Effects of Permafrost on Stream 

Channel Behavior in Arctic Alaska

GEOLOGICAL SURVEY PROFESSIONAL PAPER 1068
Effects of Permafrost on Stream Channel 

Behavior in Arctic Alaska

By KEVIN M. SCOTT

GEOLOGICAL SURVEY PROFESSIONAL PAPER 106

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON: 1978
UNITED STATES DEPARTMENT OF THE INTERIOR 

CECIL D. ANDRUS, Secretary

GEOLOGICAL SURVEY 

H. William Menard, Director

Library of Congress Catalog - card No. 78-600058

For sale by the Superintendent of Documents, U.S. Government Printing Office

Washington, D.C. 20402

Stock No. 024-001-03088-8
CONTENTS

Page

Abstract__________________________———-- 1

Introduction _______________ ————————————— —— 1

Acknowledgments __________________________ 2

Description of streams ___ ___ _—————— —————— —— 2 

Atigun River ____________—— ——— __—————— 6

Sagavanirktok River ______———— ——————— --—— 6 

Kuparuk River __________——— ——— —————— —— 6

Happy Creek _______ _———— —————— —— —— 7

Putuligayuk River ——— — —————— ——— ————— —— 7

Data collection.._________-____—__ ——————— —— 7 

Measurement techniques __ ___————_————— 7 

Rates of thaw in beds and banks —— —————— ————— 8 

Thaw and erosion rates compared ———— ————— __—— 9 

Thermo-erosional niching—————__ ——— ————— ——— 10 

Bank slumping and vegetation effects __________———— 10

Variations in timing of bank slumping ——————— —— 11

Page

Data collection—Continued

Sagavanirktok River, sites S4 and S5 ——————————— 12 

Bed scour _____—______ ————— ————— — _————— 13

Channel processes during remainder of 1976 runoff season 13 

Atigun River ——— —— _ ———————— ———— ——— 14 

Sagavanirktok River _____————— —— —— ——— 14

Sediment transport——————__ ——— —— _ ——— —— —— 14 

Data interpretation ____ —————— ————— —— ————— ——— 15

Effects of permafrost ——————————————————————— 15

The varying importance of breakup processes ___————— 16

Summary and discussion ————_ — __ ———— ———— — __ 16

Effects of major floods ——————————————————— 17

Long-term erosion rates __ ———————— _ ——— __—— 17

Hypothesized uniformity of erosion rates____ ————— __ 18

References cited ____ — ___ ———— --- —————— --—————— 18

ILLUSTRATIONS

FIGURE 1. Map showing locations of study sites ____________________----_-_-_-------------_------------------------------- 3

2. Seasonal hydrographs of streams in arctic Alaska ______-_-__-__-----_------------------------------------------ 5

3. Graphs comparing thaw depths in submersed bed and bank material of differing texture at three sites on the Atigun

River ___________________________-__________________------------------------------------------------------ °

4. Graphs comparing thaw depths in emersed bed and bank material of differing texture at three sites on the Atigun River..._ 8

5. Plot of erosion versus thaw for points in non-cohesive bed and bank material-____-___-___-__---_------------------ 9

6. Plot of erosion versus thaw for points in cohesive bed and bank material--.--------------------------------------- 9

7. Cross sections of undercut banks ______--_-_____________--------------_----------------------- — --------------- 10

8. Plot of percentage of bank failures at site SI, water temperature at site SI, maximum daily air temperature at Happy

Valley Camp, and mean daily discharge at gaging station on the Sagavanirktok River ________________________ 11

9. Plot of percentage of sample sinuous or meandering reaches showing bank failure versus percentage of time interval

between breakup and peak of breakup flooding-__.._____________________------------------------------------ 12

10. Cross sections of bed surfaces and thaw surfaces in the thalweg at Atigun River site A3 __________________________ 13

11. Plot of percentage of sample reaches showing bank failure at peak of breakup flooding, versus drainage area-------- 16

TABLES

TABLE 1. The study sites and their characteristics----------------------

2. Peak discharges measured at crest-stage gage on Happy Creek

HI

EFFECTS OF PERMAFROST ON STREAM CHANNEL BEHAVIOR

IN ARCTIC ALASKA

By KEVIN M. SCOTT

ABSTRACT

Sites with drainage areas ranging from 88 to 12,200 km2 were 

monitored on five streams in northern Alaska during the breakup in 

1976 to determine (1) the effects of frozen bed and bank material on 

channel behavior, and (2) the importance of the annual breakup flood 

in forming the channels of arctic streams.

The thawing and concomitant erosion of channels varied with 

changes in bed-material size, channel pattern, drainage area, and 

climate. The response of channels to breakup flooding ranged from 

total permafrost control of channel processes, including both bed 

scour and lateral erosion, to only brief restriction of channel behavior 

early in the rise of the flooding. The watershed characteristic that 

appears to explain much of this variation is size of drainage area.

Similar variation of channel response with these factors is believed 

to be responsible for diametrically opposite results reported from 

recent studies of the two problems posed above. Permafrost has been 

cited both as the cause of extreme stability and as the cause of unusual instability in arctic streams as compared to those elsewhere. 

That permafrost simplifies the study of arctic stream channels 

through its domination of the effects of other variables has been a 

common assumption. As a consequence, however, the generalizations 

based on a single stream or on similar streams have led to a spectrum 

of inconsistent results. This spectrum of previous results now appears potentially consistent.

Comparisons of absolute rates of lateral erosion are not feasible, 

but it is likely that the net effect of the permafrost environment is to 

create greater channel stability than is found in unregulated 

streams of similar size in nonpermafrost environments. Combinations of factors, particularly those that encourage high rates of thermo-erosional niching, can nevertheless cause rates of erosion that 

dictate caution in engineering design.

INTRODUCTION

Many studies of arctic streams have been made since 

1970 in response to an increased concern about development, particularly of oil and gas resources, in regions of continuous permafrost. This expansion of 

interest has occurred simultaneously in Alaska, 

Canada, and the U.S.S.R. It has focused especially on 

the relationships of permafrost to the hydrology of 

streams and to the behavior of stream channels. Both 

these aspects are fruitful areas of geomorphic research 

with applications to the engineering of structures in 

flood plains and stream channels. It is the latter aspect, 

however—how and when channels in the permafrost 

environment respond to flow—to which this investigation is confined.

Two questions have been of recurring interest in 

studies of the behavior of arctic streams: (1) What is 

the effect of permafrost on rates of bank erosion—does 

permafrost inhibit lateral migration, or do the streams 

of permafrost areas have higher lateral erosion rates 

than might be expected for channels that contain effective flow less than half the year? And, (2) what is the 

role of the spring breakup flooding, compared to flow 

during the rest of the runoff period, in determining the 

channel characteristics; that is, what is the annual 

timing of the channel-forming processes?

In addressing the first question several recent workers have agreed with Leffingwell's (1919, p. 172) conclusion that lateral erosion is generally retarded by 

frozen ground. Cooper and Hollingshead (1973, p. 276) 

note that permafrost stabilizes bank material that, if 

unfrozen, would offer little resistance to erosion. 

McDonald and Lewis (1973, p. 2) believe that permafrost retards erosion over short time spans. It also has 

been suggested that, because the annual floods of some 

Alaskan streams may occur while channels are frozen, 

the channel size is anomalously small in relation to 

flood discharges (see comment by C. R. Neill in Kane 

and Slaughter, 1972, p. 540).

Walker and Arnborg (1966, p. 171) and Ritchie and 

Walker (1974, p. 545), on the other hand, emphasize 

the role of permafrost in promoting notable lateral erosion (an average of 10 m in one area) by lateral undercutting during breakup of the Colville River in northern Alaska. Walker (1973, p. 73) notes a general belief 

that lateral erosion is "especially important" in arctic 

Alaska because of permafrost. High rates of erosion of 

bedrock permafrost by streams have been attributed to 

an "ice rind," an easily eroded surficial zone that has 

been shattered by ground ice (Biidel, 1972, p. 9).

The engineering reports dealing with pipeline construction also manifest differing opinions on the stability of arctic streams. One 1975 unpublished report cites 

the "extreme" stability of a reach of the Sagavanirktok 

River in northern Alaska as being related to the presence of permafrost in the flood plain. Another, compiled 

in 1974 and dealing with rivers on the Arctic Coastal
EFFECTS OF PERMAFROST ON STREAM CHANNEL BEHAVIOR IN ARCTIC ALASKA

Plain east of the Sagavanirktok River, notes that such 

rivers have a "greater tendency . . . for lateral migration . . . than is the case for most rivers in other 

physiographic regions." The significance of these comments is less in their substance than in the fact that 

they exist without contradiction in the scientific literature.

Which opinion is valid—are the streams of arctic 

Alaska extremely stable or do they in fact have 

greater-than-average rates of lateral migration? More 

definitive work is plainly needed. An example of the 

practical reasons for the recent concern over rates of 

lateral erosion is the placement of pipeline sag points, 

the sites of downward inflection of the pipe to depths 

safe from exposure by scour (short-term vertical erosion of bed) or longer term degradation of the 

streambed. Because deeper burial is costly and involves other problems, the length of deep burial is 

minimized in compliance with estimates of the probable rate of stream migration at each point of streampipeline proximity.

