of this
attests to the relative antiquity of this extensive early
glaciation. From the Boulder Creek and McKee Creek C-14
dates and the general degree of weathering on the highest-
level glacial deposits, the pre-Atlin I phase is considered
to be early Wisconsinan, and probably greater than 40,000
years B.P. Specifically it is suggested as probably lower
Altonian in the mid-continent chronology. This interpreta-
tion is consistent with the Alaska composite of Péwé (1975)
and the Yukon Shakwak glaciation of Denton and Stuiver (1967)
noted in Table V. It is also consistent with Oxygen-isotOpe
ratio trends over the past 100,000 years revealed by deep
ice cores in the Greenland Ice-sheet (Langway, et al, 1973,
and Dansgaard, et al, 1973).

Much emphasis has been put on the Boulder Creek and
McKee Creek sites in the age assignment of this phase.
Therefore, in the interest of scientific integrity it must

be kept in mind that these dates are considered to be near

137

the maximum range of C-14 dating and, indeed, that the pre—
Atlin I phase could be even older than suggested.* On the
other hand the amount of weathering on high—level drifts in
the Fourth of July Creek area is much less than would be
expected if the drift had been exposed for the entire
Sangamonian. This may be explained, of course, by the
presence of cool-dry conditions in this region even during
Sangamonian time, conditions which may have impeded or
slowed down weathering rates. Though the character of the
Sangamonian in this region is not well understood, it is
nevertheless assumed that there must have been long periods
of warmer climatic conditions that would have produced more
extensive weathering than is found in the cited drifts.

Comparison with the Taku District Chronoquy in the Alaskan
Panhandle and Southern Atlin Area

 

Of particular significance to the study area is the
Quaternary glacial record or the Juneau Icefield and its
peripheral sectors which were the source area for the Atlin
Ice. On the basis of extensive and moderately well-weathered
drift and recognized cirque sequences, the pre-Atlin phase
is correlated with Miller's (1956) early Atlin stage (Early
Juneau/Atlin stage in Miller, 1975a). The C1 level cirque
system, Table III (also see Miller, 1961), is suggested to
have been initiated in late-Illinoian or early-Wisconsinan

* Further samples of the critical Boulder Creek peat
are currently undergoing isotOpic interpretation by Geochron

(Kreuger Enterprises) in Cambridge, Mass., and will be re-
ported upon in the final publication of this report.

138

time. Evidence for interglacial climates on the Juneau
Icefield have been destroyed by later glaciations and the
only real evidence for an interglacial following this stage
is a change in elevation of active cirques in consequence

of a rise in elevation of the freezing level (a warming
condition). This interpretation is somewhat restricted in
that cirque levels were occupied at many different times
during the Pleistocene and their Wisconsinan initiation is
difficult to correlate as to time, thus forcing the analysis
to emphasize only the sequential develOpment. For these
reasons the lowest cirque level is correlated to the most
extensive Atlin region glaciations and assigned an age on
the basis of known C-14 datings. These refer not only to
the Boulder Creek/McKee Creek information in the Atlin
region, but also to pre-Woodfordian C-14 dates reported from

the southwestern Yukon region by Denton and Stuiver (1967).

Time of Onset of Wisconsinan Glaciation

Because of the mountainous character of this region it
is probable that the early Wisconsinan glaciation began here
before the end of the classical mid-continent Sangamonian
age (ca. 70-90,000 years B. P.) and that it existed for a
longer period of time than in more southerly regions (per-
haps even up to 30,000 years). Such an earlier onset of
glaciation in this northern Cordilleran region is suggested
in view of the earlier accumulation at higher elevations as
freezing levels fall. It is also suggested because of the

higher latitude position - i.e. in this sub-Arctic region.

139

Details for fluctuations within this phase would have been
removed by subsequent glaciation, as noted diagrammatically
in Table V, but by inference it is believed that major
glacials and intraglacials undoubtedly existed during this
early Wisconsinan time frame. This is also consistent with
the suggested divisions of the Wisconsinan derived from
Greenland ice cores (Dansgaard, et al, 1969, 1971).

Very little work that can be correlated with Gladys
Ice has been done in the region east of the study area. For
the north central British Columbia area Rutter, (1975) has
documented a very early advance, which he suggests as possibly
pre-wisconsinan. This may actually be a correlate of the
early Wisconsinan phase shown here for the Atlin-Taku
region. It should be noted, thought, that in the area of
Gladys drift evidence for a pre-Atlin I phase has been
completely removed or covered by the effects of extensive
Gladys I glaciation.

Regardless of details, it is certain that the Greater
Mountain Ice-sheets which were involved in the pre-Atlin I
glaciations were the most extensive ice sheets which have

affected the Atlin region during Wisconsinan time.
Boulder Creek Intraglacial

The evidence used in this study that there was a sub-
stantial retreat of ice from the Atlin region sometime
between 35 and 50 thousand years ago is based on the cited

Boulder Creek/McKee Creek C-l4 dates, abetted by the C-14

140
dates of Denton and Stuiver (1967) relating to the Silver
non-glacial of the southwestern Yukon Territory. That this.
stage actually existed is not questioned, only its precise
date and details of related intraglacial variations. The
dating would be less tentative if more locations with
organics had been found.

