the evolution of a pulsating supraglacial stream
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University of Calgary
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Graduate Studies The Vault: Electronic Theses and Dissertations
2015-09-17
The Evolution of a Pulsating Supraglacial Stream
St. Germain, Sarah
St. Germain, S. (2015). The Evolution of a Pulsating Supraglacial Stream (Unpublished master's
thesis). University of Calgary, Calgary, AB. doi:10.11575/PRISM/28140
http://hdl.handle.net/11023/2460
master thesis
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UNIVERSITY OF CALGARY
The Evolution of a Pulsating Supraglacial Stream
by
Sarah St. Germain
A THESIS
SUBMITTED TO THE FACULTY OF GRADUATE STUDIES
IN PARTIAL FULFILMENT OF THE REQUIREMENTS FOR THE
DEGREE OF MASTER OF SCIENCE
GRADUATE PROGRAM IN GEOGRAPHY
CALGARY, ALBERTA
SEPTEMBER, 2015
© Sarah St. Germain 2015
ii
Abstract
Supraglacial streams are a significant part of the glacial hydrological regime and
important for understanding the dynamics between glacial hydrology, glacier dynamics,
and climate change. During the 2014 ablation period on Bylot Island, Nunavut, a
supraglacial stream, which flowed over a 13 m high waterfall at the front of Fountain
Glacier, began to pulsate. The pulsating phenomenon involved the stoppage of flow over
the waterfall for 10-15 s, with a total period of 27 s. The objective of this research was to
determine the factors that influenced the evolution of the supraglacial stream. Results
show a change in the weather, where multiple rainfall events occurred during the second
week of the study. Analysis suggests that the rainfall caused the formation of a step-pool
sequence within the streambed. In conclusion, the formation of the step-pool sequence
and constructive interference from changes in flow conditions caused the distinctive
pulsating.
iii
Acknowledgements
I would like to express my gratitude to a number of people for their support,
encouragement, and inspiration throughout my master’s degree.
First and foremost I would like to thank Dr. Brian Moorman. His guidance and
advice over the last two years has been invaluable. For the countless hours spent
enthusiastically listening to me babble on about my data, and teaching me more in an
hour long meeting than I could learn in a month on my own, I thank you! Remarkably, I
feel this learning process has been the perfect balance between figuring out things for
yourself and being steered in the right direction when I needed it. His ability to know
when a student needs assistance and always make time for them is what makes him a
great supervisor. Also, I’m not sure how many graduate students can claim that they have
fun during thesis meetings, but I can! I honestly could not have asked for a better
supervisor.
I must thank the professors, instructors, and lab techs that have helped me along
my journey. Firstly, I would not be where I am today without the countless reference
letters and support from Dan Patterson. He taught me everything I know about GIS and
made my undergrad a truly enjoyable experience. I wish to thank the members of my
proposal committee Chris Hugenholtz and Darren Sjogren for their feedback and
constructive criticisms. I would also like to thank Derek Wilson for his assistance in
prepping equipment and teaching me to use many of the instruments required to my
conduct research.
Many thanks to Michelle Blade as “we were in this together”. The many life
conversations and the time we spent in the field on Bylot Island was a meaningful part of
iv
my last two years. Further, my officemates Elena, Ellie, Alison, and Mari have been
amazingly encouraging and have given me great advice on the thesis procedure, arctic
research, and publishing academic papers.
To the British researchers that joined our Bylot Island expedition team, many
thanks must be expressed. Thanks to Mike Hambrey for sharing some of his infinite
glaciology (especially in the area of structural glaciology), Martin Smart for his
assistance during fieldwork and on the journey back to camp, Tristram Irvine-Fynn for
his help in creating a new thesis topic in the matter of hours, and Richard Waller for
being the most cheerful and upbeat person in camp despite the rain.
I would also like to thank professor Gilles Gauthier and Marie-Christine Cadieux
from Laval University. These faithful Bylot Island researchers have been collecting
precipitation data since 1994 and have graciously allowed me to use their data.
I would also like to express my gratitude towards my family and friends.
Specifically, I would like to thank my parents for facilitating my interest in the arctic, my
mother for editing every important academic paper I’ve written, my dad for taking me on
countless camping trips, and my three brothers for all the outdoor adventures. I would
also like to thank my friend Jacob, and Aunt Sue and family for being my inspiration for
exercise, health, and continued education.
Lastly, this research would not have been possible without the generous financial
and logistical support from: Natural Sciences and Engineering Research Council of
Canada (NSERC), Polar Continental Shelf Project (PCSP), Parks Canada, Northern
Scientific Training Program (NSTP), Arctic Institute of North America (AINA),
University of Calgary, Department of Geography, and the Hamlet of Pond Inlet.
v
Table of Contents
Abstract .............................................................................................................................. ii
Acknowledgements .......................................................................................................... iii
Table of Contents ...............................................................................................................v
List of Tables .................................................................................................................... ix
List of Figures and Illustrations .......................................................................................x
List of Symbols, Abbreviations and Nomenclature...........................................................xv
CHAPTER ONE: INTRODUCTION ..............................................................................1
1.1 Context ..........................................................................................................................1
1.2 Objectives......................................................................................................................5
1.3 Hypothesis .....................................................................................................................5
1.4 Outline ...........................................................................................................................6
LITERATURE REVIEW .................................................................7
2.1 Energy Balance.............................................................................................................7
2.1.1 Net Radiation .........................................................................................................8
2.1.2 Sensible and Latent Heat Transfers ....................................................................10
2.1.3 Turbulent Heat Fluxes ........................................................................................12
2.2 Thermal Regime .........................................................................................................13
2.3 Glacial Hydrology ......................................................................................................14
2.4 Supraglacial Streams .................................................................................................15
2.4.1 Development and Evolution ................................................................................16
2.4.2 Stream Characteristics .........................................................................................17
2.4.3 Stream Morphology .............................................................................................19
STUDY AREA .............................................................................21
3.1 Bylot Island, Nunavut ................................................................................................21
3.2 Fountain Glacier ........................................................................................................25
3.3 Central Supraglacial Stream ....................................................................................28
METHODS .....................................................................................31
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4.1 Ablation .......................................................................................................................31
4.2 Albedo .........................................................................................................................32
4.3 Surface Roughness .....................................................................................................33
4.4 Streambed Profiles .....................................................................................................34
4.5 Supraglacial Stream Discharge ................................................................................34
4.5.1 Velocity-Area Method ..........................................................................................35
4.5.2 Depth Measurements ...........................................................................................35
4.5.3 Time-Lapse Imagery ............................................................................................37
4.6 Positional and Topographic Information ................................................................38
4.7 Meteorological Data ...................................................................................................39
CHAPTER FIVE: RESULTS .........................................................................................41
5.1 Meteorological Data ...................................................................................................41
5.1.1 Air Temperature ...................................................................................................41
5.1.2 Net Radiation .......................................................................................................41
5.1.3 Relative Humidity .................................................................................................42
5.1.4 Water Balance ......................................................................................................43
5.2 Glacier Surface Characteristics ................................................................................46
5.2.1 Structure ...............................................................................................................46
5.2.2 Ablation ................................................................................................................47
5.2.3 Relative Albedo .....................................................................................................51
5.2.4 Surface Roughness ..............................................................................................59
5.3 Supraglacial Stream Characteristics........................................................................62
5.3.1 Watershed / Stream Diamensions .......................................................................62
5.3.2 Stream Meandering .............................................................................................63
5.3.3 Streambed Erosion ...............................................................................................64
5.3.4 Sinuosity ...............................................................................................................66
5.3.5 Slope .....................................................................................................................67
5.3.6 Step-pool Sequence ..............................................................................................68
5.3.7 Stream Temperature ............................................................................................70
5.3.8 Discharge ..............................................................................................................71
5.3.9 Froude number ....................................................................................................74
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5.3.10 Reynolds number ...............................................................................................75
5.3.11 Pulsating .............................................................................................................76
CHAPTER SIX: ANALYSIS ..........................................................................................82
6.1 Changes to the Glacier Surface Characteristics .....................................................82
6.1.1 Ablation ................................................................................................................82
6.1.2 Albedo ...................................................................................................................84
6.1.3 Roughness ............................................................................................................85
6.1.4 Changes to the Surface Characteristics and Effects on Supraglacial Stream ..86
6.2 Evolution of Stream Characteristics ........................................................................87
6.2.1 Stream Temperature ............................................................................................87
6.2.2 Streambed Erosion ...............................................................................................89
6.2.3 Discharge ..............................................................................................................91
6.2.4 Step-pool Sequence ..............................................................................................92
6.2.5 Pulsating ...............................................................................................................93
6.2.6 Summary of Changes and Stream Characteristic Connections .........................95
CHAPTER SEVEN: DISCUSSION ...............................................................................97
7.1 Surface Characteristic and Stream Characteristic Connections...........................97
7.1.1 Watershed Runoff Contributions ........................................................................97
7.1.2 Stream Incision ....................................................................................................99
7.2 Development of Step-pool Sequence and Pulsating Phenomenon .......................101
7.2.1 Formation of the Step-pools ..............................................................................102
7.2.2 Pulsating Phenomena ........................................................................................107
7.3 Discussion of Uncertainty ........................................................................................112
7.3.1 Meteorology ........................................................................................................112
7.3.2 Ablation ..............................................................................................................113
7.3.3 Albedo .................................................................................................................114
7.3.4 Roughness ..........................................................................................................116
7.3.5 Stream Temperature ..........................................................................................117
7.3.6 Discharge ............................................................................................................118
7.3.7 Stream position/Step-pools/Streambed Erosion................................................119
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7.3.8 Pulsating .............................................................................................................120
CHAPTER EIGHT: CONCLUSION ..........................................................................121
8.1 Summary and Implications .....................................................................................121
8.2 Suggestions for Future Work ..................................................................................122
References………………………………………………………………………...……………………125
ix
List of Tables
Table 5.1. Evaporation, rainfall, and net water balance amounts for the Reach 2
watershed from July 7-23, 2014. ............................................................................... 43
Table 5.2. Measured ablation, average daily ablation, and amount of water
equivalent within the Reach 2 watershed from July 8-22. ........................................ 48
Table 5.3. Relative albedo computed from various sources and the estimated daily
relative albedo from July 7-23. ................................................................................. 58
Table 5.4. Measured stream erosion from the RTK survey and at the bottom depth
sensor, used for a combined total streambed erosion per day and amount of
water equivalent along the stream from July 8-24. .................................................. 65
Table 5.5. Step-pool characteristics including step height and pool depth on July 13
and July 22-23, and water depth on July 23. ............................................................ 69
Table 5.6. Froude number calculated for the top of Reach 2 from July 8-22. ................. 75
Table 5.7. Reynolds number calculated for the top of Reach 2 from July 8-22................ 76
Table 7.1. Comparison between relative albedo in various directions for July 16, 22,
and 23. ..................................................................................................................... 115
x
List of Figures and Illustrations
Figure 2.1. Glacier energy balance. ................................................................................... 8
Figure 2.2. Schematic representation of a thermal profile of a temperate and
polythermal glacier. Note that melt will occur at 0oC or PMP. ............................... 14
Figure 3.1. Study Area; A) Location map of Bylot Island within Canada; B) Close-up
map of Bylot Island illustrating the surrounding area; C) Landsat 7 image of
Fountain Glacier. ...................................................................................................... 24
Figure 3.2. Temporal variation of Fountain Glacier's terminus. The retreat rate has
increased considerably since 1982 when a collapse feature on the northern half
of the snout was first observed (from Wainstien, 2011). ........................................... 26
Figure 3.3. A) Changes in ice thickness measured from July 1, 2010 to July 2, 2011.
Increases in thickness to the east of the terminus reflect changes to the
proglacial icing; B) Horizontal flow speed and flow direction between July 1,
2010 and July 2, 2011 (from Whitehead et al., 2013). .............................................. 27
Figure 3.4. The stream studied was located near the terminus of Fountain Glacier.