Evidence on the second question—the importance of 

breakup flooding and of the annual timing of the processes that significantly influence the beds and banks of 

streams—is just as diverse. Walker and Arnborg (1966, 

p. 168-171) found that most lateral erosion occurs during or shortly after breakup flooding on the Colville 

River, as did Outhet (1974, p. 308) for each of five 

categories of bank shape in the Mackenzie River delta. 

At the other extreme, Miles (1976, p. 196 and 199) 

found that breakup on Banks Island in the Canadian 

Arctic was relatively unimportant in this respect, and 

that significant change in the channels did not begin 

until summer storms occurred several weeks later. To 

complete the possible variations, Abramov (1957, p. 

112) reported that bank slumping along the lower Lena 

River in Siberia was greatest toward the end of summer.

These conclusions suggest that the behavior of arctic 

streams, in response to a relatively simple hydrologic 

regime of spring breakup flooding followed by low 

summer flow with occasional rainstorm runoff events, 

is far from uniform. Some studies of arctic streams 

have implied the extension of conclusions representing 

perhaps a single stream over one season to arctic processes in general. This study is not an exception to this 

tendency, but it approaches the problem through the 

study of multiple sites on a group of five streams representing mountain, foothill, and coastal-plain drainage 

areas. A degree of perspective is added to the observations by flow records of from 5 to 7 years duration on all 

but one of the five streams. The study periods were 

May 5 to June 14, 1976, an interval that included the 

rise and recession of breakup flooding in most of the

streams, and May 24 to June 1, 1977.

The study attempts to show the effect of permafrost 

on streams by concomitant measurement of lateral 

erosion and of the rate of retreat of the thawed zone in 

the eroding bank during the period of breakup. Care 

was taken to define two factors that are believed to be 

closely related to any observed variation in erosion 

rates—size of bed and bank material as a cause, and 

channel pattern as an effect. Although any of several 

factors can affect channel pattern, Brice (1964, p. 32) 

regards bank erodibility as the most significant single 

variable. Bank erodibility in turn depends mainly on 

the size distribution of the bed and bank material (see 

Schumm, 1960, p. 17 and 22-23), as well as on the 

vegetation of the banks (Mackin, 1956; Smith, 1976), 

and, in the case of cohesive material, on the moisture 

content of the bank in combination with freezing and 

thawing (Wolman, 1959, p. 214).

ACKNOWLEDGMENTS

The study was inspired by the late Don M. Culbertson, whose advice and council made it possible. 

Logistical support was arranged by the Alaska District, Water Resources Division, U.S. Geological Survey, and the Alyeska Pipeline Service Co. Hal Hume, 

Alyeska Camp Manager at Happy Valley, expedited 

the work in several helpful respects. Among the reviewers of the draft manuscript, those making especially helpful contributions included R. K. Fahnestock, 

State University of New York at Fredonia, and J. C. 

Brice, J. M. Childers, H. P. Guy, Reuben Kachadoorian, and J. R. Williams of the U.S. Geological Survey.

DESCRIPTION OF STREAMS

The five study watersheds drain the almost treeless 

desert or semidesert that constitutes much of the north 

slope of the Brooks Range (fig. 1). The area is one of 

continuous permafrost in which the active layer, or 

zone of seasonal thaw, reaches a late-summer thickness of less than 0.3 m in tundra-covered cohesive 

material and as much as 2 m in unvegetated noncohesive material. Modern glaciers, present only in high 

mountain valleys above 1500 m altitude, cover less 

than one percent of the drainage area of the Atigun 

River—on average the highest and most mountainous 

of the watersheds studied.

Mean annual precipitation, although not well 

documented, is probably less than 200 mm over much 

of the coastal plain and increases with altitude to 

larger totals near the crest of the Brooks Range. Approximately half of the precipitation falls as snow from 

September through May, and the resulting meltwater 

runoff is largely concentrated in the breakup flooding
DESCRIPTION OF STREAMS

of late May and early June. Summer rains are associated with the semipermanent arctic front that oscillates irregularly between the Brooks Range and the 

Arctic Ocean. The temperature gradient with latitude 

may be pronounced and, associated with the effects of 

coastal fog, can result in delay of breakup in streams 

draining the coastal areas by several weeks beyond 

that of streams with their headwaters in the 

mountains and foothills. The temperature gradient reverses with the season, however; winter temperatures

150 148

EXPLANATION

f~S2 

Study site on stream

0 10 20 30 KILOMETERS

I I I I

FIGURE 1.—Map showing locations of study sites.

are frequently between —26° and — 35°C at the coast 

and —30° and — 40°C inland, and summer temperatures are commonly between 5° and 13°C near the coast 

and 10° and 18°C inland.

Stream valleys within the Brooks Range were the 

sites of late Pleistocene mountain glaciers that extended as much as 50 km north of the range front 

(Hamilton and Porter, 1975, p. 480). Unglaciated 

uplands and flood plains were profoundly affected by 

periglacial processes such as solifluction and the formation of patterned ground—processes continuous 

today but less intensively active. The larger valleys in 

the Sagavanirktok River headwaters are glacially 

sculptured, floored with glacial-outwash sediment, and 

bordered by alluvial fans formed by debris flows from 

tributary hanging valleys.

The channel patterns vary in a downstream direction with changes in several factors, the most obvious 

of which is bank erodibility. The Sagavanirktok River, 

for example, changes from braided to sinous or meandering and back to braided during its course through 

the foothills. A similar sequence in pattern occurs on 

the Atigun River and can be ascribed to the same basic 

situation—a lower gradient in reaches upstream from 

points of control by bedrock or moraines has resulted in 

deposition of fine-grained alluvium. The cohesive nature of the fine-grained sediment results in reduced 

bank erodibility that is accompanied by the development of a sinuous or meandering pattern. A braided 

pattern occurs downstream as well as upstream from 

the sections with fine-grained alluvium. In most instances the stream pattern is clearly braided or 

sinuous-meandering, but in others the pattern appears 

to be transitional, similar in some respects to the pattern called wandering by Mollard (1973, p. 347 and fig. 

4). Channels may be generally braided at lower stages 

and clearly sinuous or meandering at intermediate and 

high stages. Channel patterns are described in table 1 

according to a scheme (J. C. Brice, written commun., 

1977) in which streams with a sinuous pattern have a 

sinuosity of 1.06 to 1.25, and those with a meandering 

pattern have a sinuosity between 1.25 and 2.0. Sinuosity is the ratio of length of channel to length of 

meander-belt axis, or to length of valley axis if no 

meander belt can be defined (Brice, 1964, p. 25).

The fine-grained alluvium forming the banks of 

sinuous and meandering channels generally consists of 

cohesive silt and clay with organic-rich layers and associated soil development in the upper 0.3 to 0.7 m. 

Gravel stringers are commonly intercalated in the 

basal part of the silt-clay alluvium, which may be as 

much as 4 m in thickness, and the unit overlies coarse 

channel deposits with a texture closely approximating 

that of the bed material of the stream. The unit may
EFFECTS OF PERMAFROST ON STREAM CHANNEL BEHAVIOR IN ARCTIC ALASKA

thin laterally, and locally may consist only of a predominantly organic tundra mat. Much of the finegrained alluvium apparently consists of reworked loess 

deposited during overbank flooding and stabilized by 

tundra vegetation.

The most obvious form of lateral erosion in the 

streams is thermo-erosional niching—the process of 

undercutting of frozen banks by thawing and erosion at 

the same time—described from the Colville River delta 

by Walker and Arnborg (1966, p. 166-167). Banks may 

be undercut, or "niched," to surprising distances by 

breakup meltwaters.

All observations indicated that the banks and beds 

at the study sites were frozen at the start of the runoff 

period. A possible exception was the downstream part 

of the Sagavanirktok River (site S5), where measurements could not be obtained and the year-round existence of a shallow aquifer in the stream alluvium has 

been indicated (Sherman, 1973). The degree to which 

the Sagavanirktok River channel is underlain by permafrost is controversial, and the observations here of a 

frozen bed at the start of a breakup at all sites 

upstream from site S5 do not unequivocally indicate 

immediately underlying permafrost, although the evidence is strong. The observed frozen bed material, even 

in the thalweg of the channel, conceivably could represent seasonal frost above an underlying unfrozen zone. 

As far as stream channel behavior during the breakup

period is concerned, the difference is not important. 

The typical case for at least all but the largest coastalplain streams in their downstream sections is probably 

that observed in the Shaviovik River, 35 km east, beneath which winter freezing is complete but permafrost temperatures are significantly higher than at locations away from channels (Brewer, 1958, p. 26).

Streamflow in Alaska north of the Brooks Range is 

measured at present in three major streams—the Sagavanirktok, Kuparuk, and Putuligayuk Rivers (see 

measurement sites in fig. 1). Flow records since 1970- 

71 (fig. 2) indicate that the maximum yearly discharge 

in these streams usually occurs during the breakup 

flooding and that the breakup peak has occurred most 

commonly in early June. This flow pattern is characteristic of a "nival" runoff regime (Church, 1974), in 

which the spring breakup flood is usually the dominant 

hyrdrologic event. The term "breakup flood" is used in 

this discussion as an equivalent of "spring," "nival," or 

"snowmelt" flood—terms that are likewise in wide use.