A tentative correlation with the Port Talbot Intra-
glacial (Flint, 1971; Terasmae and Dreimanis, 1975) of some
50 to 60,000 years B. P. in the Great Lakes region and
Easterbrook's (1967) corresponding Salmon Springs Inter-
glacial (Intraglacial would be a more apprOpriate designa-
tion) of the Puget Lowland is made on the basis of the above-
cited radio—isotOpe ages. Miller (1975a) has also reCOgnized
the existence of this intraglacial on the Juneau Icefield
(see Table V) and for that sector considers it to have been
of short duration and relatively cold. He has suggested
this on the basis that sub-Arctic mountain environments have
been involved throughout the Pleistocene, making even the
intraglacial conditions more severe than along the southern

periphery of the Cordilleran and Laurentide Ice-sheets of

the Pleistocene.
Atlin I ~ Gladys I Glaciation

This stage of glaciation followed the Boulder Creek
Intraglacial and represents the first glaciation in the
region for which there is in-valley morphological evidence.

Atlin I Ice was thicker and assumed to be more thermOphysically

141
polar in character. This is inferred from the steeply dip-
ping gradient Of its terminal moraine remnants. The
rationale is that ice flows more slowly at temperatures
below the pressure-melting point, as demonstrated in the
laboratory by Glen (1955) by Nye (1953) and by Miller
(1956, 1975a) with field measurements on the Juneau Ice-
field. This results in steep terminal snouts. (In fact in
Greenland and the Antarctic such termini today are found to
be greatly oversteepened compared to those Of temperate
glaciers). This suggests that cold high-level ice came into
the region from the Juneau Icefield sector to the south as
well as from local high-elevation accumulation. Extensive
ice from the Cassiar source to the southwest and as well
much accumulation in local highlands in that sector indi-
cates that in this part of the Pleistocene the Arctic Front
had moved well east of its mean present position. This would
have more readily permitted maritime air to make its way into
the interior highlands. During this major glaciation ice is
thought to have completely filled the valleys between the
Coast Range and the Cassiar Mountains, resulting in a very
extensive accumulation on the Teslin and Taku Plateaux
(Bostock, 1948, and also Fig. 1 of this present report).

The Atlin I - Gladys I Stage is correlated with the
upper Woodfordian of the mid-continent and Salmon Springs
Stage II in the Puget Lowland. The Boutellier nonglacial
in southwestern Yukon Territory is C-14 dated by Denton and

Stuiver (1967) at between 30 and 40 thousand years B. P. As

142
such it agrees well with the Boulder Creek/McKee Creek C—14
dates and hence appears to have been contemporaneous with
the early part of Atlin I — Gladys I glaciation. Only with
more complete and absolute dates from the Atlin region, can
the apparent out of phase relationship found in the study
area be fully understood. Also because of some uncertainty
in the radiocarbon dates it is even possible that the begin-
ning of Atlin I was earlier than represented on the correla-
tion chart (Table V).

The more highly weathered moraines for Atlin I ice
than those of Gladys I age suggest that Atlin Ice built up
as the Arctic Front moved inland and was influenced by in-
creased maritimity, and retreated while Gladys I ice con-
tinued to build up primarily from out-flow from the less
maritime sector of the broad region in the Cassiar Mountains.
It is suggested that the Gladys I build-up correlates morpho-
genetically with the McConnell Glaciation in the southeastern
Yukon Territory (Hughes et al., 1969). Both of these gla-
ciations had an easterly source area and the Gladys I ice
lobe certainly extended well north into the Yukon Territory
beyond the present Yukon—British Columbia border (Figs. 1
and 2).

There is no clear evidence that Gladys I ice completely
withdrew from the Gladys Lake depression between Gladys I and
Gladys II time. Another possibility is to correlate the
Gladys I stage with Atlin II (v. Table V) which subsequent

research in the area south and east of Gladys Lake may sup-

port.

143

Evidence from the.luneau Icefield suggests that the
Atlin I-Gladys I stage is a part of the lower Gastineau/
Sloko glaciation (Miller, 1956) which was also morpho-
genetically a Greater Mountain Ice-sheet. Evidence for
this stage is present only in the sequence of erosional
landforms of cirques and berms, particularly the initiation
of low—elevation cirques (C-2, Miller, 1956, 1961, 1975a)
on the Alaskan coast at a mean elevation of 1100 feet (300
meters).

Also in the Juneau Icefield sector, no depositional
evidence exists for the Farmdalian Interglacial as the
extensive late Wisconsinan glaciation was so intense that
it removed all depositional evidence of any preceding de-
glaciation phase in the presently glacierized areas. From
the character of Atlin I moraines it is further suggested
that the originating ice in this glacial phase included a
build up of thermOphysically Polar ice which late in the
stage, in consequence of climatic amelioration, probably
changed its character to a more sub-Polar ice mass. Subse-
quent to this, it may be presumed to have been followed by
build up of a more Temperate ice character in the Boundary
Range, leaving no remnants whatsoever of the former colder
phase. Such warmer conditions naturally resulted in re-
treat of ice from the Fourth of July Creek Valley, eventually
to be followed by re-invasion of ice from the main Atlin

'Valley during the Atlin II stage.