Arrows denote the 4 reaches. Key structural elements and stream characteristics
are also shown. ......................................................................................................... 29
Figure 3.5. A) Previous stream locations in 2010, 2011, and 2014 overlaid on an
orthoimage from 2011; B) Close-up map of the bottom section of the stream and
glacier ice flow amounts between July 2, 2011 and July 24, 2014. .......................... 30
Figure 4.1. Ablation stake locations within the Reach 2 watershed. ................................ 32
Figure 4.2. Horizontal reference used for the collection of micro-topographic data
for calculation of the aerodynamic roughness length. .............................................. 33
Figure 4.3. Depth sensor installation within the streambed. Note for the purpose of
this image the security string was recoloured black, but in reality it was white so
as not to impact melt of the glacier. .......................................................................... 36
Figure 4.4. Location of the depth sensors, evaporation pan, and roughness survey
within the Reach 2 watershed. .................................................................................. 37
Figure 4.5. Time-lapse camera placed ~75 m from the base of the waterfall used to
collect pulsating frequency. Note that the images were taken one minute apart
on July 23, 2014. A) Shows the waterfall at 8:15 with high discharge; B)
Displays the waterfall at 8:16 with no discharge. .................................................... 38
Figure 4.6. Fabricated evaporation pan on the surface of Fountain Glacier. ................. 41
xi
Figure 5.1. Air Temperature recorded at Bylot-3 AWS from July 7-22, 2014. ................ 41
Figure 5.2. Net radiation recorded at Bylot-3 AWS from July 7-22, 2014....................... 42
Figure 5.3. Relative humidity recorded at Bylot-3 AWS from July 7-22, 2014. ............... 42
Figure 5.4. Net water balance and rainfall events for the Reach 2 watershed from
July 7-22, 2014. ......................................................................................................... 46
Figure 5.5. Image of terminus of Fountain Glacier from 2014 showing structural
elements. .................................................................................................................... 47
Figure 5.6. Ablation amounts at each of the 24 sites for July 9-10, 16-17, and 21-22. ... 49
Figure 5.7. Images of the sites that had the highest and lowest ablation on July 9, 16,
and 22. ....................................................................................................................... 49
Figure 5.8. Total ablation at each site from July 8-22. .................................................... 50
Figure 5.9. Relative albedo amounts at 30 sites on July 16, 22, and 23. ......................... 53
Figure 5.10. Variations in relative albedo in 4 directions on July 16, 22, and 23,
2014. .......................................................................................................................... 54
Figure 5.11. Relative albedo computed from images taken at the 24 ablation sites on
July 7, 8, and 9. ......................................................................................................... 55
Figure 5.12. Relative albedo computed from images taken at the 24 ablation sites on
July 16, 22, and 23. ................................................................................................... 55
Figure 5.13. Images used to obtain the relative albedo at site 3 on July 7, 9, 16, and
23. .............................................................................................................................. 56
Figure 5.14. Map of the relative albedo at each of the ablation sites on July 16, 22,
and 23. ....................................................................................................................... 57
Figure 5.15. Glacier surface roughness measurements taken on July 22 and 23. ........... 60
Figure 5.16. Micro-scale surface roughness estimated from images taken at the 24
ablation sites on July 9, 16, and 22. ......................................................................... 61
Figure 5.17. Roughness estimated from images take at ablation stakes 2, 7 and 8 on
July 9, 16, and 22. ..................................................................................................... 62
Figure 5.18. Schematic diagram of the watershed with stream cross-sections at the
top and bottom of the Reach 2. ................................................................................. 63
Figure 5.19. Plan-view of Reach 2 illustrating meandering and slight changes in
stream position on July 13 and 22. ........................................................................... 64
xii
Figure 5.20. Stream sinuosity. Arrows denote the 4 reaches. .......................................... 66
Figure 5.21. Stream slope. Note numbers and arrows denote stream reaches. ............... 67
Figure 5.22 Illustration of the stream features identified using the survey points ........... 68
Figure 5.23. Location of step-pools within the streambed on July 13 and July 22-24.
Note that 5 step-pools (1-5) existed in the streambed on July 13 and 26 (6-31)
step-pools were present on July 22-24. Note offset of vertical axes for separation
and clarity. ................................................................................................................ 70
Figure 5.24. Stream temperature acquired from the top depth sensor from July 8-18.
Four manual measurements were taken during the study with the velocity meter. .. 71
Figure 5.25. Uncalibrated depth obtained from the top depth sensor from July 8-21.
Note the depth sensor was not in the stream from July 18, 16:00-July 19, 15:30. ... 72
Figure 5.26. Calibrated depth from the top depth sensor from July 8-21 and
corresponding manual measurements. ..................................................................... 72
Figure 5.27. Relationship between measured discharge and measured depth................. 73
Figure 5.28. Discharge at the top depth sensor from July 8-21. ...................................... 74
Figure 5.29. Manual discharge taken at bottom of Reach 2 on July 16 at 17:30. ........... 77
Figure 5.30. Manual discharge taken at the top of Reach 2 on July 16 at 17:48. ........... 78
Figure 5.31. Range in discharge estimated of the waterfall from a video taken on July
18 at 17:05. ............................................................................................................... 78
Figure 5.32. Manual discharge taken at bottom of Reach 2 on July 22 at 10:35. ........... 79
Figure 5.33. Discharge determined from a video taken between Reach 3 and 4 on
July 22 at 10:40. ....................................................................................................... 80
Figure 5.34. Discharge estimated at the waterfall from a video taken July 22 at
11:20. ........................................................................................................................ 80
Figure 5.35. Pulsating occurrence from July 7-22. .......................................................... 81
Figure 6.1. Relationship between ablation and net radiation. ......................................... 83
Figure 6.2. Relationship between ablation and air temperature. ..................................... 83
Figure 6.3. Relationship between ablation and relative humidity. ................................... 83
Figure 6.4. Relationship between relative albedo and relative humidity. ........................ 84
xiii
Figure 6.5. Depiction of glacier surface roughness change from rainfall. ...................... 85
Figure 6.6. Relationship between net radiation and relative humidity. ........................... 86
Figure 6.7. Relationship between stream temperature and net radiation. ....................... 88
Figure 6.8. Relationship between stream temperature and air temperature. ................... 88
Figure 6.9. Relationship between stream temperature and relative humidity. ................. 89
Figure 6.10. Relationship between streambed erosion and net radiation. ....................... 90
Figure 6.11. Relationship between streambed erosion and relative humidity. ................ 91
Figure 6.12. Relationship between discharge and net radiation. ..................................... 92
Figure 6.13. Relationship between discharge and relative humidity. .............................. 92
Figure 6.14. Relationship between pulsating events and relative humidity. .................... 93
Figure 6.15. Relationship between pulsating event and net radiation. ............................ 94
Figure 6.16. Relationship between pulsating event and discharge. ................................. 94
Figure 7.1. Influences of Reach 2 watershed stream runoff contributions. ...................... 98
Figure 7.2. Comparison of Reach 2 watershed stream runoff and stream discharge. ..... 98
Figure 7.3. Stream incision. .............................................................................................. 99
Figure 7.4. Cross section of stream watershed illustrating the streambed incision
before the rainfall (July 9-13) and after multiple rainfall events (July 13-22)....... 100
Figure 7.5. Images of the stream incision all within 3 m of the bottom depth sensor;
where A) shows the stream fairly level with the bank; B) shows a slight incision
of the stream; and C) shows significant incision as the water level is much below
the surface. .............................................................................................................. 101
Figure 7.6. Helical flow step-pool formation; where A) is the beginning of a pulse
with helical flow occurring around the stream bend; B) was taken a few seconds
later during the middle of a pulse. Increased local discharge can be seen and the
water crashes down directly below the stream bend; C) was again taken a few
seconds later, in between pulses events. It can be seen where the helical flow
created the pool and subsequent lip. Images were taken on July 22 at 10:40 at
the major stream bend. ........................................................................................... 103
Figure 7.7. Step-pool locations on aerial view of the stream showing the correlation
of pools with stream meander bends and transverse fractures. .............................. 104
xiv
Figure 7.8. Image of a hydrological event on taken on July 20 where A) shows two
locations that sediment-laden water was traveling from the base of the glacier
through the transverse fractures onto the surface; and B) shows is close up
image of the water flowing from a transverse fracture. .......................................... 105
Figure 7.9. Structural step-pool formation from glacial slip events. ............................. 106
Figure 7.10. Transverse fracture formed step-pool. Image was taken on July 22
within Reach 2. ........................................................................................................ 106
Figure 7.11. Pulsating phenomena where A) shows the normal stream conditions; B-
D) shows the water sloshing back and forth within the step-pool; E) shows the
constructive interference; and F) shows the traveling waves. ................................ 109
Figure 7.12. Depiction of the wave as the amplitude and period increased
downstream. ............................................................................................................ 110
Figure 7.13. Net water balance with 0.5 mm of uncertainty for negative values
(evaporation) and 1 mm of uncertainty for positive values (rainfall). ................... 113
Figure 7.14. Albedo values under differing cloud conditions taken during the July 23
albedo survey a few min apart. ............................................................................... 115
Figure 7.15. Relative albedo uncertainty. ...................................................................... 116
Figure 7.16. Roughness measurements for July 22 with 2.5 mm uncertainty. ............... 117
Figure 7.17. Stream temperature with manual point measurements and 0.05oC error
bars. ........................................................................................................................ 118
Figure 7.18. Discharge with uncertainty shown as 7x10-4 m3/s. .................................... 119
xv
List of Symbols, Abbreviations and Nomenclature
Symbol Definition
A Cross-sectional Area
ADA Average Daily Ablation
ALrel Relative Albedo
Aw Watershed Area
AWS Automatic Weather Station
CL Channel Length
CTZ Cold‐temperate Ice Transition Zone
D0 Stream Depth
DD Downvalley Distance
DEM Digital Elevation Model
E’ac Evaporation on Snow Surface
e0 Saturation Pressure of Melting Ice
ea1 Vapour Pressure
ELA Equilibrium Line Altitude
esa Saturation Vapour Pressure
f Frequency
Fr Froude Number
g Gravity
GCP Ground Control Points
GPS Global Positioning System
h* Effective Height for the Roughness Elements
xvi
IDW Inverse Distance Weighted
Imax Maximum Digital Number
Imin Minimum Digital Number
IPCC Intergovernmental Panel on Climate Change
LWR Longwave Radiation
MA Measured Ablation
MAAT Mean Annual Air Temperature
MASL Meters Above Sea Level
MLS Minimum Linear Shift
MS Mean Stretch
Omax Overall Maximum Digital Number
Omin Overall Minimum Digital Number
P Wetted Perimeter
pi Ice Density
PMP Pressure Melting Point
pw Water Density
Qe Latent Heat Transfer
Qh Sensible Heat Transfer
Qm Energy available to Melt Ice
Qn Net Radiation
Qr Sensible Heat Flux supplied by Rain
r Pearson’s Correlation Coefficient
Re Reynolds Number
xvii
REWE Reach Erosion Water Equivalent
Rh Relative Humidity
Rh Hydraulic Radius
RTK Real Time Kinematic
s Silhouette Area of Roughness Elements
S Frequency per Unit Area
SE Stream Erosion
Si Sinuosity Index
SI Stream Incision
SWR Shortwave Radiation
Ta Air Temperature
U Velocity
u1 Wind Speed
v Kinematic Viscosity
WE Water Equivalent
WWE Watershed Water Equivalent
X Width of a Typical Element
zo Aerodynamic Roughness Length
1
CHAPTER ONE: INTRODUCTION
1.1 Context
According to the Intergovernmental Panel on Climate Change (IPCC), glaciers
provide one of the most visible indications of the effects of climate change (Lemke &
Ren, 2007). With the global climate models predicting a rise in air temperature, glacier
melt rates are expected to increase. While the exact effects of glacial melt on sea-level
remains somewhat unknown (Gardner et al., 2011), global sea level rise is expected to
have enormous environmental, social, and economic implications on all of society
(McCathy et al., 2011). Therefore, developing a comprehensive understanding of the link
between glacier dynamics and climate change is imperative.
Not only are air temperatures expected to rise from global warming, but
precipitation is very likely to increase at high latitudes (Lemke & Ren, 2007). A warmer
atmosphere means that more energy is available to drive circulation leading to greater
precipitation. Approximately 1oC of warming will increase precipitation by 5% (Lemke
& Ren, 2007). While it is understood that increased warming will cause a greater melt
rate, the impact of increased precipitation on high arctic glacier dynamics, as well as
glacier hydrology is not well understood. There are complex interactions between the
water and ice in glaciers, and our knowledge of this process is far from complete (Sharp
et al., 1998). Before the effects of climate change on glacier dynamics can be established,
a stronger foundation in glacial hydrology is essential to further our comprehension.
Changes in glacier dynamics and hydrology will not only impact society as a
whole, but also have regional effects. Understanding glacier hydrology is imperative for
water resource management, natural disaster planning, and environmental quality.
2
Alterations in the amount of glacier runoff can affect hydroelectric power, reservoir
operations, and irrigation (Hambrey & Alean, 2004). Flood predictions may also become
increasingly difficult due to variations within the channel. Changes in glacier runoff may
also impact the amount and location of gravel deposits traditionally used as a natural
resource (Anderson & McDonnell, 2005). Not only is it probable that the amount of
water will change, but also the water quality. With greater discharge, the rate of fluvial
erosion increases, and additional sediment downstream can negatively impact aquatic
environments and a wide range of species (Moody et al., 2003).
Supraglacial streams are meltwater channels that exist in the ablation zone of
glaciers during the summer season (Karlstrom et al., 2013). These streams are a
significant part of the glacial hydrological regime as they can account for a large portion
of the glacier runoff. On Bylot Island, Nunavut, during the peak ablation period of the
2014 season, an in-depth study of a small supraglacial stream on the surface of Fountain
Glacier was conducted. Methodology included measuring relative albedo, ablation, and
roughness to obtain glacier surface characteristics. Data collected in relation to stream
characteristics included: discharge, temperature, flow characteristics, and streambed
profiles. In addition, meteorological data was acquired from a number of sources.
In order for a supraglacial stream to form and exist, the streambed erosion must
be greater than the rate of the surrounding glacier ablation. The difference between the
streambed erosion and glacier ablation will also dictate the amount the stream is able to
incise into the surface of the glacier (Marston, 1983). There are a number of complex
interactions between weather parameters, the glacier surface characteristics, and the
stream hydrological characteristics that impact the development and evolution of a
3
supraglacial stream. Air temperature and net radiation impact the amount of ablation
(Dozier, 1974), and rainfall increases the amount of streambed erosion (Karlstrom et al.,
2013). Surface characteristics such as albedo, aspect, and topography influence the
amount and location of meltwater, while the surface roughness impacts the amount of
time it takes water to travel into the stream (Knighton, 1972). Important stream
characteristics include: discharge, stream temperature, flow conditions, sinuosity, and
streambed morphology.