The hydrographs of the main runoff periods in figure 

2 show the variations in amount and distribution of 

runoff seasonally, from year to year, and between stations. The available records of flow, chemical analyses, 

water temperatures, and suspended-sediment loads are 

reported in the annual series of the U.S. Geological 

Survey, "Water Resources Data for Alaska." Each of 

the five study watersheds is discussed below.

TABLE 1.—The study sites and their characteristics

[Extent of glaciation mainly after Keller, Morris, and Detterman (1961) and Hamilton and Porter (1975). Where late Pleistocene glaciation is indicated,

earlier glaciation may also have occurred]

Stream

Atigun River

— — Do——— ———

——— _Do __________

Sagavanirktok 

River. 

——Do _ _______

— — Do———— —

___ Do— __ _____

______Do __________

Happy Creek __ _

— — Do———— —

River.

Site 

No.

Al

A2

A3

- Kl-6

SI 

S2

S3

S4

S5

HI

H2

PI

Drainage 

area 

(km2)

435

600

760

132-146

4,680 

4,830

5,870

12,200

tary 

88

89

*450

Median diameter 

of bed 

material 

(mm)

22

.41

.38

56-77

80 

86

45

18

15

51

35

14

Channel 

slope 

(m/m)

0.00088

^OOl

i.OOl

.0059

.0032 

.0023

.0025

.0018

.00053

.009

.01

.00047

Channel 

pattern

____do __

___ _do ____

Braided

braided 

__ do ——

Meandering

Glacial history

Multiple late

Pleistocene

———do ______

_— __do

Pre-late

Pleistocene

Single late 

Pleistocene 

Possible pre-late

Pleistocene

T Tn orl Q _"*i Q -.MI

—— __do

———do

———do ______

———do ______

___ do ______

Site location

18.2 km upstream from

pipeline crossing. 

1.1 km upstream from

pipeline crossing.

pipeline crossing. 

0.06 km upstream to 2.4

km downstream from 

pipeline crossing. 

25.4 km upstream 

from Lupine River. 

9.3 km upstream

from Lupine River.

from Lupine River.

taries bifurcate.

coast on west channel.

Sagavanirktok River.

Sagavanirktok River.

gaging station.

'Estimated.
DESCRIPTION OF STREAMS

15

September

SAGAVANIRKTOK RIVER 

40 -| 

Q

I

LU

w 120 -i

DC 

LU

Q

Z

O 2000 -

O

W 1500-

Sj 1000 -

CL

£ 500 -

H 0-

5 1

n

II

I

\ 1 000 -

I \

No / \ 500-

record V.

5 15 15 1

A

X^V-_

I I I I I 

5 15 15

May June July May June July

y 1971 1974

OQ

p. 1500 -T

U

E 1000 -

Q 500 -

CC

< 0 -

A 1000-

jvj \ 500 -

/ V_^^

' ~~- n

ii ii i u

/ ^\v_^v_^

/ *^ ^

i i i i i

5 15 15 15 15 15 15

W May n June July May June July

Q 2000 -

> 1500 -

^ 1000 -

< 500-

LU

^ 0 -

1972 1975

1500 -

1000-

/ \ 500 -

/ V/V _

ii II 1 u

A

J I

/ ^ _ i i i i i

15 15 15 15 15 15

May June July May June July

1973 1976

80-

40 -

m 0

15 15 

June July 

1970

120 -i

80-

40 -

120 -,

80 -

< 40

15 15 

June July 

1971

80-i

40-

15 

June 

1974

15 

July

15 

120-, June,

15 15 

June July 

1972

80-

40-

1975

Jufy

KUPARUK RIVER

15 15 

June July 

1973

15 15 

June July 

1976

PUTULIGAYUK RIVER

FIGURE 2.—Seasonal hydrographs of streams in arctic Alaska. Drainage areas at gaging stations: Sagavanirktok River, 5, 719 

km2; Kuparuk River, 8,107 km2 ;Putuligayuk River, 456 km2 ; Discharge volumes include both water and sediment.
EFFECTS OF PERMAFROST ON STREAM CHANNEL BEHAVIOR IN ARCTIC ALASKA

ATIGUN RIVER

A major headwater tributary of the Sagavanirktok 

River, the Atigun River occupies a broad glaciated valley in the Brooks Range. The upper section of the valley is a glacial outwash plain; downstream, the river 

meanders between generally cohesive banks in a valley flat where the study sites are located. The flat is the 

site of glacial-outwash deposition resulting from lake 

impoundment or reduced stream gradient caused by 

the late Pleistocene recessional moraines located north 

of Galbraith Lake. Channel slope in these reaches is 

now controlled by the head of a bedrock gorge into 

which flow of the Atigun River was diverted eastward 

to the Sagavanirktok drainage when its pre-glacial 

course as a tributary of the Itkillik River or Kuparuk 

River was disrupted by moraines.

Breakup of the Atigun River occurred in 1976 approximately on April 30. This unusually early breakup 

in comparison with that of nearby streams was in part 

due to the effect of wind-blown dust from the adjacent 

pipeline-construction corridor on the albedo of the valley snowpack. An initial peak on May 6 was followed 

by relatively constant lower discharges throughout 

May and then by successively higher peaks on June 5 

and 12.

A major summer rainstorm occurred in the headwaters of the Atigun River valley on June 30, 1976. 

Measured storm precipitation was small (<40 mm) at a 

station within the valley (Atigun Camp) and at a station across the crest of the Brooks Range (Chandalar 

Camp). Nevertheless, local intensities must have been 

substantially greater in order to cause the significant 

runoff that resulted. The peak discharge exceeded that 

corresponding to bankfull flow throughout the Atigun 

River valley and, over 80 km downstream, caused the 

peak of July 30-31 on the Sagavanirktok River hydrograph (fig. 2).

SAGAVANIRKTOK RIVER

The major stream east of the Colville River in arctic 

Alaska, this river drains a large part of the Brooks 

Range. In the foothills the channel pattern is generally 

sinuous or meandering and the banks are cohesive, but 

downstream from the junction with the Lupine River, 

and beyond the maximum extent of glaciation, the 

channel becomes progressively more complexly braided and the banks more noncohesive. At site S4 the 

river splits into two branches, with the main channel 

entering the Beaufort Sea about 12 km east of the west 

channel.

A discharge of less than 0.1 m3/s is characteristic of 

the Sagavanirktok River at the gaging station (fig. 1) 

through the winter months. Substantial flow over ice

began before May 3 in 1976, reflecting the early 

breakup on the tributary Atigun River. The initial 

peak of over-ice flow on May 6, evident in the hydrograph of figure 2, was largely the result of this flow. 

Breakup in the channel between sites SI and S3 did 

not occur until May 31, however, and breakup flooding 

did not peak in the same reaches until June 5. As has 

been historically typical of the stream since measurement began, flow remained at reduced but substantial 

levels throughout most of June and July in 1976, and 

rose to similar levels both on June 28 and in response 

to the summer rainstorm of July 30 in the drainage of 

the tributary Atigun River.

In two of six years the peak mean daily discharge on 

the Sagavanirktok River has occurred in response to 

rainstorms in July or August. In other years the more 

typical "nival" pattern prevailed, in which the breakup 

flood was the most significant flow of the year. The 

hydrographs (fig. 2) show that important summer 

storms have been relatively common in the watershed 

upstream from the gaging station.

KUPARUK RIVER

Although the headwaters of this stream are confined 

to the foothills .of the Brooks Range, it is the largest 

drainage reaching the arctic coast between the Colville 

River on the west and the Sagavanirktok River on the 

east. The study sites are in the foothills where the 

Kuparuk River meanders across a 0.7-km-wide valley 

flat with cohesive alluvial fill and scattered areas of 

coarse glacial detritus in which sorted polygons have 

developed.

Flow over ice at the study sites began in 1976 approximately on May 2, but breakup did not occur until 

May 23. Breakup flow first peaked on May 26 and was 

followed by a second peak of nearly identical stage on 

June 6. The discharge was greatly reduced by the time 

of final observations on June 13. Much of this early 

flow in the upper foothills was not recorded at the gaging station near the mouth (fig. 2), probably because of 

storage in the form of icings at intermediate locations. 

A sharp temperature gradient existed northward 

across the coastal plain throughout May.

At the gaging station near the mouth of the Kuparuk 

River, the characteristic flow pattern since 1971 has 

been one of peaks in breakup flow during early June, 

followed by reduction in flow to low summer levels by 

late June. In 1976 breakup at the gaging station occurred on June 3 or 4, and the peak discharge was 

measured on June 15.