144

A Suggested Pine Creek Intraglacial

As discussed above, the more extensive weathering of
Atlin I drift indicates a relatively important intra-
glacial in the area of Atlin drift which would make it the
equivalent of the Farmdalian intraglacial in the mid-continent
chronolOgy, and of the Olympia Interglacial of the Puget Low-
lands. For the Atlin and Taku areas this has been termed the
Pine Creek Interglacial, in reCOgnition of a substantial
weathering profile on drift sheets exposed in the valleys
of Pine Creek, Spruce Creek, Boulder Creek and McKee Creek
in the Atlin valley. No time correlate of the Pine Creek
Intraglacial has been designated in either the studies of
Denton and Stuiver (1967) or Hughes et al., (1969). From
this lack of evidence for an intraglacial found to date in
adjoining regions of the Yukon, as well as on the.Juneau
Icefield* and in the Gladys Lake sectors, it is assumed that
this phase, if indeed the correlation proves to be valid,
was not as long nor as warm as the Farmdalian Intraglacial

in the mid-continent region.

Atlin II-Gladys II Glaciation

 

The advanced position of Atlin II ice and the indicated
lower level of Gladys II ice suggests that the major accumula-

tion area in this phase was the Boundary Range especially its

* Except to cite the sub-stages (moraine entities) in
the Sloko Stage noted for the Taku District in Table V.

145

higher elevation nevés as zones of maximum snow accumula-
tion. Although still an important accumulation area, the
Cassiar Mountain region and its bordering Yukon and Stikine
Plateaus, did not foster as extensive a glaciation. This is
based on the reCOgnition that this stage was characterized
by a major still-stand at an elevation lower than that
reached by Gladys I ice. Continued field work east of the
study area could produce evidence for a more extensive re-
treat and re—advance of Gladys ice at this time. The
corresponding maximum advance of Atlin II ice within the
Fourth of July Creek Valley is also not clear, because of
the blocking effect of Gladys II ice plus the locally de-
rived ice from adjacent cirque basins. Thus all of this
interpretation has to stem from the observed moraine sequence.

In the foregoing context, there are also moraines on the
southeast flank of Mount Ewing which appear to be related to
Atlin Ice. Since the rest of the moraines in the upper
Fourth of July Creek Valley area are related to the Gladys
and/or local plateau ice, it is suggested that some of this
was ice nourished from the large 4000 foot (1300 meter)
cirque between Caribou Ridge and Mount Ewing. This cirque
basin is presently filled with a tarn lake, as shown in
Figure 3.

The Gladys II stage is correlated with the upper
Glatineau/SlOko stage of the Greater Mountain Ice-sheet in
the Taku District and hence to the lower to middle Wood-

fordian of the mid-continent stratigraphy. In this stage,

146
the Puget Sound area was not glaciated as early or as exten-
sively and both the Evans Stade and the Vashon Stade,
separated by an interstadial, (Table V) are correlated
within the Atlin II - Gladys II time period. The Puget
Lowland, of course, was much farther south and as well the
glacier termini there were a much greater distance from their
source areas. Are we to conclude that the Boundary Range and
the Cassiar highlands accumulation areas were not as sensi—
tive to climatic perturbations which produced the distinct
stages within the Woodfordian in the mid-continent area and
in the Puget Lowlands? This question is yet to be revealed.

Work done in the southern Yukon Territory delineates
no substage from maximum Atlin I to the end Of the Wiscon-
sinan. The Naptowne Glaciation of Cook Inlet (Karlstrom,
1964) may however, have a correlative stage in the Skilak
Lake stage noted in Table V.

Again, it may be pointed out that variation of climatic
conditions is not so clearly manifested in the drift records
described in this dissertation because evidence for earlier
fluctuations were so often erradicated by the next maximum
advance and also modified by the effects of later alpine ice.
It is also conceivable that climatic variation in the accumu—
lation zones only caused minor changes in accumulation levels
and not in the total amount of ice channeled into the Fourth

of July Creek Valley.

147

A Late — Wisconsinan Climatic Amelioration

During and following the demise of the Atlin II
glaciation, there was an extensive develOpment of glacio-
fluvial terraces produced by melt waters flowing in massive
quantities from the downwasting Gladys II ice—sheet in the
upper Fourth of July Creek Valley. This is illustrated by
the highest fluvial terrace shown in Figure 11 for the
intermediate valley section of the Fourth of.July Creek
Valley. The position of this related "intraglacial" is
also suggested schematically in Table V. Following this
time and the related sequence of glacio-fluvial events
there was a return to cooler and wetter conditions, which
in turn led to the major resurgence of ice in Atlin II and

Gladys II time.
Atlin III, IV and V - Glagys III and IV Glaciations

Because of the youthful tOpography and minimal weather-
ing found on the glaciO-fluvial terraces the Atlin III, IV
and V stages and the Gladys III and IV stages are considered
to be late-Wisconsinan in age. The evolution of this sequence
is indicated as follows.