During the second week of the two week study period, there was a large variation
in a number of the weather parameters. In particular, multiple successive rain events
occurred during the second week. According to Karlstrom et al. (2013), the development
of supraglacial streams is sometimes augmented by rainwater. However, the exact
impacts have rarely been quantified. Importantly, a feedback exists where, as the stream
incises deeper into the glacier surface, the stream also propagates farther up-glacier. This
increases the watershed area, and further facilitates the growth of the stream (Gabler et
al., 1999). As supraglacial streams are a significant part of the glacial hydrological
regime, their development and evolution is important for understanding the connections
between glacial hydrology, glacier dynamics, and climate change.
In addition to the variation in weather parameters, changes to the glacier surface
characteristics and stream characteristics were observed. The glacier surface appeared
cleaner, having less fine sediment and cryoconite material, and became extremely smooth
as the cryoconite holes were eroded away. Within the studied supraglacial stream, a
number of step-pools suddenly formed. A step-pool sequence is stair-like in appearance;
where an erosional pool is formed directly underneath a step from the added energy in the
4
waterfall and a promontory (or lip) forms directly downstream from the pool due to the
dispersion of energy within the pool. The lip forms a reverse bed slope within the stream
profile, and thus forces water to travel against gravity to travel farther downstream (Vatne
& Refsnes, 2003).
Step-pools have been observed in mountainous, bedrock, arid, and supraglacial
streams, and are a common element of fluvial environments (Knighton, 1998). Although
these formations occur in a number of environments, the mechanics allowing them to
form are quite different. Step-pools in alluvial rivers develop during extreme floods.
Vatne and Refsnes (2003) assume that step-pools also form in glacier meltwater streams
during high discharge. However, they suggest that the cause is more from the added
energy in the form of heat from frictional dissipation versus the additional water. A few
studies have been undertaken regarding step-pool formation within englacial flowpaths,
however it has rarely been investigated in supraglacial streams on polythermal glaciers
(Irvine-Fynn et al., 2011).
In succession to the changes that occurred within the streambed, the supraglacial
stream (which flowed over a 13 m high waterfall at the front of the glacier) began to
pulsate. The pulsating phenomenon involved the complete stoppage (or extreme sudden
reduction) of discharge from the waterfall and the subsequent restart of flow.
Pulsating flow within supraglacial streams has only been noted in literature on a
few occasions. As the occurrence of pulsating flow is rare and sporadic in nature, it is
challenging to study this unique phenomenon. To date the literature has focused on
characterizing the flow pattern and is lacking completeness. The majority of the literature
5
fails to explain the mechanisms leading to the pulsating and the processes involved in this
phenomenon.
As the weather parameters, glacier surface characteristics, and stream
characteristics altered drastically during the study period, the research aim of this project
is to understand the factors that influenced the evolution of the small supraglacial stream
located on Fountain Glacier.
1.2 Objectives
In order to achieve a comprehensive understanding of the factors that influence
the evolution of the supraglacial stream, the following minor objectives have been
identified:
1) Establish the glacier surface characteristics and changes that occurred.
2) Determine the stream characteristics and evolution.
3) Examine the stream incision rate and subsequent impacts under various weather
conditions.
4) Understand the factors that influenced the rapid streambed evolution and the
processes that caused the pulsating phenomenon.
1.3 Hypothesis
Multiple successive rainfall events, and subsequent cloudy conditions, impacted
the glacier surface characteristics and stream characteristics. This in turn caused an
evolution in the streambed morphology and the pulsating flow to occur.
6
1.4 Outline
This thesis is comprised of a total of seven chapters. Chapter 1 is the introduction,
which provides context, objectives, an overall hypothesis, as well as this outline. Chapter
2 presents a literature review on the glacier energy balance, thermal regime, glacier
hydrology, and supraglacial streams. The study area is described in Chapter 3, while
Chapter 4 documents the methodology used to complete the objectives of this research.
Results, including meteorological data, glacier surface characteristics, and stream
characteristics are presented in Chapter 5. The analysis (Chapter 6), examines the factors
that affect the individual changes to the glacier surface characteristics and the evolution
of stream characteristics. Chapter 7, the discussion, is comprised of three major sections
including: surface characteristic and stream characteristic connections, development of
step-pool sequence and pulsating phenomenon, and discussion of error. Lastly, Chapter 8
includes a summary/implications section, as well as a suggestions for future work section.
7
LITERATURE REVIEW
This literature review will include sections on glacier energy balance, thermal
regime, glacial hydrology, and supraglacial streams. The combined understanding of each
of these sections is important for this study. Background information on the surface
energy balance explains the components that impact glaciers in general, the watershed of
supraglacial streams, and in turn the amount of runoff. The thermal regime section
provides information on the development of the surface layer and the glacial hydrology
section describes how this layer influences the overall hydrology and evolution of
supraglacial streams.
2.1 Energy Balance
The glacier energy balance involves a number of complex components and may
be expressed as:
𝑄𝑚 = 𝑄𝑛 + 𝑄ℎ + 𝑄𝑟 + 𝑄𝑒 [Eq. 2.1]
where 𝑄𝑚 is energy available to melt ice, 𝑄𝑛 is net radiation, 𝑄ℎ is sensible heat transfer,
𝑄𝑟 is the sensible heat flux supplied by rain, and 𝑄𝑒 is latent heat transfer (Benn & Evans,
1998; Hock, 2005).
Figure 2.1 illustrates the five main components of the glacier energy balance.
Within net radiation, there are two wavelength bands which are referred to as shortwave
radiation (SWR) and longwave radiation (LWR). Firstly, solar radiation is the primary
source of energy for the world’s climate system and travels to earth as SWR. LWR is the
infrared radiation that originates from the Earth’s surface (upward LWR) and the
atmosphere (downward LWR). Sensible heat is the exchange of thermal energy between
the glacier surface and the atmosphere. In terms of small scale studies, the sensible heat
8
supplied by rain can also be an important contributor to the overall energy balance.
Lastly, latent heat can be transferred to the glacier surface by condensation and from the
glacier surface by evaporation.
Figure 2.1. Glacier energy balance.
2.1.1 Net Radiation
Shortwave radiation is the main controlling factor when it comes to the source of
melt energy on the surface of glaciers (Pellicciotti et al., 2005). SWR can reach the
surface as direct sunlight or diffuse radiation, and influences the amount of melt and
evaporation at the glacier surface.
A portion of the SWR is reflected off the surface of the earth and this portion is
known as the broadband surface albedo or reflectivity. The albedo of the surface of a
glacier can vary drastically throughout the year and spatially over the glacier surface. It
can be as low as 0.1 for debris-covered ice to 0.9 for fresh dry snow. Factors affecting
this process include: grain size of the ice or snow, water located on the surface, presence
of impurities, surface roughness, crystal orientation, and glacial structure (Hock, 2005).
Albedo is extremely difficult to measure as it changes with the time of day, angle of
downward radiation, and cloudy conditions (Hock, 2005; Cuffey & Paterson, 2010).
9
Clear-skies cause anisotropic diffuse radiation; however, cloudy conditions cause
isotropic diffuse radiation. As a result of the radiation being uniformly reflected in all
orientations by cloudy conditions the measured albedo value can appear to be higher
(Hock, 2005).
Ablation is the term used to describe the mass losses of the glacier in the form of
runoff, evaporation, sublimation, calving, or avalanching (Glasser & Hubbard, 2005).
However, the term ablation is most often used to explain the amount of ice or snow melt
from the surface of the glacier, usually expressed in water equivalence (Konzelmann &
Braithwaite, 1995). Ablation is an important aspect of the surface melt and is dependent
on air temperature, net radiation flux, and rainfall. Rainfall is an important factor as it
adds heat from the atmosphere and promotes larger amounts of runoff (Benn & Evans,
1998); however, there tends to be less ablation on cloudy and rainy days (Dozier, 1974).
The albedo directly affects the ablation of the glacier surface. In fact, small scale spatial
changes in albedo can cause varying rates of ablation (Hock, 2005). The difference
between water and ice on the surface of the glacier changes the albedo and often creates a
feedback loop for continued melt. As water has a lower albedo, this will facilitate melt,
causing more water on the surface, further lowering the albedo.
A unique characteristic of non-temperate glaciers in the arctic is the development
of a weathering crust. By definition, the weathering crust is a porous layer of ice with
loosely interlocking crystals (Müller & Keeler, 1969). This results from preferential melt
along grain boundaries. Development of the weathering curst occurs after the melt of the
snow cover and also directly affects the albedo and rate of ablation.
10
The weathering crust is an essential part of the glacier surface as it can enable
transient storage of water, delay drainage of seasonal runoff, and can restrict water from
percolating to the basal regions (Larson, 1977; Irvine-Fynn et al., 2011). The weathering
crust is caused by cryoconite formation (Müller & Keeler, 1969). Cryoconite holes are
formed by the aeolian deposit of dust upon glacier surfaces (Edwards et al., 2013). The
dust is a “biologically active aggregate of microorganisms, mineral particles and organic
matter” termed a cryoconite granule. (Langford et al., 2014). The cryoconite granule
cause localized melting and form cylindrical melt-water/dust filled holes due to the low
albedo of the dark organic matter (Edwards et al., 2013).
Longwave radiation is another important component in the energy balance of
glaciers. Snow, ice, and water are near perfect emitters, meaning that there can be a
significant loss of energy at the surface. Some of the downward LWR escapes to space;
however, a portion is absorbed or/and emitted from clouds, water vapor, and carbon
dioxide. Ozone, methane, and other greenhouse gases in the lower atmosphere cause a
substantial amount of the longwave radiation to be reradiated and the energy can return to
the Earth’s surface (Cuffey & Patterson, 2010). Importantly, weather parameters have a
strong influence on the energy loss or gain. Cloudy and humid conditions increase LWR
in the lower atmosphere and cause continued heating of the glacier surface; whereas
clear, dry conditions cause the surface to cool (Benn & Evans, 1998).
2.1.2 Sensible and Latent Heat Transfers
Sensible heat transfer occurs when the air temperature is warmer than the glacier
surface; a temperature gradient is created and promotes vertical air exchange (Benn &
11
Evans, 1998). Heat exchange increases when there are strong winds and turbulence from
a rough glacier surface (Benn & Evans, 1998; Cuffey & Paterson, 2010).
The sensible heat flux of rain is often neglected in an overall surface energy
balance of a glacier. However, the sensible heat flux may be an important factor in short-
term studies. Under normal daily conditions there is an interaction between the glacier ice
surface and the air directly above the surface. During the summer when the air
temperature is greater than 0oC, the surface often warms to the pressure melting point
(PMP). Importantly, the heat capacity of water (4186 J/kg oC) is roughly four times
greater than air (1005 J/kg °C). This means that under rainy conditions, the liquid water
adds a lot more heat to the ice surface. Water has roughly double the specific heat
capacity of ice (2093 J/kg °C); as such as rainwater hits the ice it takes a significant
amount of time to cool to the temperature of the surface. As the water is cooled and
frozen, latent heat transfer will occur, which in turn warms the ice (Benn & Evans, 2010).
Latent heat is the amount of energy consumed or lost during the change of state
between ice, water, and vapour (Benn & Evans, 1998). The amount of humidity in the air
and wind speed determines whether the type of latent heat energy transfer will be
evaporation, condensation, or sublimation (Benn & Evans, 2010). The latent heat of
fusion, which is the change of state between a solid or liquid, takes 334 J g-1 of energy.
The latent heat of evaporation is the change of state between a liquid and gas, and uses
2500 J g-1 of energy (Benn & Evans, 1998). Evaporation and condensation use over seven
times more energy than freezing and melting. This is important as evaporation reduces
the energy available to melt the glacier surface because of its high energy consumption
(Hock, 2005).
12
2.1.3 Turbulent Heat Fluxes
The combination of sensible heat transfer and latent heat transfer is known as the
turbulent heat flux. The transfer of energy is dependent on the wind speed, surface
roughness of the glacier (small scale), and the stability of the atmosphere (large scale)
(Pellicciotti et al., 2005).
On a small scale, as the wind blows across the glacier surface, the air is vertically
mixed by turbulent eddies and increases the gradient near the surface. The roughness of
the glacier surface creates drag and affects the dynamics of the turbulent eddies. Glacier
surface roughness can be determined through the collection of micro-topographic data or
high-resolution surface profiles. In glaciology, the aerodynamic roughness length
parameter (zo) is calculated from the micro-topographic data and defined as “the height
above a surface at which the extrapolated horizontal wind speed profile reaches zero”
(Brock et al., 2006). zo is an important control on the rate of turbulent heat transfer
between a glacier surface and the air above it. It is also an important factor as the
roughness impacts the ability of water to travel on the surface of the glacier (Knighton,
1972).
On a large scale, the turbulent heat flux and stability of the atmosphere is affected
by the exchange of sensible heat and water vapor over the glacier surface. An inversion
over the surface often occurs where the air density increases with decreasing temperature,
causing the cold air to sink below the warm air mass. This layer of chilled air (up to 100
m thick) flows down-glacier as a gravity current forms what is known as katabatic winds.
Generally katabatic winds are not felt on the terminus of Fountain Glacier, presumably
due to the ‘L’ form of the glacier.