In contrast with the Sagavanirktok River, the 

Kuparuk River drainage yields much higher rates of 

runoff per unit area during breakup, and significant
DATA COLLECTION

flow is largely concentrated in the June breakup period 

(fig. 2). Breakup occurs more uniformly in the Kuparuk 

River basin than in the Sagavanirktok River watershed because of the less diverse relief in the former. In 

addition, the runoff from summer rainstorms, which 

are most intense in the mountainous southern sections 

of the Sagavanirktok River basin, is greatly retarded 

by the tundra vegetation of the foothills and coastal 

plain once the active layer begins to thaw. The thick 

sod mat and hummock surfaces of the foothill and lowland tundra create both a great capacity for absorption 

of precipitation and a high resistance to flow. Thus, 

rainstorms in the Kuparuk River basin have not 

caused significant peaks during the period of flow 

measurement.

HAPPY CREEK

This small drainage of 89 km2 enters the 

Sagavanirktok River from the west at Happy Valley 

Camp. It is typical of the foothill area; little valley flat 

is evident and the channel is broadly sinuous or meandering, with partial bedrock control. The bed material 

is coarse gravel, and the banks consist of a tundra mat 

over a thin veneer of gravel-rich soil.

Meltwater discharges in Happy Creek rise and recede rapidly. In 1976 breakup flow increased from 0 to 

39.4 m3/s between June 3 and June 6 and receded to 

low levels by June 8. Annual peak discharges before 

1976 ranged from 2.5 to 16.2 m3/s since crest-stage 

measurements began in 1972 (table 2).

PUTULIGAYUK RIVER

The drainage of this stream is confined to the Arctic 

Coastal Plain, an area of low relief with numerous 

oriented lakes similar to those so intensively studied in 

the vicinity of Barrow. The meandering channel has 

bed material of fine gravel and banks of cohesive silt 

and clay with soil development overlying fine gravel.

Measurements since 1970 indicate that the 

Putuligayuk River generally peaks rapidly, rising 

from zero flow during 1 to 2 weeks in early June and 

falling continuously to low summer levels in about the

TABLE 2.—Peak discharges measured at crest-stage gage on Happy 

Creek. Drainage area above gage is 89 km2

Year Date Discharge

(m3/s)

1972

1973

1974

1975

1976

Unknown

Unknown

During May

June 6

June 6

14.9

2.5

1 11.3

16.2

39.4

'Estimated; ice effects.

same length of time. The same highly attenuated 

summer flow regime that is characteristic of the 

Kuparuk River basin is present to an even greater degree in the Putuligayuk River basin. To the hydrologic 

effects of the tundra are added the effects of poorer 

drainage and the presence of standing water in a 

greater proportion of the watershed.

Summer storms historically have caused little 

runoff; through 1976 no discharge in excess of 3 m3/s 

has been recorded following recession of the breakup 

flood. Significant flow, mainly over ice, began in 1976 

approximately on June 5, and the peak discharge occurred on June 18. Observations of the channel were 

ended on June 12.

DATA COLLECTION

MEASUREMENT TECHNIQUES

Erosion rates in cohesive bank materials were 

measured with erosion pins (6 x 178 mm) emplaced 

horizontally at appropriate intervals in a bank cross 

section (see Wolman, 1959, p. 207). The sites selected 

were those at which maximum amounts of bank erosion could be expected, such as the outside of a meander bend and other locations where nearshore current 

velocities were high and banks were known from field 

evidence to be eroding. Erosion rates at some sites were 

sufficiently high to require the periodic replacement of 

a few pins.

Erosion rates in noncohesive bed and bank materials 

were determined in a similar manner through measurements of the channel surface relative to lengths of 

reinforcing bar (9.5 mm x 1 m) driven to depths that 

penetrated frozen material. Scour measurements were 

made in the same way. Horizontal control was established by tape survey at the time of emplacement; vertical control was made by referencing the tops of the 

reinforcing bars and the bed surface to stage, which in 

turn was recorded with improvised staff gages. As the 

noncohesive beds and banks thawed during the study 

period, it was necessary periodically to drive the bars 

deeper into frozen material, each time referencing the 

new height for vertical control.

Rates of migration of the frozen-unfrozen interface 

were obtained by driving lengths of reinforcing bar 

into the bank or bed at the location of each erosion pin 

or bar. This procedure worked well after verification by 

excavation and repeated measurements. With practice 

it was possible to distinguish the nature of whatever 

rigid surface was encountered—whether frozen bed 

material, segregated ice, or a large clast—by the "feel" 

of the probe when driven against the obstacle. The difficulties described by Mackay (1977, p. 327) in probing
8 EFFECTS OF PERMAFROST ON STREAM CHANNEL BEHAVIOR IN ARCTIC ALASKA

the depth of the active layer in fine-grained soils with 

little ground ice were not encountered. In both cohesive 

and noncohesive sediment the boundary could be determined accurately. Where the bed material was too 

coarse to use a probe, thaw rates were determined by 

successive excavations that were offset sufficiently to 

avoid any thermal disturbance from the preceding excavation.

RATES OF THAW IN BEDS AND BANKS

The variation in thaw rates with sediment size was 

determined in a section of the Atigun River along 

which bed material changed in size but discharge and 

thermal conditions were relatively constant. Figure 3 

shows the variation in thawing rate of submersed bed 

and bank material plotted with change in midday 

water temperature. Each data set represents the condi60

50

40

30

20

10

0

50

40

30

I 20 

a.

UJ

D 10

r o 

H 20

10

0

Dm = 22 mm (site A1)

Dm = 0.38 mm (site A3)

Dm < 0.0625 mm (site A2)

m DC LU

51 O Q

5LU <

LU Q D:

10

tc~ h

IU UJ 2

I- tC ^ -5 

O

*s 10 15 20 25 30 

May

1976

9 14 19

June

FIGURE 3.—Thaw depths in submersed bed and bank material of 

differing texture at three sites on the Atigun River. Water temperature was measured at site A3. Size was determined by field measurements at site Al, by sieving at site A2.

tions at a single erosion pin or bar in the channel, and 

one that is typical of many similar sets of data, each 

corresponding to the location of a pin or bar in the cross 

section. Sediment in the silt (0.004-0.0625 mm) and 

clay (less than 0.004 mm) size ranges at site A2 is bank 

material; the gravel (over 2 mm) at site Al and the 

sand (0.0625-2 mm) at site A3 are bed material. At the 

latter two sites major sections of the banks are composed of the same material as the beds, and the submersed thaw rates in those parts of the banks were 

similar to the rates measured in the beds.

< H

CCZ 

UJ LU 

0. O 

5

LU W 

h UJ

I I I I I I I I

20

15

10

5

0

-5

-10

10 15 

May

20 25

1976

30 9 

June

14 19

FIGURE 4.—Thaw depths in emersed bed and bank material of differing texture at three sites on the Atigun River. Air temperature 

was measured at Galbraith Lake.
DATA COLLECTION

A similar set of measurements was obtained for 

emersed sediment of the same sizes at the same three 

locations. Each point (fig. 4) is also typical of a group of 

points at each site. Air temperatures shown are the 

daily maxima at Galbraith Lake, 10 km northwest of 

site A3.

Texture proved to be the dominant factor in the variations in thawing, given the approximately equal flow 

and thermal conditions at the three sites. Thaw rates 

were proportional to grain size whether sediment was 

exposed to water or air. In the absence of other factors, 

the effects of the latent heat of ice should produce the 

lowest thaw rates in sediment with the greatest ice 

content. That this was not seen probably reflects the 

efficiency of the finer, more cohesive sediment, once it 

has thawed, in insulating the frozen interface and 

thereby retarding further thawing.

Position in the cross section—differing flow depths 

and whether the measurement was in bed or bank 

material—was of relatively minor importance in explaining variations in thaw rates of submersed sediment. Emersed thaw rates proved to be moderately 

greater than submersed rates for each of the three textures observed during the period. When occurring coincident with erosion, however, the potential for much 

greater rates of thawing in submersed material is obvious.

Although air temperature trended downward between May 1 and 21, the thaw depths in emersed, noncohesive material increased rapidly during that period.

It is likely that incident solar energy increased 

throughout that time interval. At only one point at one 

site (A3) was any decrease in thaw depth observed—in 

this case associated with the unusual cold spell of May 

20-21. At other points the active layer thickened continuously.

THAW AND EROSION RATES COMPARED

The question of permafrost control of stream behavior hinges on the relative amounts of thaw and erosion 

at individual points. If erosion uniformly proceeds 

more slowly than the observed rates of thaw, for example, there is clearly no direct effect of frozen bed or 

bank material on erosion, and thus on stream behavior. If frozen bed or bank material is in contact with the 

flow after breakup, it can be assumed that erosion is 

being retarded by the frozen condition of the material.

Figures 5 and 6 summarize the measurements of 

thaw and erosion, with the data divided according to 

whether the material is cohesive or noncohesive. The 

distinction in sediment type is based on field examination. The measurements are confined to the time interval between breakup and the peak of breakup flooding 

in each stream. A large number of points representing 

no erosion and minuscule amounts of thaw are not 

shown. Such points are typical of the smaller streams.

The results show that the general circumstance is for 

erosion to proceed more slowly than thaw, with the 

exception of two specific locations where sediment was

80

60

40

20

Points at apexes of 

thermo-erosional niches.