A rather extensive retreat of Atlin II ice preceded
the Atlin III phase during which time the massive kame
moraine (Ruffner moraine — outwash complex) was built at the
juncture of the Porter Lake Valley and the Fourth of July
Creek Valley (Fig. 17). Following this, Gladys III ice

downwasted and retreated from the valley to the northeast,

148

the melt-waters scouring several large runoff channels as
the ice retreated. Three well-defined glacier levels lie
about 500 feet (152 meters) apart but the steep gradient of
their projected surfaces, as well as their quite similar
morphology and weathering, indicate that the represented
substages of Gladys III were very close in time. They have
thus been grouped as one stage.

The Ruffner moraine - outwash complex at the valley
junction is correlated with the zone of massive embankment
moraines near Jakes Corner on the Alaskan Highway. These
lie some 60 miles (100 kilometers) north of the village of
Atlin. Extrapolating from Carbon - 14 dates in bog sedi-
ments, Anderson (1970) has suggested that the Jakes Corner
moraines are of maximum late Wisconsinan age and hence were
built by 10,200 years B. P. Thus a similar age is assigned
to the Ruffner Moraine Complex. Dates from the upper Fourth
of July Creek plateau indicate that alder was growing 9800
C-14 years B. P. It is to be noted that both of these dates
are minimum. With respect to the mid-continent stratigraphy
and that of the Colorado Front Range, this stage (termed At-
lin III) is correlated with the late Woodfordian or the Port
Huron stage in Michigan and the Middle Pinedale Stage in the
Rockies. The McConnell Glaciation and Kluane Glaciation of
the eastern and south western Yukon Territory reSpectively
were present the entire period from Atlin/Gladys II to final
deglaciation at the end of Atlin V time. Data from Cook In-

let (Karlstrom, 1964) suggest that the Skilak Lake glaciation

149
ended about the same time as the final stage of the Wiscon-
sinan and was followed by an early Holocene glaciation, the
Tanya (Table v).

In the Boundary Range, the Intermediate Mountain Ice-
sheet glaciation, called the Douglas/Inklin stage by Miller
(Table V) is thought to be equivalent to the Atlin III -
Gladys III glacial resurgence. Evidences for this stage in
the Juneau Icefield region are the modification of cirque
levels 2 and 3 at the 1100 and 1800 foot (300 and 545 meter)
elevations, the presence of a pronounced erosional berm
(level 3) in the highland and the existence of notable
moraine embankments and terraces in the Inklin junction
region of the upper Taku River Valley (Miller, 1963).

The final downwasting and retreat of Atlin III ice is
documented by a recessional moraine complex near Lower Mc-
Donald Lake. This is considered to correlate with a huge
kame terrace complex at the lower end (north) of Atlin Lake,
(Fig. 2) 30 miles (48 kilometers) north of the Fourth of July
Creek mouth. By this time, the Atlin Ice was confined to
the main Atlin valley, with a flattish gradient of only 16
feet per mile (3 meters per kilometer). From the study of bOg
sedimentation and palynolOgy, Anderson (1970) has dated the
Atlin Lake kame terrace at Mile 24 on the Atlin Road, some
30 miles (48 kilometers) north of the Fourth of July Creek
Valley, at 9000 to 10,000 years B. P., which would correlate
to the Sumas Stage in the Puget Sound region of Washington

State and the waning Valders Stage Of the mid-continent

150

stratigraphy in the Great Lakes region of the mid-continent
Laurentide Glaciation. The final Wisconsinan glaciation

in the coastal and inland sector of the Taku District
(Juneau Icefield) is designated the Salmon Creek/Zohini
stage by Miller, again based on deltaic terraces in the
Salmon Creek valley near Juneau and on moraine and kame
terrace sequences near Zohini Creek in the upper Taki River
tributary sector of the southeastern Atlin region (Tables

V and VI). Thus Atlin IV - Gladys IV would be a correlate
of the lower Salmon Creek/Zohini stage, all of this being
associated with a lesser Mountain Ice-sheet in the highlands
of the Coast Mountains. The primary evidence for this stage
on the Juneau Icefield itself is the occupation and erosion
of Cirque level 4 at 2500 feet (757 meters) and the erosion
of berm level 4 in the wide-strathed glacial-filled valleys,
such as occupied by the present—day Taku and Llewellyn
Glaciers. Evidences at lower elevations along the Alaska
coast and in the Taku Valley and the southeastern Atlin
region lie in the presence of gravels stratigraphically
above marine till and diamictons in the Coastal area, and
corresponding pebble terraces at the mouths of interior
valleys in the upper Taku River valley. Upper Salmon Creek/
Zohini stage deposits are represented by undifferentiated
outwash gravels and kame terraces at Tulsequah and Zohini
Creek where the Taku River crosses the Alaska-British
Columbia border (Fig. l and Miller, 1956).

More precise teleconnectional study of the latest

151
Wisconsinan features in the adjoining Taku River region
could be profitable as the sequences do appear similar to
that described in the Atlin IV and V stages. Some of the
Taku sequences may, however, predate the early Holocene
(Miller, personal communication).