13
2.2 Thermal Regime
Temperature variations between different glaciers and the thermal regime within
one glacier vary considerably. A temperate glacier is characterized as an ice body at the
PMP. The entire glacier has a consistent temperature at or close to 0oC, except for a thin
surface layer called the transient thermal layer (Benn & Evans, 1998). The transient
thermal layer is subject to seasonal temperature fluctuations. In winter, air temperatures
allow for below 0oC temperatures to penetrate into the surface. Conversely, in the
summer, warm air temperatures cause melt and a release of latent heat, which cause the
surface to remain at the PMP (Irvine-Fynn, 2004) (Figure 2.2).
Polythermal glaciers are defined as ice masses containing both temperate and cold
ice throughout the year (Blatter & Hutter, 1991). As such, they have a section below the
PMP, as well as a section at the PMP. Characteristically, a cold‐temperate ice transition
zone (CTZ) exists (Irvine-Fynn et al., 2011), whereby frequently a vertical temperature
gradient in the ablation zone contains a cold layer covering a temperate ice layer
(Wainstein, 2011). Near the margins, the cold ice layer generally reaches the glacier bed
such that it is often frozen to the underlying sediments (Rabus & Echelmeyer, 1997).
Typically, polythermal glaciers exist in regions with extended subfreezing winter air
temperatures and annual near-surface temperature gradients that produce a net conduction
of heat away from the glacier. As a result, there is a deeper penetration of the winter cold
wave (Björnsson et al., 1996, Blatter & Hutter, 1991). In a polythermal glacier only a
small portion of the transient thermal layer is at the pressure melting point, depending on
the season. Winter air temperatures cause the transient thermal layer to be below 0oC for
14
much of the year. Throughout the spring and summer the transient thermal layer is slowly
warmed and eventually reaches 0oC by the end of the summer (Figure 2.2).
Figure 2.2. Schematic representation of a thermal profile of a temperate and polythermal
glacier. Note that melt will occur at 0oC or PMP.
2.3 Glacial Hydrology
Glacial hydrology differs from terrestrial hydrology as complexities that exist as
part of the alluvial river regime such as sediment transport, bed and bank material
composition, and vegetation are absent (Knighton, 1985). As a result, the characteristics
of the glacier ice significantly influence the magnitude, timing, and variability of
streamflow (Anderson & McDonnell, 2005). In general, glacier hydrology is grouped into
three zones: supraglacial, englacial, and subglacial. However, the thermal regime of the
glacier has a direct influence on the hydrology of the glacier and affects the ability of
water to migrate into the different hydrologic zones. Reviews on temperate glacial
hydrology are written by Hubbard & Nienow (1997) and Fountain & Walder (1998) and
the most recent review of polythermal glacial hydrology is written by Irvine-Fynn et al.
15
(2011). As a result, I will focus on supraglacial hydrology in temperate and polythermal
glaciers in this section.
The supraglacial hydrology is heavily influenced by the transient thermal layer.
The polycrystalline structure of temperate ice allows for water to exist at the boundaries
of the ice grains and facilitates the development of veins (Hambrey & Alean, 2004). This
means that in the locations of temperate ice, water will travel both over the surface and
through the ice. Ice below the PMP creates a somewhat impermeable layer preventing
water from easily percolating through the ice (Irvine-Fynn et al., 2011). This causes
greater amounts of overland flow. During this study, the depth to the ice below 0oC was
approximately 1 m across much of the ablation area of the glacier. Water was observed
flowing both over the surface and within the upper 1 m of the surface layer. Importantly,
the depth of the transient thermal layer affects how, and how much water travels into a
supraglacial stream, and will also influence the amount the stream is able to incise into
the glacier.
2.4 Supraglacial Streams
Supraglacial streams are meltwater channels that exist in the ablation zone of
glaciers during the summer season. Once all the snow has melted, the water discharge
volume can be explained by three main water sources. These include: summer
precipitation, ablation of the overall glacier surface, and melting of the streambed
(Marston, 1983). Many comparisons to alluvial channels have been performed.
Supraglacial streams can have meanders, cutoff loops, anastomosing channels, and
propagating knickpoints. However, the major difference is the mechanism and timescale
of vertical and horizontal adjustment (Karlstrom et al., 2013).
16
2.4.1 Development and Evolution
In terms of the development and evolution of supraglacial streams, the glacier
structure influences the development, while the size and shape, and incision rate impact
the evolution. These will be discussed below.
2.4.1.1 Structural Influences
The structure, topography, and glacier motion greatly affect supraglacial stream
formation, morphology, and drainage. Glacier drainage is influenced by foliation and
crevasses. In areas with little slope, small rills form parallel to stratifications and
foliation, allowing for water to gather, and flow to commence (Hambrey, 1977). On some
glaciers, crevasses act as zones of weakness, facilitating the creation of moulins, and
allow for water to infiltrate the glacier (Hambrey & Alean, 2004). In other glaciers with
few crevasses, like Fountain Glacier, they play little to no role in the englacial hydrology.
Differential melt rates from various ice types form characteristic ridge and furrow
topography. This causes water to flow in certain directions within the drainage basin. In
relation to the stream, shear planes normal to the stream have been linked to the
formation of knickpoints and bed-shear stress influences the amplitude of meanders in the
stream (Knighton, 1981).
2.4.1.2 Stream Size and Shape
Supraglacial streams vary greatly in size and shape ranging anywhere from a few
centimetres in tiny rills to several meters wide (Hambrey & Alean, 2004). The drainage
structures can form either a dendritic pattern or a meander pattern (Hambrey & Alean,
2004). However, on Fountain and neighbouring Stagnation Glacier, the streams follow a
17
more parallel configuration. There are a number of parameters that influence the
development and establishment of the supraglacial drainage structures (Kostrzewski &
Zwoliñski, 1995). These include: drainage densities, incision rate, channel size, slope,
and discharge (Irvine-Fynn et al., 2011). There are complicated, yet direct relationships
between the incision rate, channel width, slope, and temperature (Fountain & Walder,
1998). Channel width is a function of summer air temperatures, and thermal and
hydraulic conductivity of the weathering crust (Leopold & Maddock, 1953). In addition,
increases to supraglacial stream velocity and width occur in response to higher discharges
and steeper gradients.
2.4.1.3 Stream Incision
Supraglacial streams exist in areas where the vertical channel incision is greater
than the rate of the glacier ablation (Knighton, 1981; Marston, 1983). Streambed erosion
occurs due to a combination of radiative, thermal, and mechanical melting. Radiative
melting occurs from solar radiation penetrating through the flowing water which allows
for melting of the streambed. Thermal melting from the heat exchange between the
warmer water and ice also occurs. Lastly, frictional/mechanical erosion can take place
along the streambed boundary as saltating sediment is transported downstream by the
flow.
2.4.2 Stream Characteristics
In terms of this study, the significant stream characteristics included discharge,
flow conditions such as the Reynold’s number and Froude number, and water
temperature.
18
2.4.2.1 Stream Discharge
Stream discharge is one of the most important factors in the evolution and life
span of a supraglacial stream. Discharge is derived from three sources: snow or ice melt
from the glacier surface, melt from the boundary of the channel itself, and rainfall during
the ablation season (Dozier, 1974). Major changes occur seasonally; as well as diurnally.
Early in the ablation season, when discharge is high, meanders develop and modifications
to existing channels can take place rapidly. When discharge rates decrease later in the
season, stream meanders have a tendency to straighten out (Hambrey, 1977). On a diurnal
time scale, the discharge is affected by the daily weather and strongly influenced by the
net radiation. Often, there is a two hour lag between peak downward radiation and peak
daily discharge, with a maximum discharge occurring between 15:00-17:00 local time
(Dozier, 1974). This lag is due to the time required for the latent heat of fusion to melt the
glacier ice and the water to travel to the stream (Marston, 1983).
2.4.2.2 Flow Conditions
The Froude number is the ratio of a characteristic velocity to a gravitational wave
velocity. It is an important aspect of the dynamics of the supraglacial stream as it allows
for the flow to be classified as subcritical (<1) or supercritical (>1). Streams flowing at
subcritical velocity commonly have a series of standing waves at the surface. Standing
waves or hydraulic jumps mark the sudden transition from subcritical to supercritical
flow (Dingman, 2009). Supercritical flow means that the velocity is greater than the
velocity of a translator wave, in this instance gravity waves can establish themselves in a
curve (Anderson & McDonnell, 2005). Lastly, when the Froude number is >2, roll waves
are capable of forming (Carver et al., 1994).
19
The Reynolds number is another important stream characteristic as it describes
whether the flow was laminar (<500) or turbulent (>2000). Laminar flow is characterized
by smooth, constant flow, while turbulent flow can produce chaotic vortices and eddies.
In supraglacial streams when the flow is considered turbulent, the three dimensional flow
patterns become unstable (Camporeale & Ridolfi, 2012). This can then impact the energy
dispersion and can significantly increase erosion rates.
2.4.2.3 Water Temperature
Water temperature is another important factor on the evolution of supraglacial
streams, as it influences the melt rate of the channel walls. Theory predicts stream
temperatures to be between 0.02-0.05oC, however field observations often range between
0 to 0.4oC. The temperature of the water is dependent on the air temperature, solar
radiation, discharge, and stream size and slope. In addition, sediments on the streambed
or within the water have been linked to the variations between the predicted and observed
values (Isenko et al., 2005).
2.4.3 Stream Morphology
Two intriguing supraglacial streams morphologies, which will be discussed
below, include the formation and development of meanders and step-pool sequences.
2.4.3.1 Sinuosity / Meandering
There is a large amount of evidence to support that supraglacial streams can
adjust their position very rapidly. Sinuosity, the ratio between the channel distance and
geodetic distance, is the calculation that determines how straight or curvy a stream is.
Meandering is the most common supraglacial channel pattern (Knighton, 1972), leading
20
to the assumption that this pattern is close to equilibrium (Dozier, 1974). The distribution
of shear against the stream banks, specifically asymmetrical helical flow, leads to
differential melting within the streambed, and therefore the meandering that occurs
(Marston, 1983; Parker, 1975). Erosive power causes the walls to undercut on the outside
bends (Hambrey & Alean, 2004) causing meanders to typically resemble a series of sine-
curves. The meander system is capable of migrating downstream leaving a visual record
of meander belts on channel walls and cutoff meander loops (Marston, 1983).
2.4.3.2 Step-pool Sequence
Another feature of supraglacial streams is the step-pool sequence. Step-pools have
been observed in mountainous, bedrock, arid, and supraglacial streams and are a common
element of fluvial environments (Knighton, 1998). In mountainous streams step-pools are
formed during extreme flood events, with a recurrence interval of approximately 20-50
years. High discharge is needed for mobilization of the rocks or large woody debris. A
small dam is formed from the trapped debris and a pool is scoured directly downstream.
Vatne and Refsnes (2003) assume that step-pools also form in glacial meltwater
streams during high discharge. However, they suggest that the cause is more from the
added energy, in the form of heat from frictional dissipation, versus the additional water.
In supraglacial streams an erosional pool is formed directly underneath a step from the
added energy in the waterfall, and a promontory (or lip) forms directly downstream from
the pool due to the dispersion of energy within the pool. The lip forms a reverse bed slope
within the stream profile, and thus forces water to travel against gravity to travel farther
downstream (Vatne & Refsnes, 2003).
21
STUDY AREA
The study area was on a small arctic glacier in Nunavut, Canada. A small
supraglacial stream in the centre of Fountain Glacier, on Bylot Island was the subject of
this research. Bylot Island is situated on the eastern margin of the Canadian Arctic
Archipelago, which extends from the northern point of Ellesmere Island to the southern
point of Baffin Island. The Archipelago is composed of a series of icecaps and icefields.
With the threat of global warming, the Archipelago has been identified as a highly
sensitive area (Gardner et al., 2011).
Bylot Island is the ideal study location as it is a protected island under the
jurisdiction of Sirmilik National Park. In addition, a number of studies have already been
undertaken on Bylot Island and a considerable amount of baseline data exists. A general
overview of the location, physical characteristics, and climate will be presented in regards
to Bylot Island. This will be followed by detailed information about Fountain Glacier and
the central supraglacial stream study area.
3.1 Bylot Island, Nunavut
Bylot Island is located directly north of Baffin Island; at latitudes of 72.5o and
74oN, and longitudes 76o and 81oW (Figure 3.1A) (Dowdeswell et al., 2007). It is
approximately 180 km along its NW-SE axis and 120 km at its widest point along its NE-
SW axis (Wainstein, 2011). It’s separated from Baffin Island by Eclipse Sound on the
southeast and Navy Board Inlet on the southwest. Bylot Island is an uninhabited island.
Pond Inlet, is the closest community and Environment Canada weather station, located 30
km away on Baffin Island.
22
Bylot Island has a mountainous region located at its centre and decreases in
elevation towards the coastal lowlands. The centre of the island is highly glaciated with a
large of number of nunataks. The Byam Martin Mountain chain follows the NW-SE axis
of the island, with the highest mountains being Angilaaq and Malik with elevations of
1844 m and 1905 m, respectively (Dowdeswell et al., 2007). The mountainous region
consists of Proterozoic igneous and metamorphic Canadian Shield bedrock. The lowlands
are composed of poorly consolidated non-marine shale, sandstone, and mudstone from
the Helikian and Cretaceous-Tertiary age (Wainstein et al., 2008; Irvine-Fynn, 2004).