80

V) 60

DC

Ul I01

40

20

All points on line are at apexes 

of thermo-erosional niches

20 40 60

THAW DEPTH, IN CENTIMETERS

80 20 40 60

THAW DEPTH, IN CENTIMETERS

80

FIGURE 5.—Plot of erosion versus thaw for points in noncohesive bed 

and bank material. Points without detectable erosion (41) are excluded.

FIGURE 6.—Plot of erosion versus thaw for points in cohesive bed and 

bank material. Points without detectable erosion (25) are excluded.
10 EFFECTS OF PERMAFROST ON STREAM CHANNEL BEHAVIOR IN ARCTIC ALASKA

frozen at the bed surfaces. These were (1) the thalwegs 

of streams—sites of the largest amounts of vertical 

scour; and (2) the innermost surfaces, or apexes, of 

thermo-erosional niches—sites of the most intense lateral erosion. Most such points were not frozen; yet the 

only points that remained frozen after breakup on the 

larger streams were at those locations. In plotting figures 5 and 6, it was assumed that the retreat of a solidly frozen surface represented equal amounts of thaw 

and erosion.

That maximum rates of erosion are less in cohesive 

material than in noncohesive material is evident from 

a comparison of figures 5 and 6. This difference is not 

the sole function of the lesser rates of thaw in cohesive 

sediment seen in figures 3 and 4. Were this the case, 

control of erosion by permafrost in cohesive banks 

would be greater than in noncohesive banks, and in 

figures 5 and 6 the points representing cohesive sediment would be located nearer the line representing 

equal amounts of thaw and erosion. Although a comparison of the figures suggests that permafrost control 

of erosion is more likely in cohesive sediment, the 

greater resistance of the finer sediment to erosion must 

also be attributed to the properties that cause that size 

of bed and bank material to thaw more slowly—the low 

permeability that produces low rates of thawing is a 

function of the cohesive bonding of clay minerals that 

increases the resistance to erosion.

Measurements on the larger streams (Sagavanirktok River sites Sl-3 and all Atigun River sites) 

showed that direct permafrost control of erosion in 

most cases ended well before the peak discharge occurred. The situation was much different for smaller 

streams like the upstream Kuparuk River and Happy 

Creek, where frozen material remained at or near the 

surface, especially in cohesive bank material, after recession of the breakup flood.

THERMO-EROSIONAL NICHING

Bank undercutting that can be characterized as 

thermo-erosional niching is the dominant cause of 

bank retreat in the streams with banks of mainly cohesive material, a category that includes nearly all 

streams with sinuous or meandering channels in the 

study area. In typical thermo-erosional niches in 

homogeneous deposits, the top of the niche develops at 

sustained levels of high flow. The development of 

niches is discouraged by heterogeneity of banks and by 

fluctuations in stage. Although the bank undercutting 

in most study streams is less regular than that illustrated from a noncohesive bank (fig. IB), both vertically and longitudinally, the process of bank undercutting by concomitant thaw and erosion is the same. It

should be noted that bank undercutting and subsequent failure are unique in the arctic setting only in 

their prevalence and degree of development. The increased strength provided by permafrost in the banks 

permits greater amounts of undercutting, which in 

turn produces larger slump blocks.

The niching process is more rapid in the layers of 

sand and gravel that are sporadically interbedded in 

cohesive banks (fig. 7A), as would be expected from the 

greater rates of thaw and erosion in the coarser sediment. Where banks are uniformly of silt and clay, 

niche development may be concentrated at the base of 

the cohesive fill, within the underlying gravel. A corollary to these observations is that the relative susceptibility of cohesive banks to erosion can be assessed from 

the stratigraphy of the alluvial fill, observable either 

in bank cross sections at times of low flow or by excavation.

Niching is also present in braided channels. Where 

channels braid in predominantly sand-size material, 

the niches are developed most dramatically (fig. IB} 

and correspond in morphology to the previously described examples (Gusev, 1952; Walker and Arnborg, 

1966). In small channels in sandy reaches of the Antigun River an entire anabranch was seen to disappear 

within a niche at some sites. With coarsening of bed 

and bank material above the size range of sand, niching becomes progressively less well developed. Consequently, the process is of lesser importance in the 

braided streams than in those with sinuous or meandering patterns.

BANK SLUMPING AND VEGETATION EFFECTS

The failure of cohesive banks in response to undercutting is generally a gradual process over hours or

Base of 

thawed layer/*

1 METER

Gravel

Bank of silt and clay interbedded with gravel. Sagavanirktok River, site S2

*- 5-20-76

Bank of sand. Atigun River, 

near site A2

FIGURE 7.—Cross sections of undercut banks. A, Cohesive. B, Noncohesive. Dates indicate water levels at time of survey.
DATA COLLECTION 11

several days. Sloughing of small blocks from the vertical face of noncohesive banks (Outhet, 1974, p. 307) is 

an important mode of failure as is the sudden collapse 

of both types of banks upon fracture through frozen 

material weakened by thawing (Walker and McCloy, 

1969, p. 76). True slumping is not particularly common, but the term may be appropriate, as a simplification, for the assemblage of mass-wasting processes involving bank failure. Failures probably tend to be 

accelerated during flow recession (Inglis, 1949, p. 152), 

by the reversal of seepage forces and the loss of the 

resisting pressure provided by flow in the channel.

Ice wedges associated with the raised-edge polygonal 

ground on the coastal plain may thaw rapidly and provide planes of weakness along which bank failure can 

occur. Segregated ice masses thaw more rapidly than 

most cohesive sediment but appear to thaw less rapidly 

than some highly erodible noncohesive sediment. This 

seeming contradiction to the effects of the latent heat 

of ice on thaw rates probably reflects the ability of the 

cohesive sediment to insulate the thaw interface. Ice 

wedges occur in all flood-plain deposits of the study

> 00

100

50

DC BCI "-

ujO WM 

Q- 0 

UJ 20 

, .cog 

rr DC HI HI < 

ujiil DC HI DC - -— 10

O 0 

LU 20

w -J til <

HI O

(- 300

>• Hi" HI

200

« 100

I I I I I I I I I I I I I

\ \ i i i i i i i r 

SiteSI

i i i 1 i i i i

i i i i i

Gaging station

29 30 31 1 2 3 4 5 6 7 8 9 10 11 12 13 14

May June 

1976

FIGURE 8.—Plot of percentage of bank failures at site SI, water temperature at site Si, maximum daily air temperature at Happy 

Valley Camp, and mean daily discharge at gaging station on the 

Sagavanirktok River. Discharge records are estimated for the 

period shown.

area but are of quantitative importance only in the 

coastal plain.

The bonding provided by the tundra root mat can 

significantly retard slumping. It is also possible, because cold climates retard the decay of roots, that lateral erosion may be retarded directly by root structures 

to an unusual extent (Smith, 1976). Because the shallow root depth of the modern tundra vegetation is limited by the depth of active layer (less than 0.3 m in 

vegetated cohesive material), roots are not important 

in limiting bank erosion in the larger streams of the 

study area. Older organic layers at depths up to 1.0m 

are preserved in a frozen condition and may retard erosion directly in the lower banks of the smaller streams. 

Bank erosion was limited by vegetation during the record discharge in Happy Creek, but other important factors were sediment size and depth of thaw. Vegetation 

on the banks of Happy Creek was important in the 

indirect sense that it greatly increased the insulation 

properties of the bank surfaces and thereby decreased 

both the thaw rate and subsequent erosion.

The direct role of vegetation in most study streams 

was mainly that of retarding the failure of undercut 

banks to varying degrees. The tundra mat was commonly observed to remain attached to the bank, folded 

downward into the stream on the top of a failed bank 

and thus acting temporarily to prevent further erosion, 

as described by Walker and McCloy (1969, p. 74).

VARIATIONS IN TIMING OF 

BANK SLUMPING

The percentage of the reach showing slumping of the 

bank at site Si on the Sagavanirktok River is plotted 

with time in figure 8, along with discharge and associated air and water temperatures. The sudden increase 

in slumping seen on June 4 was related to niching that 

was accelerated by rising water temperatures on June 

3 and 4. There was clearly little time lag between the 

niching and bank failure at this site.

The importance of undercutting and subsequent 

bank failure in the study streams varies with size of 

the drainage area when the comparison is limited to a 

single channel pattern and sediment type—that is, to 

sinuous and meandering channels with banks of cohesive sediment. Figure 9 shows the prevalence of slumping along particular sinuous and meandering reaches 

within the time period between breakup and the 

breakup peak. Each reach represents approximately 

the outside bank of a single meander bend between the 

points of inflection.

Figure 9 indicates that the rise of breakup flooding 

was vastly more important on the Sagavanirktok River 

than on even the Atigun River, the next largest. No
12 EFFECTS OF PERMAFROST ON STREAM CHANNEL BEHAVIOR IN ARCTIC ALASKA

bank failures of any kind were observed in association 

with breakup flooding on Happy Creek, the smallest 

drainage studied. And, at the Kuparuk River sites only 

a few small failures were initiated near the end of the 

period. Limited observations on the Putuligayuk River 

revealed only a slight amount of bank slumping beginning as of June 12, and similarity in pattern and bank 

material suggested the likelihood of behavior like that 

of the study sites on the Kuparuk River. The causes 

and implications of this variation with drainage area 

are discussed at length in a subsequent section.