Once confined to their immediate valleys, near the
study area of this report, Gladys and Atlin Ice downwasted
and retreated rapidly. This has been shown by Anderson,
(1970) and also by the general lack of recessional moraines
in the immediate valleys. Atlin Ice retreated to near the
present glacial position of the Llewellyn and Willison
Glaciers, just south of the head of Atlin Lake, (Figs. 1
and 3). Following this, all Gladys Ice, which was by then
completely devoid of a source area, disappeared by down-
wasting and stagnation. This dessication left large kettle
ponds and pit lakes, as well as massive deposits of melt-
water sedimentation (largely sands and gravels) in the area

between Gladys and Teslin Lakes (Fig. 3).

CHAPTER IX

THE HOLOCENE

The Last 10 to 11,000 Years

Holocene climatic trends as discussed by Anderson
(1970), Miller and Anderson (1974) and Miller (1975a) show
a rather continuous warming trend from about 9500 years
B. P. to about 6000 B. P. This culminated in the Thermal
Maximum (Table II) when mean annual temperatures were some
20F (30C) higher than present. During this interval, the
most intensive of periglacial processes migrated to higher
elevations and the greatly retracted Juneau Icefield re-
ceived its maximum accumulation only on the highest neve
and at cirque levels 5 and 6 (i.e., at 3200 and 3900 feet;
975 and 1190 meters as shown in Table III). The latter is
some 700 feet (210 meters) higher than the lowest cirque
filled with ice today. Evidence for the Thermal Maximum
in the Atlin region is also given by a recent C-14 date of
3472 i 87 years B. P. on a silty peat horizon in the fore-

land area about a mile below the Cathedral Glacier terminus

(Miller 1975a; Jones, 1975; also see Radiocarbon, 1976, Vol.

18, NO. 1, sample AU-77A). This site lies between two re-

cessional moraines of early Holocene age.

152

153

Near the end of the Thermal Maximum, the Arctic front
maintained an average position well west of the inner
channel-ways of the Alexander Archipelago and the south-
eastern Alaska coast. Then tOO, wetter conditions and
increased storminess prevailed on the continental flank of
the Boundary Range and well over into the Atlin region.

The Thermal Maximum (Hypsithermal Interval or Climatic
Optimum of some writers) was followed by a major cooling
trend as the Arctic Front once again moved inland. This
was coincident with a decreased storminess and drier condi-
tions accompanied by general cooling in the Atlin district.
Concurrently, on the coast, relatively cooler and wetter
conditions dominated the Juneau Icefield and the Taku
District. The nature of these out-of-phase climatic trends
has been documented and explained by Miller and Anderson in
their 1974 report on glacial and palynolOgical changes in
this region during the approximately 10,000 years of the
Holocene.*

From these studies it is also apparent that the early
NeOglacial time interval from about 2500 to 1200 years B. P.
(600 B. C. to 700 A. D.) was as much as 4°F (2°C) cooler
than today. This was a period of extremely intense peri-
glacial activity in the upper Fourth of July Creek Valley as
described from the field evidence discussed in Chapter VII.

Then a short warm interval persisted from about the end of

* Defining the Holocene as the interval since the last
major glacial or large-magnitude climatic event. On this
definition the Holocene may represent quite different time
spans in different regions of the globe.

154
the 7th century (1200 years B. P.) to the 13th century
(750 B. P.) (Miller and Egan, 1968). This is now well-
documented for the Alaskan coast by significant C—14 dates
on overridden trees at the terminus of Mendenhall Glacier
(1,197 i 153 years B. P.) and of the Taku Glacier (970 i
100 years B. P.) at sites studied by Miller (1975a); also
see Radiocarbon 1976, sample AU-100).

Since about the 14th century, the Atlin region has
experienced a relatively and wetter climate compared to the
short warm interval of the Middle Ages noted above. With
these cooler conditions, in the Atlin region nivation hollows
have been periodically filled with firn as the lesser

climatic oscillations continued.

Quaternary Peat Deposits in Boulder Creek Valley

Carbon 14 dates obtained in the Boulder Creek Valley
further substantiate the late Holocene climatic changes. In
Figure 37 is illustrated the position of two late Holocene
organic horizons stratigraphically above an organic horizon
representing the beginning of the Holocene and which lies on
the latest Wisconsinan drift surface comprised of unweathered
ground moraine and basal till. The organics are well-formed
intercallated with lacustrine and glacio-fluvial layers.

The lower of these late-Holocene peat horizons is dated
2770 C-14 years B. P. (Geochron, 1973) which correlates
fairly well with the Cathedral Glacier peat bog date of
3472 i 87 C-14 years B. P. It represents the end of the

Thermal Maximum and beginning of Anderson's (1970) cool dry

155

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156
Neoglacial period. As such it is related in time to Neoé
glacial advances of the Mendenhall Glacier near Juneau dated
by several selected samples which pre-date the NeOglacial
advance at 1800 to 2200 C-14 years B. P. (Heusser, 1952;
Miller, 1956; Cross, 1968; Miller, 1975a). The upper
horizon at 1100 C-14 years B. P. (Geochron, 1973) correlates
well with the later Mendenhall buried forest (1197 i 153 years
B. P.) near Juneau, cited in the previous section. It also
may possibly correlate with an overridden forest of mid-
NeOglacial age near the Davidson Glacier on the western side
of Lynn Canal some 10 miles southeast of Skagway, Alaska
(Egan, 1971; and Fig. 1).