Approximately 43% (4,783 km2) of Bylot Island is covered by glaciers, with the
largest glacier being 49 km long and 6.5 km wide (Dowdeswell et al., 2007). Sixteen
major glaciers have accumulation areas within the mountainous centre of the island and
flow through deeply carved valleys towards the coastal lowlands (Moorman & Michel,
2000a). For the most part, these glaciers terminate on land or in lakes; however, two of
the glaciers still calve into the sea (Dowdeswell et al., 2007). It is believed that at least
some of the glaciers on Bylot Island are polythermal (Irvine-Fynn et al., 2011).
According to previous research (Klassen, 1993), Bylot Island has been subjected
to four regional (foreign) and three local (native) ice sheet glaciations. The four major
foreign glaciation events that occurred were the Baffin, the Eclipse, the Button and the
Cape Hatt glaciation periods (Klassen, 1993). The Baffin glaciation was the most
widespread and caused massive ice movement on Bylot Island. During the Eclipse
glaciation, the marine channels surrounding the island were affected, whereas the marine
channels and coastal regions were affected during the Button glaciation. Lastly, the Cape
23
Hatt glaciation extended just south of Bylot Island and did not have much influence on
the island (Wainstein, 2011).
The three native glaciation periods include the Bylot, Aktineq, and Neoglacial.
The native glaciations were important as they caused the latest evolution in the landscape.
It is estimated that the Neoglacial period was the latest glacial period and occurred
approximately 120 years ago (Wainstein, 2011). At the present time, the glaciers are
either at or retreating from their last glacial maximum positions (Moorman, 2005). A
detailed description of the glacial history and quaternary geology of Bylot Island can be
found in Klassen (1993).
The climate of Bylot Island is cold and dry, and can technically be considered an
arctic desert (Wainstein et al., 2008). Bylot Island is in a zone of continuous permafrost
and it has been estimated that the permafrost range is between 200-400 m, while the
active layer ranges from 30-50 cm in thickness (Moorman, 2005). From the automatic
weather station located at the terminus of Fountain Glacier, the mean annual air
temperature (between 2000 – 2012) was approximately -11oC, while the mean annual
ground temperature was roughly -7.6oC. Pond Inlet receives an average annual
precipitation of less than 225 mm and the snow pack thickness at the Fountain Glacier
terminus is less than 80 cm in the winter (Moorman, 2005).
24
Figure 3.1. Study Area; A) Location map of Bylot Island within Canada; B) Close-up map of Bylot Island illustrating the
surrounding area; C) Landsat 7 image of Fountain Glacier.
25
3.2 Fountain Glacier
Fountain Glacier is officially designated as Glacier B26 by the Canadian Glaciers
Atlas of Canada (Inland Waters Branch, 1969). This glacier is located on the southern
half of Bylot Island, southwest of Stagnation Glacier (B28), and directly across Eclipse
Sound from the town of Pond Inlet, Baffin Island (Figure 3.1B). Fountain Glacier is
approximately 16 km long, 1.5 km wide at the terminus, and has a catchment area of 72
km2 (Wainstein et al., 2008). The elevation ranges from 245 meters above sea level
(MASL) to 1750 MASL and it has an average surface slope of 5.4° (Wainstein, 2011).
Fountain Glacier is somewhat unusual as it resembles an ‘L’ in form. The top two-thirds
of the glacier flow in a north to south direction, with a 90° turn occurring in the bottom
one-third. This causes the direction of flow to then become west to east (Figure 3.1C)
(Whitehead, 2013).
Research has shown that Fountain Glacier is a polythermal glacier with “cold
margins frozen to the glacial bed and a core of warmer ice” (Moorman & Michel, 2000a).
Today, most of the surface of Fountain Glacier is smooth and gently undulating, with
very few moulins or crevasses. There are two well defined canyons that have incised by
more than 20 m into the glacier. These canyons were formed by supraglacial streams
which discharge a large percent of the supraficial water. In addition to these two large
supraglacial streams, there are a number of smaller supraglacial streams in the terminus
region.
According to Whitehead (2013), there has been little net accumulation and an
increase in elevation of the equilibrium line altitude (ELA). Fountain Glacier has been
26
found to have a strongly negative mass-balance. Between 1958 and 2010, the terminus
has lost 35 - 45 m of ice and has retreated almost 250 m (Figure 3.2). Within the last 15
years, there has been a sudden change to this glacier’s equilibrium state. Due to dry
calving, the terminus has now become a 20-30 m high vertical wall, instead of a gentle
slope that was once walkable (Wainstein, 2011).
Figure 3.2. Temporal variation of Fountain Glacier's terminus. The retreat rate has
increased considerably since 1982 when a collapse feature on the northern half of the
snout was first observed (from Wainstien, 2011).
27
Ice flow is generally down-
glacier with rates ranging from near
zero for the marginal regions and up
to 8 m/year in the centre of the
glacier. Observations made over the
summers of 2009 and 2010, show
vertical ice loss in the terminus region
averaged between 2 m - 3 m over the
melt season, with an average rate of
ice loss, in late June/early July, being
around 3 cm/day (Whitehead, 2013).
Figure 3.3 gives specific mass
balance details between 2010 and
2011 including loss of ice thickness
and flow speed and direction
(Whitehead et al., 2013).
Another unique feature of
Fountain Glacier is the large
proglacial icing located at the
terminus (Wainstein, 2011). The icing
dominates the glacial outwash plain and has been observed over 11 km down-valley of
the glacier terminus (Moorman & Michel, 2000b); however, in 2011 the extent was only
1.2 km with a thicknesses of 3.6 m on average (Whitehead, 2013). During the summer,
Figure 3.3. A) Changes in ice thickness measured
from July 1, 2010 to July 2, 2011. Increases in
thickness to the east of the terminus reflect
changes to the proglacial icing; B) Horizontal
flow speed and flow direction between July 1,
2010 and July 2, 2011 (from Whitehead et al.,
2013).
28
the icing decreases in size from melt and reforms perennially during the winter.
Wainstein (2011) suggests that the preservation of the icing depends on the interactions
between the glacier, permafrost, and the proglacial valley. Fountain Glacier has a well-
developed subglacial hydraulic network and a large proglacial talik. This combination of
features allows the storage of pressurized water and conduction of water towards the
outwash plain. This is essential for the regeneration of the icing during the winter months
(Wainstein et al., 2014).
3.3 Central Supraglacial Stream
Research was conducted on the third largest supraglacial stream on Fountain
Glacier. This supraglacial stream is located in the center of the glacier and flows off the
front of the glacier in the form of a 13 m waterfall (Figure 3.4). In July of 2014, the
stream had a total length of 1190 m, with an average width of 40 cm, and depth of 8 cm.
As significant changes in sinuosity and slope occurred, the stream has been divided into 4
reaches for analysis purposes.
Reach 1 is the upper section of the stream. Reach 2 was bounded by two depth
sensors, one placed at the top of the reach and the other placed at the bottom of the reach.
Many of the measurements were conducted within Reach 2 as it was a representative
section of stream and did not have any rocks or debris to affect the albedo or ablation
rates. Reach 2 had a watershed area of 4616 m2, and an elevation change of 22 m. Reach
3 was between the bottom depth sensor and the stream bend, while Reach 4 was from the
stream bend to the waterfall (Figure 3.4).
29
Figure 3.4. The stream studied was located near the terminus of Fountain Glacier.
Arrows denote the 4 reaches. Key structural elements and stream characteristics are also
shown.
Changes in stream location between 2010, 2011, and 2014 can be seen in Figure
3.5A. Clearly, this stream is a perennial stream as it reoccupies a pre-existing channel
year to year with only minor changes. The central supraglacial stream migrates laterally
on the surface of Fountain Glacier. The majority of the upper section of the stream did
not alter more than a metre between 2010 and 2011. Conversely, in 2014, the stream
location appears to alternate to the N and S of the 2010/2011 position, in the order of
several meters. Figure 3.5B illustrates the lower stream changes, as well as the glacier ice
flow amount and direction. The lower section of the stream is again in a similar location
in 2010 and 2011, while in 2014 the position of the stream is either in the same location
or to the S by as much as 10 m. The most notable difference in the stream position is the
30
change in the waterfall location. In 2014, the waterfall was 55 m SW of its position in
2011.
Figure 3.5. A) Previous stream locations in 2010, 2011, and 2014 overlaid on an
orthoimage from 2011; B) Close-up map of the bottom section of the stream and glacier
ice flow amounts between July 2, 2011 and July 24, 2014.
N A B
31
METHODS
A number of complementary techniques were used to investigate the supraglacial
hydrology of Fountain Glacier. These include: the collection of ablation, albedo, surface
roughness, streambed profiles, stream discharge, positional and topographic information,
and meteorological data.
4.1 Ablation
In order to calculate the amount of supraglacial melt, a miniature ablation stake
survey was completed within the watershed of Reach 2. For this survey, white wooden
dowels, 5 mm in diameter, were imbedded into holes that were drilled using a standard
battery operated drill. The drill bit was 9.5 mm; slightly wider than the doweling. As the
drill bit was 45 cm long, this was as a limiting factor on the length of the doweling, which
were deliberately cut to be the same length. A total of 25 stakes were installed throughout
the watershed between July 7 and 8 (Figure 4.1). Stakes were placed in strategic locations
to be representative of microscale changes in the ice surface (Konzelmann & Braithwaite,
1995). The stake locations were chosen based on differences in elevation, orientation,
distance from the stream, as well as being placed in white and clear ice areas.
Ablation measurements were taken every 2-3 days within the study period.
Typically, stakes are surrounded by an ablation hollow (Konzelmann & Braithwaite,
1995). For this reason, the distance from the top of the stake to the ice surface was always
measured (using a measuring tape) to the nearest half cm on the up-glacier side of the
stake. As the stakes were composed of wood they were occasionally floating within the
ablation holes. Before measuring the stakes, they were pushed down to eliminate any
32
error in measurement from the buoyancy. Images and a written description of the area
directly surrounding each stake were taken three times throughout the study period; this
was to characterize changes in the ice surface. Also, the location of each stake was
acquired using a global positioning system (GPS). On July 11 and 16, the holes were re-
drilled in the exact same location to prevent the stakes from melting out.
The average ablation over the study period was converted into ice water
equivalent. The total water volume within the watershed area was calculated to determine
the amount that drained into the supraglacial stream.
Figure 4.1. Ablation stake locations within the Reach 2 watershed.
4.2 Albedo
Three albedo surveys were conducted throughout the study period by collecting
digital images using a Panasonic Lumix digital camera. A total of 120 images were
obtained on July 16, 22, and 23. To acquire an albedo value that was representative of the
Reach 2 watershed area, four near vertical images were taken at 30 locations. The four
33
images were taken in a circular pattern (NW, SW, SE, and NE), approximately 2 m away
from each of the 25 ablation stakes. The remaining five locations were taken specifically
of the water in the stream; one at the top, one at the bottom, and three in the middle of the
reach. Care was used to ensure pictures were taken at a similar height and angle.
Upon returning from the field, the images were processed using public domain
software “ImageJ”. The minimum, maximum, and mean digital number for each image
was obtained, and used to calculate the reflectance or relative albedo.
4.3 Surface Roughness
From micro-topographic data (high-
resolution surface profiles) aerodynamic
roughness length (zo) can be calculated. On
July 22 and 23, a manual micro-topographic
survey was conducted in the centre of the
Reach 2 watershed (Figure 4.2). Nails were
hammered into the glacier, with a 5 m long
horizontal reference string attached in between.
Measurements with a ruler were made between
the string and the ice surface at 10 cm
intervals, to the nearest mm. According to
Irvine-Fynn et al. (2014), using this technique will result in an accuracy of ±2.5 mm,
when measuring the distance between the ice surface and horizontal line.
Figure 4.2. Horizontal reference used
for the collection of micro-topographic
data for calculation of the
aerodynamic roughness length.
34
4.4 Streambed Profiles
Supraglacial streambed profiles are commonly measured using a Trimble Real
Time Kinematic (RTK) high precision GPS system (Karlstrom et al., 2013). By
completing two streambed surveys on different days, the rate of streambed erosion was
calculated in relation to time, and the changes in the streambed morphology can be
reviewed. Reach 2 was surveyed for the first time on July 13. The second streambed
survey took three days to complete; parts of the reach were surveyed on July 21, 22, and
23. A large amount of overlap between the survey sections, allowed for the three survey
sections to be adjusted to all reflect the streambed elevation on July 22. At this time, the
RTK GPS unit was attached to the survey pole and measurement points were manually
collected in the centre of the streambed at each flexion point; whether it was a bend or a
sudden change in elevation from a step in the stream. According to Trimble, the RTK
GPS system is accurate ±1 cm in the horizontal and ±2 cm in the vertical (‘Trimble’,
2014). Stream depth was also recorded every few survey points.
4.5 Supraglacial Stream Discharge
Supraglacial stream discharge was determined using three methods. The velocity-
area method provides discharge, but is labour intensive. To increase the temporal density
of the discharge estimations, discharge rating curves were made so that discharge
estimates could be made from simple depth measurements (Oostrem & Brugman, 1991).
By combining data from the depth sensor and barometric pressure sensor, stream depth
could be obtained. Finally, two time-lapse cameras were positioned to capture changes in
stream discharge at the waterfall location.