Referencing the data of figure 9 to the period between breakup and the peak of breakup flooding provided a geomorphologically and hydrologically 

significant time interval over which to compare the 

importance of slumping in different watersheds. The 

amount of slumping proved to be a valid index of lateral erosion. Had absolute values of lateral erosion 

been averaged in each sample reach and plotted in 

figure 9, the results would have been similar, even 

though the maximum amount of erosion and the proportion of reach affected by that erosion do not necessarily correlate. The length of reach affected by slumping is a variable easily determined with successive 

ground photographs, compared to values of total erosion that can be determined only at surveyed cross sections or with aerial photography.

100

Sagavanirktok River 

(SiteSD

0 25 50 75 100

PERCENTAGE OF TIME INTERVAL BETWEEN BREAKUP 

AND PEAK OF BREAKUP FLOODING

FIGURE 9.—Plot of percentage of sample sinuous or meandering 

reaches showing bank failure, versus percentage of time interval 

between breakup and peak of breakup flooding. Only partial data 

were collected from the Putuligayuk River.

In appraising figure 9, it is worth noting that the 

time lag between undercutting and slumping of a bank 

can vary with a number of factors. It was verified, however, that the lack of slumping at the sites in the 

smaller watersheds was due to an actual lack of lateral 

erosion and not to a time lag between niching and 

failure.

SAGAVANIRKTOK RIVER, SITES S4 AND S5

It was noted in the introduction that cohesiveness, as 

a function of the particle-size distribution of bed and 

bank materials, is the main factor in determining bank 

erodibility, and that erodibility in turn can be considered to be the major variable in determining channel 

pattern. Consequently, the association between channel pattern and erosion rates corresponds to that 

between bed material and erosion rates—the greater 

erosion rates occur in braided channels, which are 

characteristically formed in noncohesive sediment. Because of anomalous local conditions, downstream 

reaches of the Sagavanirktok River were exceptions to 

both this relation and the variation with drainage 

area.

The time of breakup in watersheds north of the 

Brooks Range is the reverse of that expected from their 

altitude, with the lowest and northernmost basins 

being the last to yield flow. Dense coastal fog remained 

longer than usual in 1976 and kept temperatures low 

through mid-June in an area as far as 60 km from the 

Arctic Ocean. As a result, the response of channels to 

breakup flooding in the major, through-flowing rivers 

depended, as it would to a lesser degree in more typical 

years, on local conditions as well as on the thermal 

input of flow from the earlier breakup farther south.

From June 5 to 12 the bank material and most bed 

material in braided reaches on the lower 

Sagavanirktok River (sites S4 and S5) remained frozen 

although the peak of breakup flow had passed on June 

5 or 6. On June 11 at site S4 there was little evidence of 

any channel erosion, and neither submersed nor 

emersed sediment had thawed more than several centimeters in depth. Flow depth at site S5 prevented 

measurements in the deeper channels, but in observable respects the situation was similar to that at site S4. 

The high upstream water temperatures (5° to 7°C at 

sites S1-S3) were rapidly dissipated in the shallow, 

complexly braided channels at these two sites on the 

fog-bound coastal plain. Water temperature was only 

0.6°C at site S4 at the same time it was over 6.0°C 

upstream at site SI (fig. 8).

This variation in channel response with climate illustrates the type of local complication that can dominate any general relations like those based on bed
DATA COLLECTION 13

material or drainage area. It also emphasizes the major 

variations that are possible along the larger drainages 

that cover considerable latitude in the Arctic.

BED SCOUR

The effect of frozen bed material on scour was 

studied with repeated profiles of the bed and thaw surface across the channel at site A3 on the Atigun River. 

Bed material at this site is fine sand (median diameter =0.3 8 mm) within the size range that shows large 

increases in bed-material discharge with low water 

temperatures (Colby and Scott, 1965, fig. 19). The 

channel at that point had a sinuous pattern during the 

observation period, with tendencies toward braiding at 

the lower end of the range of observed discharges. 

Throughout most of the observation period, a single 

channel from 25 to 30 m in width and with a depth of 

less than 1 m was present.

Figure 10 portrays the channel change and the increases in thaw depth between May 20 and 31. Before 

May 26 little lateral movement occurred, and the bed 

cross section was about the same on May 7 as that 

shown on May 20 and 26. As late as May 20, frozen 

material was sporadically exposed at the bed surface in 

the thalweg (point a in figure 10). Sediment transport 

during the pre-May 20 period was evident in the form

of dunes moving across a frozen substrate.

The importance of figure 10 is that scour was being 

retarded by frozen bed material as late as 3 weeks after 

breakup when midday water temperatures had risen 

as high as 2.5°C (fig. 3). That the scour did not increase 

once the thalweg had thawed to greater depths, shown 

by the profiles of May 31, may reflect the greatly increased load in the stream that accompanied the thawing process and was promoted by the combination of 

fine-grained bed material and low water temperatures.

Figure 10 also shows the sudden increase in lateral 

mobility of the channel that occurred after May 26 and 

the corresponding change in configuration of the thaw 

surface. These changes occurred under conditions of 

relatively constant stage, excluding small diurnal 

changes.

CHANNEL PROCESSES DURING REMAINDER OF 

1976 RUNOFF SEASON

The study sites were resurveyed in May 1977 to determine the channel changes that occurred during that 

part of the 1976 runoff season following breakup flooding. At the time of the 1977 survey, breakup either had 

not occurred or was in an early stage, and the channels 

were as they existed at the time of freezeup in September 1976. The resurvey was surprising in the

so

60

t; 40

UJ

O

20

Stage: -^—5-20-76 5-26-76 -^5-31-76

10 15

METERS

FIGURE 10.—Cross sections of bed surfaces and thaw surfaces in the thalweg at Atigun River site A3. 

Stages indicated are water levels on dates of survey.
14 EFFECTS OF PERMAFROST ON STREAM CHANNEL BEHAVIOR IN ARCTIC ALASKA

small-to-nonexistent degree of change that was observed to have occurred at most sites in 1976 after at 

least partial recession of the breakup flow.

The absence of change was striking in the smaller 

watersheds like Happy Creek and the upstream 

Kuparuk River. In the latter watershed individual 

slump blocks had not moved detectably from positions 

recorded by photography following recession of 

breakup flooding in June 1976. Slumping observed 

along the Putuligayuk River was comparable to 

amounts recorded during the breakup period on the 

Kuparuk River (fig. 9), but, because observations were 

not continued through the complete period of breakup 

flooding on the Putuligayuk River, the failures there 

could not be ascribed to the flooding period or the remainder of the runoff period.

The exceptions to the observations of little change 

were the sites on the Atigun River and the downstream 

part of the Sagavanirktok River.

ATIGUN RIVER

The rainstorm flood of July 30 was a major instrument of lateral and overbank erosion throughout the 

entire length of the Atigun River valley. The peak flow 

of the July 30 flood exceeded bankfull discharge 

throughout the observed sections of the river. It locally 

approached, but did not exceed, the levels of the 

maximum evident flood measured by Childers and 

Jones (1975, table 1 and p. 16-17). On that basis, it can 

be estimated that the peak discharge was about 300 

(±50) m3/s in the reaches near site Al. The recurrence 

interval of the flow could be approximated within the 

range of 5 to 15 years.

Amounts of lateral erosion were determined by comparing ground photographs at the study sites and in 

nearby reaches. Lateral erosion was as much as 4.9 m 

in the case of a terrace level 1.4 m above the flood plain. 

The banks were eroded by a measurable amount at 

most locations where erosion would be expected— 

outsides of bends and, in straighter reaches, points of 

impingement of the thalweg against channel banks. 

Overbank flow on vegetated areas of the flood plain 

locally stripped the tundra mat from surfaces shallowly underlain by gravel.

An impressive feature of the flood results was the 

resistance to erosion of uniform cohesive deposits, especially where a cutbank was absent and vegetation 

extended to the level of the low-water channel.

SAGAVANIRKTOK RIVER

The situation at sites on the sinuous or meandering 

portion of the channel (SI and S2) in May 1977 was 

similar to that observed after partial recession of

breakup flooding in mid-June 1976. Some movement of 

individual slump blocks had occurred. Only a few had 

been removed by erosion, and many remained folded 

down into the channel with the intact tundra mat acting to prevent subsequent erosion. The small amount 

of change was unexpected, particularly so in light of 

the fact that flow continued at significant levels 

throughout much of the remaining runoff season 

(fig. 2).

Site SI was observed daily for a week during the rise 

of the breakup flood in 1977 (May 25-June 1). During 

that time it was apparent that the previous year's 

slump blocks were acting to inhibit lateral erosion. 

This result raises the possibility that more than one 

year may be necessary to remove the slumped sections 

of a bank and prepare it for another large increment of 

lateral erosion by means of niching and failure.