Because of these teleconnections it is most probable
that the two upper peat horizons in Boulder Creek Valley are
associated with regional climatic change, though one could
also suspect some relationship to changes in drainage of
Boulder Creek and associated bOgs along the shores of Sur-
prise Lake (Fig. 1). Today the vegetation above the upper
peat horizons is well drained and dry. Therefore it is
certain that at least locally wetter conditions existed during
the formation of the peat and that this was followed by
fluvial deposition.

Increased nivation and perhaps occupation of high-level
cirques by renewed glaciation could have produced the run-
off necessary for fluvial deposition after each peat-
producing period. Further stratigraphic work in this and
other valleys in the area should lead to a refined chronology

as it applies to the Holocene.

157
Based on the evidence cited in Chapters VI and VIII,

the following tabulation (Table VI) of the Glacio-
ChronolOgy reviews the Wisconsinan events in the Fourth
of.July Creek Valley. By implication it also refers to

the broader Atlin Region and the adjacent Taku District to
the south and west. In combination this summary represents
the late—Pleistocene chronology of the northern end of the

Wisconsinan Cordilleran Ice Sheet.

Approx. Yrs.

B.P.

O

OGO‘AN

12

14

16
18

20
30

4o
50
100

158

TABLE VI

Atlin Region
(this study)

Nivation and palsa

develOpment

LATE-PLEISTOCENE GLACIO-CHRONOLOGIC SUMMARY
FOR THE ATLIN REGION AND TAKU DISTRICT

Takupistrict
(Miller, 1956)

Neoglacial

 

Thermal Maximum

Intense peri-
glacial

High cirque
glaciation

 

Atlin IV & V

Gladys IV glaciation

Salmon Creek/
Zohini stage, with
3 substages

 

 

 

 

 

 

Atlin III-Gladys III Douglas/Inklin
glaciation stage

Climatic Amelioration Intraglacial
Atlin II-Gladys II Gastineau/Sloko
glaciation stage

Pine Creek Intraglacial

AtIin'I-Gladys I glaciation

Boulder Creek Intraglacial Intraglacial

Pre-Atlin I glaciation

Early’Juneau/
Atlin Stage

CHAPTER X
SOME SUMMARIZING CONCLUSIONS

The correlations attempted in the study area were
done with the axiom that if large-magnitude continent-wide
and even world-wide climatic perturbations were responsible
for major glacial advances and retreats in the mid-continent
region of the United States and Canada then major atmospheric
perturbations should also have been recorded in the sequence

of glacial landforms in the Pacific Cordilleran region.
Problems in Teleconnection

The reader was reminded early in this report that the
main climatic and/or glacial divisions might be expected to
be out-of-phase in the two localities considered, i.e., the
Atlin Valley and the Gladys Valley provenance zones. This
should be particularly apparent in the early phases of the
glacial sequence when it is expected that major accumulation
would begin somewhat earlier in the highlands of the Boundary
Range. Thus the chronO-relationship (Table V) suggests that
the earliest Wisconsinan glaciation actually began in late
Sangamonian time. That is to say major accumulations were
initiated in the regions of highest elevation and especially
near the coast where ready sources of moisture prevailed.

159

160
Both of these prerequisites are fulfilled in the Alaska-
Canada Cordillera. Although there are no time-stratigraphic
'boundaries known, the Boulder Creek (Port Talbot or "equiva-
lent") intraglacial is inferred to have been shorter in the
Cordilleran region than revealed in the midécontinent
chronology. The reason for this conclusion is again stated
to be an earlier build-up of ice during re-glaciation and a
retarded recession at the end of the early Wisconsinan (pre-
classical) glacial stage...i.e., just prior to Gladys I
glaciation.

Even in terms of sequence, there is limited evidence
for a correlate of the Farmdalian intraglacial in the
Cordilleran region, except for some notable weathering of
low-elevation tills in the Atlin Valley. This is found on
a basal till widely exposed on the sluiced down-banks of
McKee Creek, Pine and Spruce Creeks and in the Boulder Creek
and Otter Creek areas where intensive gold mining during and
since the Gold Rush days has exhumed this key horizon. In
each case, this till lies stratigraphically below a younger
and unweathered surface till. Because of particularly note-
worthy limonitic weathering of the lower unit in the Pine
Creek area (near the old mining town of DiscoverY) the name
Pine Creek Intraglacial is suggested in Tables V and VI, to
represent a retreatal phase between Atlin I and Atlin II
glaciation.

During this time of ameliorating climate in the mid-

continent region melting of the Laurentide glacial sheet was

161

greatly enhanced by increased precipitation in the summer
months as maritime sub-trOpical air masses moved up from the
Gulf of Mexico as they do today. The frontal position of
these storms was presumably very much controlled by position
of the ice margin and by the cold high pressure air masses
which were associated with the ice sheet beyond this margin.
With these increased temperatures and precipitation, the

ice receded and Farmdalian soils develOped (Ruhe, 1975).