35
4.5.1 Velocity-Area Method
The velocity-area method involves measuring the velocity in cross-sections at set
intervals depending on the width of the stream. These measurements were conducted
approximately twice a day, every second day, 1 m downstream from the depth sensors. A
Flowatch Meter was utilized to determine velocity and a measuring tape was used to
obtain stream width and depth. The average water velocity was taken every 10 cm, at a
40% depth from the bed surface. The Flowatch Meter specifications indicate the
sensitivity range to be <0.083 m/s- <0.1 m/s, with a precision of ±2% ('Flowatch', 2009).
The manual distance measurements of both width and depth were taken to the nearest cm.
Multiplication of the cross-sectional area by the velocity for each level was totalled, to
give the total discharge (Dackombe & Gardiner, 1983). Measurements were taken 16
times to ensure the range of discharges were captured. From these series of
measurements, a rating curve was calculated to provide the discharge from the water
depth alone.
4.5.2 Depth Measurements
The depth measurements were taken with a “Model 3001 Solinst Levelogger
Junior Edge”, designed for a depth of less than 5 m. Unlike a terrestrial stream, the
streambed of supraglacial streams decrease overtime due to the streambed melt erosion.
To compensate for this, a 50 mm diameter Kovac ice auger was used to drill a hole into
the centre of the supraglacial streambed. The depth sensors, originally placed on July 7,
36
were installed in the holes with rocks on top
to prevent them from floating. A string was
attached to the sensor and a rock on the edge
of the stream bank to ensure the sensor wasn’t
carried downstream (Figure 4.3). One was
installed at the top of Reach 2 and the other at
the bottom of Reach 2 (herein referred to as
the top depth sensor or bottom depth sensor)
(Figure 4.4). On a number of occasions, the
distance between the depth sensor and
streambed was measured. Importantly, this
allowed for the calculation of streambed
erosion. On July 13 and 19 the top depth
sensor was temporarily removed to re-drill the hole. Also on July 13, the depth sensor
was partially encased in plastic in an attempt to prevent it from freezing into the ice.
As the depth sensor only measures absolute pressure, a barometric pressure sensor
is required for accurate barometric compensation in order to calculate water depth
(‘Solinst’, 2014). A “Solinst Baralogger Edge” was placed at the base of Fountain Glacier
(beside Bylot-1 weather station), and recorded atmospheric pressure every minute from
July 7-24. According to Solinst, the baralogger has an accuracy of ±0.05 kPa and the
levelogger depth sensor has an accuracy of 0.3 cm. The levelogger also has a temperature
sensor that is accurate to ±0.05°C, however it should be noted that the temperature
calibration range is only between 0º-50ºC (‘Solinst’, 2014).
Figure 4.3. Depth sensor installation
within the streambed. Note for the
purpose of this image the security
string was recoloured black, but in
reality it was white so as not to impact
melt of the glacier.
37
Figure 4.4. Location of the depth sensors, evaporation pan, and roughness survey within
the Reach 2 watershed.
4.5.3 Time-Lapse Imagery
On July 10, a “Wingscapes WSCA04” outdoor time-lapse camera was positioned
in the general direction of the glacier terminus (~450 m away) and captured the location
where the supraglacial stream flows off the glacier in the form of a waterfall. The camera
was set to take images every 10 min. On July 13, the temporal frequency of this camera
was changed to take photos every 5 minutes. On July 21, an additional camera was
mounted ~75 m from the waterfall and was set to take images on a 30 s time interval. An
example of two time-lapse images can be seen in Figure 4.5, the two images were taken
on July 23, exactly one minute apart. In the first image the waterfall is flowing, while no
discharge was observed in the second image. Using all the images collected an
understanding of the stream discharge pulse frequency could be obtained.
38
Figure 4.5. Time-lapse camera placed ~75 m from the base of the waterfall used to
collect pulsating frequency. Note that the images were taken one minute apart on July 23,
2014. A) Shows the waterfall at 8:15 with high discharge; B) Displays the waterfall at
8:16 with no discharge.
4.6 Positional and Topographic Information
Positional and topographic information included the collection of written
descriptions, handheld GPS coordinates, and high precision RTK GPS data. Firstly,
detailed descriptions of Reach 2 were taken on a frequent basis and notes on any
observed changes were recorded.
Secondly, GPS coordinates within the Reach 2 watershed included the ablation
stakes, micro-topographic survey, and water balance pan locations. Although the majority
of this study occurred within the Reach 2 watershed, the location/length of the entire
39
stream was mapped using a handheld GPS, and the entire stream watershed was explored
on several occasions.
Thirdly, the Reach 2 watershed boundary was surveyed using the RTK GPS unit.
During this survey, the system was attached to a backpack and run in continuous mode,
taking readings every 5 s. The watershed boundary was obtained by walking the full
extent on July 19. The watershed is not only influenced by topographic highs, but also the
longitudinal foliation/structure of the glacier. For this reason, the first researcher walked
ahead of the second researcher who had the GPS, to ensure that the correct boundary was
acquired.
The last source of geographic information was two orthophotos (from 2010 and
2011) that were obtained from Whitehead. These were used to determine changes in the
glacier surface, locations of transverse fractures, and past positions of the supraglacial
stream.
4.7 Meteorological Data
A number of meteorological data sources ranging in spatial distribution and
temporal frequency were utilized for analysis purposes. Firstly, air temperature, relative
humidity, wind speed and direction, and net radiation on an hourly basis were obtained
from a Campbell Scientific Automatic Weather Station (AWS). While there are four
AWS on Bylot Island, data used was from Bylot-3 which is located in the next valley
over, about halfway up Stagnation Glacier (approximately 4 km away from the study
area). The second meteorological data source was a precipitation/evaporation pan
40
temporarily installed in the centre of the Reach 2 watershed, on the surface of Fountain
Glacier (Figure 4.4).
A water balance pan was fabricated on site using the bottom portion of a 5 gallon
plastic pail (diameter of 29.8 cm) and the original water level was set 5 cm from the
bottom. Measurements of the change in depth were taken on a frequent basis (every field
day on the glacier surface). Using a ruler, the depth was recorded to the nearest mm in
four equally spaced locations around the cylindrical pan and averaged for the mean depth
to account for differential melt. By comparing (subtracting or adding) the change in depth
between subsequent days the amount of precipitation/evaporation was calculated.
Differential melt of the glacier surface caused the pan to tilt and spill; on July 11,
the pan was re-fabricated by creating a hexagonal base for the cylindrical dish to sit upon
and held in place using the weight of two rocks on either side (Figure 4.6). In addition,
the amount of precipitation in the second
half of the study period caused
overflowing of the pan to occur; this
resulted in unattainable precipitation and
evaporation values. To compensate for
missing data, the daily precipitation
amount was obtained from the
Environment Canada weather station
located in Pond Inlet.
Figure 4.6. Fabricated evaporation pan on
the surface of Fountain Glacier.
41
CHAPTER FIVE: RESULTS
The results chapter is divided into three key sections: meteorological data, glacier
surface characteristics, and stream characteristics. Each individual section is crucial for
the understanding of the overall connections examined in the analysis and discussion.
5.1 Meteorological Data
5.1.1 Air Temperature
During the entire study period, the average air temperature was 4.6oC. The
maximum was 11.8oC, which occurred on July 10, and the minimum was 1.2oC on July 8.
A somewhat cyclical pattern can be seen in Figure 5.1. Peaks in temperature occurred
during the day, with lower temperatures occurring during twilight. The air temperature
remained fairly constant between July 7-13 and July 14-22. The average the first week
was 4.4oC, and increased slightly to 4.8oC the second week.
Figure 5.1. Air Temperature recorded at Bylot-3 AWS from July 7-22, 2014.
5.1.2 Net Radiation
Although the study site was north of 60o latitude, in the zone where the sun never
truly sets during the summer, the net radiation still fluctuated on a daily cycle. The net
0
2
4
6
8
10
12
14
07 08 09 10 11 12 13 14 15 16 17 18 19 20 21 22
Air
Te
mp
era
ture
(°C
)
Day (July 2014)
42
radiation ranged between -78 to 313 W m-2, with the average of 54 W m-2 for the entire
study period (Figure 5.2). The average net radiation from July 7-13 was 66 W m-2 and the
average from July 14-22 was less, with 43 W m-2.
Figure 5.2. Net radiation recorded at Bylot-3 AWS from July 7-22, 2014.
5.1.3 Relative Humidity
The average relative humidity was 82%, but the values varied significantly during
the study period (Figure 5.3). The minimum was 45% on July 10 and the maximum was
100% on a number of days during the second week of the study. Again, there was a
difference between July 7-13 and July 14-22; the average was 76% and 87% respectively.
Figure 5.3. Relative humidity recorded at Bylot-3 AWS from July 7-22, 2014.
-100
0
100
200
300
400
07 08 09 10 11 12 13 14 15 16 17 18 19 20 21 22
Ne
t R
adia
tio
n (
W m
-2)
Day (July 2014)
40
50
60
70
80
90
100
07 08 09 10 11 12 13 14 15 16 17 18 19 20 21 22
Re
lati
ve H
um
idit
y (%
)
Day (July 2014)
43
5.1.4 Water Balance
The precipitation/evaporation pan installed on Fountain Glacier suggests that on
July 13 and 23 the net water balance was negative; water was removed from the glacier
surface in the form of evaporation at the rate of 2-3 mm/day, respectively (Table 5.1).
Between July 14-16 and July 20-22, the net water balance was positive, with values of 9
and 12 mm, respectively. This was due to a number of precipitation events.
Table 5.1. Evaporation, rainfall, and net water balance amounts for the Reach 2
watershed from July 7-23, 2014.
Day
(July
2014)
Fountain
Glacier
Rain /
Evaporation
Pan (mm)
Calculated
Evaporation from
Glacier surface
(mm)
Pond
Inlet
Rain
(mm)
Net
Water
Balance
(mm)
Surface Water
Balance within
Reach 2
Watershed (m3)
07
No Data
-1 0 -1 -4.6
08 -2 0 -2 -9.2
09 -2 0 -2 -9.2
10 -2 0 -2 -9.2
11 -3 0.3 -2.7 -15.2
12 -2 0 -2 -9.2
13 -2 -2 0 -2 -9.2
14 9
-1 2.8 2 9.2
15 -1 6.1 5 23.0
16 -1 2.5 2 9.2
17 No Data -4 0 -4 -18.5
18 -1 5.8 4.6 21.2
19 -1 1.5 0.5 2.3
20 12 -2 6.1 4 18.5
21 -3 0.3 8 36.9
22 -1 0 -1 -4.6
23 -3 -3 0 -3 -13.9
44
As the data from the Fountain Glacier precipitation/evaporation pan is sporadic
and incomplete, further meteorological information was compiled. The daily evaporation
amount was calculated and precipitation data from the Environment Canada Pond Inlet
station was obtained.
According to Kojima (1979), the following equation estimates the evaporation on
a snow surface:
𝐸′𝑎𝑐 = 1x10−3𝑢1(𝑒𝑠𝑎 − 𝑒𝑎1) x 240 [Eq. 5.1]
where 𝐸′𝑎𝑐 is the evaporation on the snow surface (mm d-1), u1 is the wind speed (m s-1),
𝑒𝑠𝑎 is the saturation vapour pressure at air temperature (hPa), and 𝑒𝑎1is the vapour
pressure 1 m above the snow surface (hPa).
Ohno et al. (1992) have modified Kojima’s equation to allow for the evaporation
from a melting glacier surface to be calculated. Ohno’s equation can be seen below:
𝐸′𝑎𝑐 = 1x10−3𝑢1(𝑒0 − 𝑒𝑎1) x 280 [Eq. 5.2]
where e0 is the saturation vapour pressure of the melting ice (6·11 hPa).
Vapour pressure can be calculated using the following:
𝑒𝑎1= 𝑅ℎ x 𝑒𝑠𝑎 [Eq. 5.3]
where 𝑅ℎ is relative humidity (%) and 𝑒𝑠𝑎 can be calculated as follows:
𝑒𝑠𝑎 = 6.11x10[7.5𝑇𝑎 / (237 + 𝑇𝑎)] [Eq. 5.4]
where 𝑇𝑎 is the air temperature (oC).
Using Ohno’s equation, the average evaporation from the surface of Fountain
Glacier was estimated to be -2 mm/day from July 7-23 (Table 5.1). The evaporation was
45
the greatest on July 17 with a value of -4 mm/day, and the least on July 7, 14, 15, 16, 18,
19, and 22 with -1 mm/day.
The Pond Inlet weather data shows that from July 7-23 there was a total of 25 mm
of rainfall. Table 5.1 shows that high rainfall amounts occurred on July 15, 18, and 20
with 6.1 mm, 5.8 mm, and 6.1 mm, respectively.
The net water balance was estimated by combining the data from Fountain
Glacier, calculated evaporation, Pond Inlet data, as well as direct observations (Table
5.1). The net water balance was negative from July 7-13 and positive from July 14-21,
except for July 17. The net water balance became negative again on July 22 and 23. In
total there was 26 mm of rain that fell dispersed over 6 days and a total of 21 mm of
evaporation occurring over 10 days. The net water balance within Reach 2 can be seen in
Table 5.1 and was calculated by multiplying the net water balance by the Reach 2
watershed area (4616 m2).
Figure 5.4 presents the occurrence of rain events and the water balance in
graphical format for the Reach 2 watershed area. Time-lapse cameras indicated that it
rained on July 11, 14, 15, 16, 18, 19, 20, and 21. Rainfall generally occurred in both Pond
Inlet and in the Fountain Glacier valley. The only distinguished exception was on July 21
when Fountain Glacier was under a rain cloud, while Pond Inlet was visually observed to
have clear sunny skies above it. It should be noted that although a rain event can be seen
on July 11, the net water balance was negative due to the low amount of precipitation and
high amount of evaporation.