The preceding sequence of large change during 

breakup flow and little change during the rest of the 

runoff season was reversed at the sites on the braided 

sections of the river downstream (S4 and S5). There, 

measurable change in the position of most of the individual anabranches had occurred, indicating that the 

portion of the runoff season after recession of the 

breakup flood is probably the normal time of channel 

change. Channels at these sites were mainly frozen 

and unchanged throughout the breakup flooding of 

1976.

SEDIMENT TRANSPORT

With the frozen condition of the channels seen to last 

at least part of the way through breakup flooding, 

there are implications for the measurement and estimation of sediment discharge in arctic streams that 

involve questions of sediment availability. Sediment in 

transport has been divided conceptually for practical 

purposes into two components: the wash load (or finematerial load)—consisting of fine sediment the discharge of which reflects events upstream and is considered not to be a function of flow hydraulics; and bedmaterial load (or bed-sediment load)—generally 

coarser sediment of sizes present in the bed, the discharge of which is considered a function of flow hydraulics. The wash load in this concept must be measured 

by sampling; the bed-material discharge is commonly 

estimated for engineering purposes with relations involving flow and sediment parameters. If the bed is 

partly frozen, however, sediment may not be readily 

available at a point and the total load may thus become, like the wash load, a function of conditions and 

sediment availability upstream. This means that for 

some arctic streams at certain times the total sediment 

load can be determined only with measurements.
DATA INTERPRETATION 15

The effect of a frozen bed on sediment transport will 

be greatest for the coarsest fractions of bed material— 

those sizes most likely to move only as bedload, rolling 

or skipping along the bed surface. It will certainly be 

most noticeable in the smaller streams in which the 

frozen condition of the bed may extend through the 

entire breakup period. If such an effect is significant in 

the larger streams outside the coastal plain, it will 

probably be confined to the initial rise of the breakup 

flood. Although the coarse nature of the sediment (table 1) made direct measurement of bedload impractical, several lines of evidence can be used to determine 

first, when bed material was thawed enough to be 

mobile during the 1976 breakup flooding, and second, 

whether breakup flows were competent to move the 

available material.

At sites S4 and S5, on the Sagavanirktok River, the 

effects of a predominantly frozen bed and bank probably had considerable influence on sediment transport 

through the rise and for some time after the 1976 

breakup flooding. Thaw depths measured at upstream 

sites (Sl-3), however, indicated that the bed was probably capable of becoming mobile within 1 or 2 days 

after breakup. Consequently, bed-material transport 

there was mainly a function of the competency of the 

flow, and was entirely so later in the rise of the 

breakup flood as competence increased to the point 

where transport of some sizes present in the bed was 

possible.

On the Kuparuk River, observations indicated that 

no movement of cobble-size (64-256 mm) bed material 

accompanied the first breakup peak of May 26. Soundings with a probe during and the day after the peak 

revealed directly that cobble-size clasts were frozen in 

an immobile bed. However, extensive transport of the 

same material occurred during the second breakup 

peak of nearly equal discharge, thereby suggesting 

that the competency of the first flow was sufficient to 

cause transport. Both peaks were approximately 

equivalent to bankfull discharge.

The sand-size bed material of the Atigun River at 

site A3 was continuously transported to some degree 

throughout the breakup period. Data were inconclusive on whether transport of the gravel at site Al occurred before the peak of June 5, but the depth of thaw 

would have permitted movement had flow been sufficiently competent. Extensive movement of the coarser 

material occurred with the peaks of June 5 and 12.

A small amount of bed-material transport occurred 

with the sudden rise and sharp peak of breakup flooding in Happy Creek. The peak discharge of 39.4 m3/s 

was the maximum recorded since crest-stage measurements began in 1972. Transport of gravel near the 

lower end of the pebble-size range (4-64 mm) could be

observed in culverts below Happy Valley Camp. 

Coarser sediment was present in the bed yet was not 

transported in spite of the fact that flow near the peak 

was greatly in excess of that competent to do so, based 

on estimates of flow depth and slope and the recognized 

relations between size and shear stress on the bed 

(Baker and Ritter, 1975, fig. 1). Although observations 

of the bed during the short interval of the peak were 

not possible, it seems reasonable to suggest that the 

coarser material was frozen in place and that thaw in 

response to flow released only the finer gravel. Such 

appeared to be the case after flow receded.

DATA INTERPRETATION

EFFECTS OF PERMAFROST

Variations in thaw and erosion rates with size of bed 

material have obvious implications. Banks that are 

dominantly of cohesive material will erode more slowly 

than those with noncohesive sediment. More practically, the relative stability of streams with cohesive 

banks can be assessed by both the degree of interbedding with noncohesive sediment and the thickness of 

the cohesive fill relative to the depth of the channel. 

These measures reveal the susceptibility of a bank to 

thermo-erosional niching, the main mechanism by 

which the cohesive banks of artic stream channels retreat.

Permafrost was seen to retard bank erosion and 

scour in all the study streams; that is, frozen sediment 

was observed in direct contact with flow in channel 

thalwegs and at the apexes of thermo-erosional niches 

after breakup. The extent and duration of the effect, 

however, was strongly related to the size of the drainage basin when the observations were confined to sinuous and meandering channels with cohesive banks.

Data from all sites on the Kuparuk River and Happy 

Creek indicated that, if rise of the breakup flood is 

rapid and involves the lesser discharges and velocities 

of the smaller streams, little or no lateral erosion and 

scour may occur and bed-material discharge may effectively be prevented by the frozen bed. The data from 

the two upstream sites (SI and S2) on the 

Sagavanirktok River show that, if the breakup flood 

involves the higher discharges and velocities of the 

larger streams, lateral erosion and probably scour as 

well may be extensive. Bed-material discharge is probably only briefly retarded after the start of breakup 

flooding in such streams. During the lengthy, six-week 

interval between breakup and the peak of flooding in 

the Atigun River, a stream of intermediate size, approximately the first half of the interval was characterized by frozen bed material and the second half by 

general channel mobility.
16 EFFECTS OF PERMAFROST ON STREAM CHANNEL BEHAVIOR IN ARCTIC ALASKA

The variation in lateral erosion and thaw rates with 

watershed size reflects a basic variation of those processes as functions of stream power (proportional to the 

product of discharge and slope). The net result is doubtless highly time dependent as well, reflecting the 

shorter duration of the flooding in the smaller watersheds and the commonly lower water temperatures observed during breakup flooding in those watersheds. 

Watershed size is merely an easily determined factor 

that explains the variation in the sense of an independent variable in a regression analysis.

That there are definite differences in stream behavior related to size is indicated by the almost total absence of bed and bank processes before the breakup 

peak on the smallest streams. This observation is 

significant in the case of Happy Creek because the 

peak discharge in 1976 was by far the largest of record 

(table 2), suggesting the improbability of contradictory 

results in previous years. In streams like Happy Creek, 

the critical level of thermal input necessary for any 

significant bed or bank processes may not be reached 

before the breakup peak or even during the breakup 

flooding. It is likely, therefore, that the channels of the 

smaller streams respond significantly only to major 

floods, in contrast to those of the larger streams which 

plainly are modified by the annual breakup flooding.

A corollary to these conclusions is the effect of permafrost on lateral erosion once the peak of breakup 

flooding is past. In most of the larger channels thaw 

depths had increased to the point where little direct 

effect of permafrost on bank erosion or scour would 

occur in the absence of flood flows during the remainder of the runoff season. In fact, the effect of permafrost 

during the period after the breakup peak may be to 

facilitate the erosion of cohesive banks by assuring

OF 3CENTAGE SAMPLEREACH SHOWFAILURE ATBREAKUP PEAK BANK J> O)ONJ00 

_^O O

1 1 1 1 1 1 1 1

Happy Creek

i i i i i i m

I I I I I I I I

Sagavanirk

Atigun 

Kuparuk R' ver 

River °

O 1 1 1 1 1 1 1 1

i i i i i i 1 1

(SiteSD 

tok River 

(SiteS2)0

i i i i i i 1 1

0 100 1000 10,000

DRAINAGE AREA, IN SQUARE KILOMETERS

FIGURE 11.—Plot of percentage of sample reaches showing bank failure at peak of breakup flooding, versus drainage area.

saturation of the thawed material. Moisture content 

was found by Wolman (1959, p. 211-212) to be the 

main factor not directly related to stage that increased 

the erosion of cohesive banks. Permafrost increases the 

moisture content of banks in two ways—the direct 

thawing of permafrost acts as a continuous source of 

moisture through the summer, and the base of the active layer functions as a barrier that will prevent the 

subsurface infiltration of runoff.

THE VARYING IMPORTANCE OF 

BREAKUP PROCESSES

The plot of drainage area and the proportion of a 

study reach affected by slumping before the peak of 

breakup flow (fig. 11) illustrates at least a partial answer to why investigators have ascribed such different 

importance to the breakup period in terms of channel 

processes. That the most visible channel process in all 

arctic streams—bank slumping—can vary greatly with 

the size of the stream is a probable component of the 

explanation.