In the region of the Cordilleran ice sheet, however,
there were strong geographic and orographic influences On
the progress of deglaciation. For example, in spite of
global atmOSpheric warming, the result in the highlands of
the northern Cordellera was increased high-level accumula-
tion. This continued to nourish glaciers in the Alaskan
Panhandle and in the interior regions of British Columbia
and the Yukon. Because of the high elevation and relief in
this region, the global climatic warming was insufficient to
cause much reduction in the upland ice masses in this
Cordilleran region during the Farmdalian time—span. Only
the low-elevation valley glaciers suffered wasting, such as
those filling the valleys of McKee Creek, Pine Creek, Spruce
Creek, Otter Creek and Boulder Creek, as previously dis-
cussed. The relative locations of these key sites are
indicated On the map of Figure 1 and shown in tOpOgraphic
detail in Figure 3. In fact, it is implicit here that there
was probably not a true intraglacial manifested in the

interlobate sector of the highland study area where there

162

was such an interplay between the Atlin and Gladys ice and
where even in the waning and waxing phases local cirque
glaciers added substantially to the ice-cover.

In the final stages of Wisconsinan glaciation (Atlin
III, IV and V time) quite different conditions appear to
have existed. Certainly there is strong evidence that the
ice retreat was rapid. From the time-stratigraphic record
given by C-14 dates, it appears that these stages were con-
current with those of the mid-continent deglaciation. Miller
(1975a) has also suggested this for the adjacent Taku Dis-
trict and the Southeastern Atlin region, noting that "the
main morpho—erosional and morphO-stratigraphic sequences,
abetted by available time-stratigraphic evidence, are sub—
divided into an Early1Juneau/Atlin (early Wisconsinan) stage
and a Gastineau/Sloko (mid-Wisconsinan) stage, followed by
two late Wisconsinan stages...the Douglas/Inklin and Salmon
Creek/Zohini stages. On the basis of sequence, the latter
two should be equivalents of the Port Huron and Valders
glaciations of the mid-continent chronolOgy." This observa-
tion distinguishes theglacio-climatic character at the end
of the Wisconsinan from the apparently quite different nature
and teleconnectional relationship of earlier stages within
the Wisconsinan.

As has been previously shown, in the northern Cordil—
leran region the nature of the pre-Atlin I phase of glacia-
tion (early-Wisconsinan) is still obscure and hardly under—

stood and, indeed, in this region it may never be. But there

163

is abundant evidence for the Atlin I and Atlin II glacia-
tions, and for the final wasting of stages in Atlin III,
IV and V time.

As for the Pine Creek Intraglacial (ca. 13,000 to
14,000 B. P.), it is reiterated that this is only charac-
terized in the lower valleys of the Atlin region. The high
plateau sector of the upper Fourth of July Creek valley con-
versely experienced a build-up, or at the least was affected
by a long period of still-stand of ice. Hence on evidence
from the study area, a classical type of intraglacial cannot
be distinguished. In fact, just the Opposite apparently took
place, with effective increases in local glaciation in the
high cirque-headed valleys. This situation and its basic
significance in terms of the causal factors involved when
there are simultanious advances and retreats of glacier
lobes have been discussed at length in the early.Juneau Ice-
field Research Program reports. Some of these illustrate
particularly well the nature of so-called anomalous advances
on the Taku Glacier in contra-distinction to the retreatal
activities of the termini of nearby Norris Glacier, the Twin
Glaciers, and the Herbert and Mendenhall Glaciers on the

Juneau Icefield (Miller, 1956, 1963, 1973a, 1973b).

Conceptual Model fog Out-of-Phase Pleistocene
Glacial Variations of Large Magpitude
In view of the classic mid-continent Wisconsinan
chronology to which reference has been made in the correla-

tion chart of Table V, it is of interest that large-magnitude

164
out-of-phase variations with time-stratigraphic significance
have occurred in the regime of the Des Moines, Michigan,
Huron and Erie glacial lobes during the Pleistocene (Flint,
1971; Embleton and King, 1974). This has long perplexed
glacial geologists, especially those without experience in
the behavior of existing glaciers.

A significant contribution to the understanding of the
out-of-phase phenomenon has been made by the research teams
of the Juneau Icefield Research PrOgram which, in the mid
1940's, began to lock at the total system activity of
glaciers with multiple neves and separate nourishment pro-
venances. A detailed analysis of this problem, exemplifying
the behavior of the Norris (receding) and Taku (advancing)
glaciers, was first presented by Miller (1956, 1963). In
these, the effects of shifting zones of maximum snowfall and
changing freezing levels through time have been delineated.
Since its beginning the study has been based on solid
empirical evidence through a number of years of systematic
sampling of net accumulation changes in different sectors
and at different elevations on the Juneau Icefield. Because
the icefield is geographically adjacent to the study area
dealt with in this report these findings are of special
interest.

Subsequently, Anderson (1970), Miller and Anderson
(1975, 1974a and 1974b), Jones (1975) and Miller (1973,
1975a) and Miller (1975c) have elaborated on the conceptual

model based on recent and current regime trends on the.Juneau

165

Icefield. This has been done by extending these ideas
backward through the Holocene to explain the seemingly
enigmatic out-of-phase climatic characteristics in the
Atlin area compared to the Taku District on the Alaskan
coast. Details of these out-of-phase climatic differences
have been noted in Table II.