46
Figure 5.4. Net water balance and rainfall events for the Reach 2 watershed from July 7-
22, 2014.
5.2 Glacier Surface Characteristics
5.2.1 Structure
Figure 5.5 shows the structure on the terminus of Fountain Glacier. The structure
in the lower 1.5 km includes transverse fractures and crevasse traces. The region is
dominated by steeply dipping transverse fractures, which are parallel to the margin. The
transverse fractures are crossed by a series of closed crevasse traces that are positioned
orthogonal to the glacier margin (Hambrey, pers. comm.).
07 08 09 10 11 12 13 14 15 16 17 18 19 20 21 22
0
5
10
15
20
25
30
35
40
-20
-10
0
10
20
30
40
07 08 09 10 11 12 13 14 15 16 17 18 19 20 21
Ne
t W
ate
r B
alan
ce (
m3 )
July (2014)
Water Balance
Rain Event
47
Figure 5.5. Image of terminus of Fountain Glacier from 2014 showing structural
elements.
5.2.2 Ablation
The ablation was measured every one to three days from July 8-22 at 24 sites
within the Reach 2 watershed. As measurements were not taken on a consistent time
scale, the average daily ablation (ADA) (cm/day) was calculated in order to compare the
data:
ADA = (MA x h / 24) [Eq. 5.5]
where MA was the measured ablation (cm) and h was the number of hours between
survey measurements. MA and h were then divided by 24 hours to obtain the daily
ablation (Table 5.2).
48
Table 5.2. Measured ablation, average daily ablation, and amount of water equivalent
within the Reach 2 watershed from July 8-22.
Figure 5.6 displays an example of the daily amount of melt at each site on July 9-
10, 16-17, and 21-22. The ablation variability between different days is a function of net
radiation; however, there is also a variability in the amount of melt between different
sites on the same day. On July 9-10, site 18 had the greatest ablation with 10.1 cm, and
site 13 had the least amount of melt with 2.9 cm. The highest amount of melt was seen at
site 22 on July 16-17 with 8.9 cm of melt, and the lowest at site 6 with only 2.4 cm of
melt. On July 21-22, the greatest ablation was at site 10 with 3.3 cm, and the least at site
18 with 1.3 cm.
Date
Number
of Hours
Average
Measured
Ablation (cm)
Average Daily
Ablation
(cm/day)
Watershed Ice
Water Equivalent
(m3/day)
July 8-9 25 6.1 5.9 243.3
July 9-10 50 11.4 5.5 227.3
July 10-11 5.5 227.3
July 11-12 52 12 5.8 239.7
July 12-13 5.8 239.7
July 13-14 68.5 10.9 3.8 158.7
July 14-15 3.8 158.7
July 15-16 3.8 158.7
July 16-17 23.5 4.4 4.5 186.7
July 17-18 50.5 7.8 3.7 155.5
July 18-19 3.7 155.5
July 19-20 66 6.4 2.3 96.7
July 20-21 2.3 96.7
July 21-22 2.3 96.7
49
Figure 5.6. Ablation amounts at each of the 24 sites for July 9-10, 16-17, and 21-22.
Images of the sites that had the highest and lowest ablation rate on July 9, 16, and
22 are displayed in Figure 5.7 below. July 9 - site 13 appears to be in white ice and site
18 was located on the edge of a fracture. July 16 - site 6 and site 22 varied in ablation
rates, yet look relativity similar in appearance. Both sites had clear blue ice, with water
on the surface, and cryoconite holes surrounding the stake. On July 22 - site 6 had the
0123456789
1011
0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24
Ab
lati
on
(cm
/day
)
Site Number
July 9-10 July 16-17 July 21-22
Figure 5.7. Images of the sites that had the highest and lowest ablation on July 9, 16, and
22.
50
greatest amount of melt, while site 18 had the least. Site 6 is on the boundary between
white and clear ice, with a fair amount of cryoconite in the vicinity, whereas site 12 is
located in white ice with no cryoconite surrounding it.
As seen in Figure 5.8, the ablation amount appears to be spatially random within
the Reach 2 watershed. From July 8-22, site 12 had the greatest amount of melt with a
total of 93.4 cm, while site 16 had the least amount of melt with only 45.5 cm. Site 12 is
located in the centre of the watershed along the northern boundary, whereas site 16 is
located in the lower section of the watershed also on the northern edge. There does not
appear to be a relationship between the sites closer to the stream, farther downstream, or
on one side of the stream versus the other side.
Figure 5.8. Total ablation at each site from July 8-22.
Ablation results are typically expressed in ice water equivalent (WE) (m/day):
WE = ADA x (ρi / ρw) [Eq. 5.6]
51
where ρi is ice density (900 kg/m3), ρw is water density (1000 kg/m3), and ADA is
average daily ablation (m/day) (Cuncio, 2003). To calculate the ice water equivalent
within the watershed (WWE) (m3/day):
WWE = WE x AW [Eq. 5.7]
where AW is the watershed area (m2), in this case 4616 m2.
The average hourly ablation and Reach 2 watershed ice water equivalent for each
day are displayed in Table 5.2. Both metrics will be used for comparisons within the
discussion section. The total daily average amount of melt was 57.5 cm (2388.8 m3)
during the entire study period, with an average of 4.2 cm/day (174.5 m3/day). The
greatest amount of ablation occurred from July 8-9 with an average of 5.9 cm/day (243.3
m3/day), while the lowest amount of ablation was observed from July 19-22 with an
average of 2.3 cm/day (96.7 m3/day). Importantly, the ablation rate changed drastically
throughout the study period. From July 8-13 the average ablation was 5.4 cm/day (222.7
m3/day), whereas from July 14-22 the average ablation decreased to 3.3 cm/day (138.1
m3/day).
5.2.3 Relative Albedo
Although albedo can be calculated using a digital camera, a reference is needed
for the value to be converted into absolute albedo. Typically, either an image of a white
piece of paper is taken or an actual albedo value is obtained using a pyranometer. As no
albedo reference value was acquired during the surveys, the calculated albedo is only
considered a relative albedo value. In order to compare the various albedo surveys that
52
were taken on different days, the images had to be normalized due to variations in
lighting conditions.
Black cryoconite existed in all of the images and presumably the colour of the
cryoconite did not change over the duration of the study period. Using this assumption,
the minimum digital number (representing the black cryoconite) was obtained for each
survey day. A minimum linear shift (MLS) was given to the data,
MLS = (Imax - Imin ) - Omin [Eq. 5.8]
where Imax was the maximum digital number of the image, Imin was the minimum
digital number of the image, and Omin was the overall minimum for that survey day (or
during constant lighting conditions). This equation shifted all of the values for each
survey day to correspond to the minimum digital number to compensate for the false
values caused by daily variations in the lighting conditions.
Next, mean stretch (MS) was given to the data under the assumption that the range
in the digital numbers was also a function of the lighting conditions and that the range in
the digital numbers should not have changed throughout the study period. The following
equation was used to normalize the maximum,
MS = Omax / MLS x Imean [Eq. 5.9]
where Omax was the overall maximum digital number for that survey day and Imean was
the mean digital number of the image. Lastly, the normalized relative albedo (ALrel) was
scaled using the digital number total:
ALrel = MS / 256 [Eq. 5.10]
On July 16, 22, and 23 complete albedo surveys were conducted and included
four images of the glacier surface at 30 locations (Figure 5.9). On July 16, the relative
53
albedo ranged between 0.56-0.94, with an average of 0.72. On July 22, the average
albedo was 0.74, with a maximum of 0.88, and minimum of 0.55. The last survey,
conducted on July 23 showed the average glacier albedo was 0.60, with a maximum of
0.78, and minimum of 0.52.
Figure 5.9. Relative albedo amounts at 30 sites on July 16, 22, and 23.
Figure 5.10 shows images taken in four directions at site 3, on July 16, 22, and 23.
At each of the four sites, the albedo fluctuated slightly depending on the surface
characteristics. On July 16, the albedo ranged between 0.69-0.73 in the four directions.
The albedo was lowest in the NW direction, with the image showing widespread
cryoconite. The albedo was highest in the SE direction, with half the image having
cryoconite and the other half appearing to have white ice. On July 22, site 3 had an
albedo that ranged between 0.77-0.83. The albedo was lowest in the NE direction,
presumably due to the presence of a large fracture running through the centre of the
image. On July 23, the albedo at site 3 was lowest in the NW and NE directions with
0.57, and highest in the SE direction with 0.59. The NW and NE images do not appear to
0.4
0.5
0.6
0.7
0.8
0.9
1
0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30
Re
lati
ve A
lbe
do
Site Number
July 16 July 22 July 23
54
have any melt, while the SW and SE images appear to have greater reflectance from light
bouncing off slightly wet looking ice.
As seen from Figure 5.10, there is a clear difference in the reflectivity of the
surface between the various days. Although it would not be suggested that July 22 had
the highest relative albedo due to lighting conditions, the surface was indeed the most
reflective on this day. The ice surface was composed of white ice, with only small
amounts of cryoconite present. As seen from the July 16 images, the surface appears to
have water on the surface of clear blue ice. The relative albedo was the lowest on July 23.
The surface displayed primarily white ice, which appeared to be in a state of melting.
Figure 5.10. Variations in relative albedo in 4 directions on July 16, 22, and 23, 2014.
55
The relative albedo was also estimated using the images taken at the 24 ablation
sites. Figure 5.11 presents the variations between the different sites on July 7, 8, and 9,
while, Figure 5.12 displays the sites on July 16, 22, and 23. The average relative albedo
on July 7, 8, 9, 16, 22, and 23 was 0.57, 0.67, 0.69, 0.72, 0.77, and 0.65, respectively. As
seen in the figures, some of the sites had a large amount of deviation during the same
day. The range on July 7, 8, 9, 16, 22, and 23 was 0.05, 0.18, 0.17, 0.21, 0.3, and 0.24.
Figure 5.11. Relative albedo computed from images taken at the 24 ablation sites on July
7, 8, and 9.
Figure 5.12. Relative albedo computed from images taken at the 24 ablation sites on July
16, 22, and 23.
0.5
0.6
0.7
0.8
0.9
1.0
0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24
Re
lati
ve A
lbe
do
Day (July 2014)
July 7 July 8 July 9
0.5
0.6
0.7
0.8
0.9
1.0
0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24
Re
lati
ve A
lbe
do
Day (July 2014)
July 16 July 22 July 23
56
Figure 5.13 below shows the albedo values from site 3 on July 7, 9, 16, and 23.
The albedo on July 7 and 23 are similar, with values of 0.58 and 0.59, respectively. The
images also appear to have the same characteristics; displaying white ice and a few
cryoconite holes. The images from July 9 and 16 have a similar albedo of 0.72 and 0.75,
yet are quite different in appearance. On July 9, the surface was clean white ice; whereas
on July 16, the image displays water overtop of clear ice with cryoconite present.
Figure 5.13. Images used to obtain the relative albedo at site 3 on July 7, 9, 16, and 23.
Within the Reach 2 watershed, the distribution of albedo values was highly
scattered and evidently changed on a frequent basis (Figure 5.14). On July 16, the albedo
was highest at sites 21 and 23, and lowest at sites 4, 5, and 15. The two locations with the
highest albedo were located on the south side of the Reach 2 watershed, and the three
sites with the lowest albedo were located on the north side of the Reach 2 watershed.
57
Figure 5.14. Map of the relative albedo at each of the ablation sites on July 16, 22, and
23.
58
The highest albedo values on July 22 were recorded at sites 10 and 17, while the
lowest albedo occurred at sites 1 and 24. Lastly, on July 23, the albedo was highest at
sites 7, 8, 9 and 22, and lowest at sites 19 and 24. Interestingly, on July 23 the sites with
the highest albedo were all located in the center of the watershed, where the lowest
albedo values had been on July 16.
During the study period, the average daily albedo was estimated using the results
from the albedo and ablations surveys, as well as from additional glacier images, and in
some cases the value was inferred between days (Table 5.3). The daily albedo ranged
Table 5.3. Relative albedo computed from various sources and the estimated daily
relative albedo from July 7-23.
Date Albedo
Survey
Ablation
Survey
Additional
Images
Inferred
between days
Daily
Albedo
July 7 0.57 0.57
July 8 0.67 0.67
July 9 0.69 0.69
July 10 0.64 0.64
July 11 0.59 0.59
July 12 0.63 0.63
July 13 0.66 0.66
July 14 0.55 0.55
July 15 0.63 0.63
July 16 0.69 0.72 0.72
July 17 0.70 0.70
July 18 0.70 0.70
July 19 0.69 0.69
July 20 0.69 0.69
July 21 0.74 0.74
July 22 0.76 0.77 0.77
July 23 0.65 0.66 0.66
59
from a low of 0.59 on July 11, to a high of 0.77 on July 22. From July 7-13 the average
daily albedo was 0.63; however, from July 14-23 the average was much higher with a
value of 0.69.