As noted in the introduction, a complete spectrum of 

conclusions as to what part of the runoff period is the 

most significant has been obtained. The results of 

study of large streams like the Colville River (Walker 

and Arnborg, 1966), the Mackenzie River (Outhet, 

1974), and the upstream channel of the Sagavanirktok 

River in the study area would suggest that breakup 

processes are the dominant erosional events of the entire runoff period. At the other extreme, results from 

the relatively small drainages of Banks Island (Miles, 

1976) or the smaller streams of the study area would 

indicate that breakup flooding was of comparatively 

minor importance.

Obviously, a simple correlation like that of figure 11 

is not the complete answer to why some stream channels change significantly during breakup and others do 

not. Variation with other factors may obscure that relation. The effect of climate in the case of the sites on 

the lower Sagavanirktok River is an example of only 

one such possibility, in which channel changes on a 

large stream could be slight during breakup but pronounced thereafter. One can envision a similar set of 

circumstances that might even delay significant channel change until the later stages of the summer runoff 

period. Slumping on the lower Lena River is in fact 

concentrated in the late summer (Abramov, 1957), but 

this may only reflect a time lag between niching and 

slumping (Czudek and Demek, 1970, p. 105).

SUMMARY AND DISCUSSION

The preceding discussions provide first approximations of answers to the two problems posed at the be-
DATA INTERPRETATION 17

ginning of this paper. The results are encouraging in 

that the seemingly contradictory results of previous 

investigation can be at least partly reconciled by the 

indicated variations in timing and intensity of arctic 

stream processes with bed-material size, channel pattern, drainage area, and climate. It is logical for investigators to have expected that permafrost would be a 

simplifying element in the study of arctic stream 

channels by possibly masking or dominating the effects 

of other variables. This expectation naturally led to 

contradictory generalizations on the basis of results 

from a single stream or from similar streams. The results from this study suggest that permafrost is best 

considered as no more than an additional variable, albeit an important one, in the study of arctic streams. 

Three additional and closely related aspects of erosion in arctic streams are (1) the effects of major floods, 

(2) long-term erosion rates, and (3) the hypothesized 

uniformity of arctic erosion rates.

EFFECTS OF MAJOR FLOODS

We do not have sufficient flow records or erosion 

measurements from the Alaskan Arctic to assess the 

relative effects of major floods versus annual events 

with any certainty.

No truly outstanding flood has occurred in the 

streams with gaging stations during the period of flow 

measurement (post-1970). The peak instantaneous discharge of record of 818 m3/s at the gaging station on the 

Sagavanirktok River was the result of a summer storm 

on August 19, 1974. That value compares with a 

maximum evident flood discharge of 1750 m3/s measured by Childers and Jones (1975, table 1) at a site on 

the Sagavanirktok River 7.2 km upstream from the 

Lupine River confluence and a short distance 

downstream from site S2. Childers, Sloan, and Meckel 

(1973, p. 4) noted that the major flooding observed on 

the Sagavanirktok River in July 1961 may have 

formed the high-water marks observed in 1972 at a 

level 2 to 3 feet above the tops of the main channel 

banks. This flooding was observed approximately 20 

km upstream from site S4, downstream from the confluence with the Ivishak River.

The unusual, but not catastrophic, 1976 summer 

storm in the Atigun River watershed was seen to have 

caused substantial channel change. The flow records 

(fig. 2) show that summer storms appear to be common 

in the Sagavanirktok River watershed, of which the 

Atigun River drainage is a part. Similar summer 

storms, however, have not significantly affected the recorded runoff in streams draining only foothill or 

coastal-plain areas. Not only the effects of the rarer 

floods, but their probable origin—rainstorm runoff or

breakup—apparently may vary with physiographic 

province.

The evidence from other parts of the Arctic is likewise sketchy. On Ellesmere Island in the Canadian 

Arctic, Cogley and McCann (1976, p. 109-110) reported 

extensive shifting caused by summer-storm runoff in a 

braided channel from precipitation having a recurrence interval of perhaps 10 to 20 years. The degree 

to which the course of the main channel was changed 

was not clear, however.

There is no reason to suppose that the larger arctic 

stream channels are different from those in temperate 

regions in which the channel form is attributed to 

events of moderate frequency rather than the rare, 

catastrophic event (Wolman and Miller, 1960, p. 67). 

This frequency concept is not equally applicable to all 

streams. The smaller the stream the more variable the 

flow and the rarer will be the flood that causes channel 

change by exceeding the threshold of sediment movement or, in the case of arctic streams, causing sufficient 

thawing to release sediment for transport.

An example of the latter case may be the small 

Happy Creek drainage that showed essentially no 

channel change in spite of the fact that the 1976 

breakup flooding was the greatest recorded. At the 

same time, the Sagavanirktok River at upstream sites 

showed significant amounts of change with levels of 

breakup flooding that were only moderate, based on 

the short period of record. That in the Arctic the events 

of moderate frequency—the bankfull floods of other 

climates—in fact correspond to the breakup floods has 

been a common suggestion. The results of this study 

confirm this suggestion for at least the larger streams.

LONG-TERM EROSION RATES

It is interesting to examine long-term erosion rates 

in light of the existing flow records. Brice (1971, p. 

36-39) compared aerial photographs representing the 

period 1949-69 for 9.1 km of the Sagavanirktok River 

in the braided channel between sites S3 and S4. The 

flooding seen downstream in July 1961 would have affected these reaches during that period in the event the 

flow did not originate in the drainage basin of the 

Ivishak River. The photographs reveal a maximum of 

91 m of lateral erosion, equivalent to 4.6 m/yr. Less 

than 12 percent of the main vegetated bank was affected by erosion detectable at the scale of the photographs, however, and Brice (1971, p. 38) was able to 

conclude that in general the 1949 river course was "not 

very different" from the 1969 course.

Childers and Jones (1975, p. 8) detected no erosion 

from aerial photographs at their site between October 

17, 1969, and August 24, 1974. A brief comparison by
18 EFFECTS OF PERMAFROST ON STREAM CHANNEL BEHAVIOR IN ARCTIC ALASKA

the writer of 1969 and 1974 photographs in a series of 

reaches upstream from the Lupine River in the sinuous 

and meandering section of the stream similarly showed 

that detectable erosion in that time span was rare. The 

maximum amount seen was 34 m, equivalent to 6.8 

m/yr and comparable to the maximum of 5.0 m of lateral erosion observed on the ground during the 1976 

breakup period.

Thus, over a 27-year period surprisingly little lateral 

erosion has been observed generally along the course of 

the Sagavanirktok River, although significant 

amounts have occurred locally. This same result was 

found in the Canadian Arctic, where few areas of rapid 

lateral erosion in the previous 16 to 20 years were seen 

by Lewis and McDonald (1973, p. 268) in aerial photographs of streams draining the north slope of the 

Yukon Territory. The Alaskan results suggest, but by 

no means establish, that most of the erosion has taken 

place in small increments during breakup flooding. 

What erosion has occurred apparently has been concentrated in specific reaches where texture or stratigraphy of the bank is favorable for thermo-erosional 

niching. Other reaches have remained relatively unchanged since 1969, and possibly since 1949.

The above maximum amounts of lateral erosion cannot be compared directly with those of streams in other 

climates because of the variability of those data (see 

Wolman and Leopold, 1957, p. 96 and table 4). Certainly, however, the long-term rates of change in the 

channels of the streams in the study area appear to be 

small in relation to those of unregulated stream channels in most nonpermafrost environments. Assuming 

that absolute rates of erosion are in fact generally less 

than those in comparable channels elsewhere, the difference may be due more fundamentally to the climatic 

and hydrologic character of the permafrost environment than to the presence of permafrost in channel 

banks.

HYPOTHESIZED UNIFORMITY OF EROSION RATES

This study generally lends validity to the idea that 

arctic stream erosion may be comparatively uniform on 

an annual basis (Gill, 1972, p. 130; Cooper and Hollingshead, 1973, p. 276). Factors discussed above that 

support this hypothesis include: (1) the demonstrated 

effect of at least partial permafrost control of breakup 

channel processes; (2) the probability that breakup 

flooding is the most important channel-forming process 

in most larger watersheds; (3) what little is known of 

long-term erosion rates in the study streams; and (4) 

the similarity of those long-term rates with the maximum erosion observed during the breakup in 1976.

Uniformity in annual rates could occur as a result of

comparable amounts of thermo-erosional niching and 

slumping in response to a relatively uniform annual 

thermal input to banks during breakup floods. The 

concentration of about half the annual precipitation in 

the breakup runoff of most streams is a phenomenon 

tending to smooth the vagaries of individual storms. 

Also, the variability of the precipitation in arctic icefree barrens is reportedly smaller than that of dry 

coastal areas elsewhere (Giovinetto, 1974, p. 20).

Further work could establish that an average erosion rate determined from aerial photographs may represent a rate that could more confidently be extrapolated than the rates observed in nonpermafrost areas. 

There are a variety of obvious factors, however, that 

will mitigate against such extrapolation: susceptibility 

to thermal niching may vary laterally with the texture 

and ice content of the flood-plain deposits; icings and, to 

a lesser extent, ice jams may divert flow and cause 

large amounts of local erosion; and normal channel 

shifts will change the loci of maximum erosion.

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