The results of the writer's current studies in the
Fourth of July Creek Valley tend to substantiate the
validity of this conceptual model in terms of its appli-
cation to those realms of the Wisconsinan time-scale which
predate the Holocene. The conceptual model has also helped
to explain the nature of the out-of-phase oscillations which
the glacio-morphostratigraphy and the glacio-fluvial record
gives in the Porter Lake Valley and in the intermediate
segment of the Fourth of July Creek Valley. This includes
that most significant interlobate zone above and below the
Ruffner Moraine Complex. Thus the fundamental interpreta-
tions of this study have leaned on key results of the cited
previous investigations of Holocene stratigraphy in the
Atlin valley and on the NeOglacial and Little Ice Age re-
search on the.Juneau Icefield prOper. But also the current
study has built upon these foundation concepts, and it is
hOped that they had added some verification to them.

It is also sincerely hOped that this report will serve
as a ledge upon which other researchers may stand as further

detailed investigations unfold.

 

CHAPTER XI

FURTHER RESEARCH PROSPECTS AND
ADDITIONAL STUDIES
Although this study was initially focused on the
relatively small arena of the Fourth of July Creek Valley,
it became apparent early in the investigation that it was
essential to look beyond the immediate field area in order
to come to grips with some of the basic problems involved.
This is why, a few Observations from surrounding areas and
from the larger adjoining region have been included. Be-
cause the considerations involved this larger area which
much of which is characterized by inaccessibility, high
relief, limited exposures and, in some cases even ground
cover of dense undergrowth, it has been difficult to encom-
pass. The constraints are not nearly as severe as those
occasioned by the extremely rugged terrain and dense
mantling of rain forest which obscures so much of the glacial
geology in the Alaskan coastal sectors of the Taku District.
Still, because of this situation and in terms of the chrono—
logical interpretation, the results of this study must be
looked upon as fundamentally reconnaissance in nature. But
some satisfaction is derived from the fact that an important
step has been made, for in any new region reconnaissance

geology is essential before detailed work can follow. This

166

 

167
difficulty has been clearly reCOgnized in other studies
referenced from the Alaska—British Columbia-Yukon region -
e.g. those of.Johnson, Denny, Hughes, Karlstrom, Denton
and Stuiver, Rutter, Miller et a1.

Although some details of the chronolOgy may not
stand severe scrutiny of future investigations, the main
tenets of the chronolOgy should remain valid. Some con-
fidence stems from the gross similarity of the conclusions
with the main lines of interpretation found in adjoining
regions - albeit some of the interpretations from other
regions are quite incomplete. The hope is that interest
will now be generated to encourage further work in other
critical valleys of the Cassiar and Coast Mountains andin
the Atlin region. Such investigations should include a de-
tailed analysis of soils, and of soils-vegetation relation-
ships along the line of the reconnaissance work of Anderson
(1970) and Lietzke (1969) in this area. Also further time-
stratigraphic information must be diligently searched for
because, to date, too few buried peat horizons have been
located. When such additional research is accomplished,
hOpefully more complete teleconnectional correlations can
be made from valley to valley.

In the broader context it is the writer's belief that
this reconnaissance study has established a framework for
the Wisconsinan stratigraphy of the region from which future
workers can begin to attack the more specific details. The

most urgent needs for amplifying these results is additional

168
loontrol on dating of the glaciO-climatic events, and further
soil-stratigraphic evidence. Therefore, even though we are
a long way from where we were in this investigation several
jyears ago, there is still much to be done and many questions
'to be answered. As more absolute time relationships become
available, however, some of the key questions should be

answered and the reCOgnized gaps in the chronology closed.

APPENDIX

GLOS SARY

Diffluent ice: The condition of glaciation when ice
is confined to major valleys.

Freezing level: The elevation at which permanent
snow accumulates. This level rises with increased tempera-
tures and fall with temperature lowering.

Holocene: The time interval since the last major
glacial or large-magnitude climatic event. The Holocene
may represent quite different time spans in different
regions Of the globe.

Kame moraine: A moraine which is mantled by ice
contact glacio-fluvial drift.

Storm paths: Peripheral winds in the interaction
zone between the low-pressure maritime cells and high-
pressure continental cells along the north Pacific Coast.
This is distinct from the "cyclonic tracts" which refers
to the movement of the center of low-pressure systems which
is normal to the storm paths.

Transfluent ice: The condition of glaciation when
all available cols are used and ice spilling from one

valley to another, ice.

169

BI BLI OGRAPHY

 

BIBLIOGRAPHY

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Aitken, J. D. 1959. Atlin map-area, British Columbia.
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Anderson, J. H. 1970. A Geobotanical Study in the Atlin
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Anderson, J. H., Miller, M. M. and Tallman, A. M. 1975.
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Geological Survey of Canada Memoir 312: 156 p.

 

 

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( :11 SOUTH- EASTERN PUGET LOWLAND GREAT LAKES WISCONSIN
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m \\ Atlinll I Glad sl‘ 9 3 Early (3malnt.liunlts, ‘8
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Silver ~—’ 1 "_T Y _ Intra lacial ‘
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\ 2 upper raS anngarno tan)
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