5.2.4 Surface Roughness
Aerodynamic roughness length parameter (zo) is an important variable in the
energy balance as it affects the rate of turbulent heat transfer between a glacier surface
and the air above it. From the micro-topographic data/high-resolution surface profiles
taken on July 22 and 23, the zo can be calculated as follows:
𝑧𝑜 = 0.5ℎ∗ (𝑠
𝑆) (Eq. 5.11)
where h* is the effective height for the roughness elements calculated as twice the
standard deviation of the elevations with the mean elevation set to 0 (Brock et al., 2006),
s is the silhouette area of roughness elements (area measured in a vertical plane
perpendicular to the wind direction), and S is the frequency per unit area (Munro, 1989).
Subsequently, the silhouette area can is solved by:
s = h* x X / 2f (Eq. 5.12)
where X is the width of a typical element and is defined as the length of the traverse, and
f is the frequency. Lastly, the frequency per unit area can be calculated using the
following:
S = (X / f)2 (Eq. 5.13)
Figure 5.15 displays the manual surface roughness measurements and the average
surface height for July 22 and 23. From the measurements taken above the surface
reference pole, the average surface height was determined to be 9.6 cm on July 22, and
60
10.2 cm on July 23. On July 22, the maximum height was 14.5 cm, with a minimum of
7.2 cm, leading to a range of 7.3 cm. On July 23, the range was 10.1 cm, as the maximum
was 17 cm, and the minimum was 6.9 cm.
Along the 500 cm profile, the aerodynamic roughness length parameter was
calculated to be 1x10-4 m on July 22. On July 23, the aerodynamic roughness was
noticeably higher with a value of 6.7x10-4 m.
Figure 5.15. Glacier surface roughness measurements taken on July 22 and 23.
Aerodynamic roughness takes into consideration both the glacier topography, as
well as the microscale changes in the weathering crust. Unfortunately, surveys to
calculate the aerodynamic roughness length parameter were only completed on July 22
and 23. It is assumed that the glacier topography did not significantly change during the
study. However, during the study period the dynamics of the weathering crust changed
drastically. As a result, a discussion of the microscale roughness is necessary.
The microscale roughness was evaluated using the images collected at each of the
ablation stake sites on July 9, 16, and 22. The roughness was assessed on a scale from 1-
4, 1 being “not rough” and 4 being “very rough”. As seen in Figure 5.16, many of the
-20-18-16-14-12-10
-8-6-4-20
0 50 100 150 200 250 300 350 400 450 500
Ele
vati
on
(cm
)
Profile Length (cm)
July 22July 23
61
sites on July 9 were “very rough”. On July 16 most of the sites were “slightly rough”,
while on July 22 the roughness ranged from “not rough” to “rough”.
Figure 5.16. Micro-scale surface roughness estimated from images taken at the 24
ablation sites on July 9, 16, and 22.
Figure 5.17 below shows images used to assess the roughness values from sites 2,
7 and 8 on July 9, 16, and 22. On July 9, deep cryoconite holes existed on the surface of
the glacier. Sediment and water were located within the holes. On July 16, the majority of
the holes had disappeared leaving sediment in topographic lows. On July 22, very few
cryoconite holes were present, the glacier surface was extremely smooth, and much of the
sediment had disappeared.
0.5
1.5
2.5
3.5
0 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24
Surf
ace
Ro
ugh
ne
ss
Site Number
July 9 July 16 July 22Very
Rough
Rough
SlightlyRough
NotRough
62
5.3 Supraglacial Stream Characteristics
5.3.1 Watershed / Stream Dimensions
As mentioned in the study area section, the Reach 2 watershed area was 4616 m2
and had a stream length of 228 m. Figure 5.18 is a schematic diagram of the Reach 2
watershed illustrating cross-sections of the stream at the top and bottom of the reach.
At the top of Reach 2 (top depth sensor), the average stream width was 38 cm and
the average stream depth was 8 cm. The bottom of the reach (bottom depth sensor) had a
larger average width of 41 cm, and a slightly lower average depth of 7.5 cm. The stream
bank height (from the top of the ice to the water level) varied along the length of the
stream, with the south side always greater than the north side. At the top of the reach, the
Figure 5.17. Roughness estimated from images take at ablation stakes 2, 7 and 8 on July
9, 16, and 22.
63
bank height to the south was 45 cm and 10 cm to the north, whereas at the bottom it was
37 cm to the south and 25 cm to the north.
Figure 5.18. Schematic diagram of the watershed with stream cross-sections at the top
and bottom of the Reach 2.
5.3.2 Stream Meandering
On July 13 and 22, positional data of the stream was collected for Reach 2. This
can be seen in Figure 5.19; which illustrates the plan-view of the stream and the
meandering that occurred. The position of the stream remained similar with the two
profiles overlapping; however, slight differences can be seen on the outside of the stream
bends where erosion was occurring.
64
Figure 5.19. Plan-view of Reach 2 illustrating meandering and slight changes in stream
position on July 13 and 22.
5.3.3 Streambed Erosion
The difference in elevation between the RTK surveys and the difference between
streambed measurements taken at the bottom depth sensor give dispersed rates of
streambed erosion. However, by combining the two data sets, the daily rate of streambed
erosion can be obtained.
Similarly to the ablation results, the stream erosion (SE) (m/day) was converted to
be expressed as ice water equivalent. To calculate the ice water equivalent along the
Reach 2 (REWE) (m3):
REWE = SE x As [Eq. 5.14]
where As is the stream area covered by water (m2). As was calculated using the average
stream width (0.4 m) multiplied by the Reach 2 length (228 m).
During the entire study period, the streambed eroded a total of 142 cm (116.7 m3)
(Table 5.4). Estimates show that the melt rate was lowest on July 8, and greatest from
July 20-22. From July 8-13, the upper section of the streambed eroded a total of 37.6
65
cm/day (30.9 m3/day). Conversely, from July 13-24, the streambed eroded a total of
104.6 cm/day (85.8 m3/day).
Table 5.4. Measured stream erosion from the RTK survey and at the bottom depth sensor,
used for a combined total streambed erosion per day and amount of water equivalent
along the stream from July 8-24.
Date
(2014)
Streambed
Erosion (cm)
measured using
RTK
Streambed Erosion
(cm) measured at
Bottom Depth
Sensor
Combined
Streambed
Erosion
(cm/day)
Reach 2
Water
Equivalent
(m3/day)
July 8 34 4 3.3
July 9 30 6.8 5.9
July 10 6.8 5.6
July 11 6.8 5.6
July 12 6.8 5.6
July 13 2.8 5.3
97 60 3.6
July 14 9.4 7.7
July 15 9.4 7.7
July 16 9.4 7.7
July 17 9.4 7.7
July 18 9.4 7.7
July 19 9.4 7.7
July 20 10.5 8.6
July 21 10.5 8.6
July 22 10.5 8.6
July 23 5.5 6.8
11.2 2.8
July 24 8.4 6.9
Total 142 116.7
66
5.3.4 Sinuosity
The sinuosity index is expressed as a value greater than 1.0 and is calculated as:
𝑆𝑖 = 𝐶𝐿/ 𝐷𝐷 [Eq. 5.15]
where CL is channel length and DD is down-valley distance. The ratio allows for the
stream to be considered straight (Si <1.05) or meandering (Si ≤ 1.50) (Babar, 2005).
Using the stream profile collected with the GPS on July 13, the sinuosity index has been
calculated and is shown in Figure 5.20. The overall sinuosity for the stream was
Figure 5.20. Stream sinuosity. Arrows denote the 4 reaches.
67
calculated to be 1.07. Nevertheless, the sinuosity of the stream changed depending on the
reach. Reach 1 had the least sinuosity with a value of 1.04. The sinuosity for the Reach 2
was 1.09, with the upper half having a value of 1.08 and the lower half having a higher
value of 1.10. The sinuosity of Reach 3 was 1.05. Lastly, Reach 4 had the greatest
sinuosity with a value of 1.18.
5.3.5 Slope
Slope percentage was simply calculated by dividing the elevation change by the
distance downstream (Figure 5.21). The overall slope of the stream was 15%; however,
this was not consistent throughout the entire length. Reach 1 had a slope of 13%, while
Reach 2 had a slope of 10%. The upper and lower half of Reach 2 did not have the same
slope. The lower half of the reach had a larger slope than the upper half, with 11% and
9%, respectively. The bottom section of the stream was by far the steepest; Reach 3 had a
slope of 20%, while Reach 4 had a slope of 36%.
Figure 5.21. Stream slope. Note numbers and arrows denote stream reaches.
68
5.3.6 Step-pool Sequence
Using the RTK data from the stream profiles, the elevation was subtracted from
each subsequent survey point. A lip was identified at any location where the elevation
increased instead of decreased. By subtracting the elevation between the lip survey point
and the previous survey point, the pool depth was determined. The step height was then
calculated by subtracting the pool survey point by the preceding survey point (Figure
5.22).
Figure 5.22. Illustration of the stream features identified using the survey points.
Results show that there was a significant increase in step-pools between the July
13 and July 22-24 survey profiles (Table 5.5). The July 13 stream profile had a total of 5
step-pools; however, there were a total of 26 step-pools identified from the July 22-24
survey. On July 13, the pool depth average was 7.5 cm, while the step height was 39 cm.
The July 22 survey was taken of the upper half of Reach 2, and had an average pool
depth of 8.2 cm, and a step height of 17 cm. The July 23 survey was taken at the lower
half of the Reach 2, and had an average pool depth of 8.9 cm, and a step height of 30.2
cm. Interestingly, the average step height was 13 cm greater for the July 23 profile
69
Table 5.5. Step-pool characteristics including step height and pool depth on July 13 and
July 22-23, and water depth on July 23.
Pool ID Date Step Height (cm) Pool Depth (cm) Water Depth (cm)
1
July 13
23.9 4.6
2 44.9 10.3
3 25 8.8
4 47 7.5
5 54.6 6.3
6
July 22
11.7 12.7
7 23.1 0.8
8 24.3 15.8
9 10.7 7.2
10 6.1 5.6
11 27.1 22.3
12 3.2 2.5
13 26.9 7.1
14 21.8 5.1
15 19.3 4.5
16 21.5 13.6
17 8.4 1.2
18 11.7 3.6
19
July 23
17 1.9 15
20 24.6 15.3 20
21 40.2 21.7 30
22 73.4 1.5 20
23 19.1 1.1 18
24 21.5 13 22
25 46.2 3 27
26 25.6 5.5 11
27 18.8 6.6 16
28 5 10.6 15
29 50.8 13.5 25
30 20.3 18 31
31 56.2 9.1 16
70
survey. The water depth at each survey point was recorded on July 23. The average water
depth along the Reach 2 was 15 cm; however, within the pools the average was 20 cm,
with the highest water depth within a pool being 31 cm.
By creating a line out of the RTK survey points, the step-pools can be seen along
the stream profile in Figure 5.23. The location of the step-pools can be seen for the
profile on July 13 and the profile from July 22-24. On July 13, the majority of the step-
pools are farther downstream at the very end of Reach 2. The July 22-24 profile shows
that most of the step-pools were located in the bottom two-thirds of Reach 2.
Figure 5.23. Location of step-pools within the streambed on July 13 and July 22-24. Note
that 5 step-pools (1-5) existed in the streambed on July 13 and 26 (6-31) step-pools were
present on July 22-24. Note offset of vertical axes for separation and clarity.
5.3.7 Stream Temperature
Stream temperature was continuously measured; however, during the time the
sensor was frozen into the bed, the data was clearly incorrect and had to be disregarded.
71
Four manual temperature measurements were made using the velocity meter and
correspond quite accurately to the depth sensor data (Figure 5.24). From July 8-14, the
temperature of the stream varied on a daily cycle, ranging from 0.4oC to -0.3oC. On July
15, the stream temperature increases to a maximum of 2.4oC and fluctuated around 2oC
for the next two days. On July 17 and 18, the stream temperature decreased to 0.4oC and
no daily fluctuations were observed.
Figure 5.24. Stream temperature acquired from the top depth sensor from July 8-18.
Four manual measurements were taken during the study with the velocity meter.
5.3.8 Discharge
Unfortunately, results only include discharge from the top depth sensor, as the
bottom depth sensor was not retrieved from the stream. Figure 5.25 presents the raw
untransformed depth for the top depth sensor. Two calibration techniques were used to
acquire true depth measurements. A known issue with placing a depth sensor in a
supraglacial stream is the constant change in depth from streambed erosion. As a result, a
linear shift was given to all of the segments of data (segments were based on times
between streambed erosion measurements). Essentially, using the linear equation for each
line segment, the slope was set to 0. This removed the depth changes from erosion and
-0.5
0
0.5
1
1.5
2
2.5
07 08 09 10 11 12 13 14 15 16 17 18 19 20
Tem
pe
ratu
re (
oC
)
Day (July 2014)
Stream Temp Manual Measurement
72
left the true water depth variations. The second issue, was the freezing and thawing of the
depth sensor within the drilled hole. During the time when the depth sensor was frozen,
the sensor experienced an extreme change in pressure. To compensate for the change, the
frozen segments of data required a vertical shift downwards. As manual measurements in
depth were acquired on a frequent basis, these measurements could be used to ensure the
given calibrations were correct. Figure 5.26 shows the calibrated depth and manual depth
measurements.
Figure 5.25. Uncalibrated depth obtained from the top depth sensor from July 8-21. Note
the depth sensor was not in the stream from July 18, 16:00-July 19, 15:30.
Figure 5.26. Calibrated depth from the top depth sensor from July 8-21 and
corresponding manual measurements.
0
1000
2000
3000
4000
5000
6000
08 09 10 11 12 13 14 15 16 17 18 19 20
Un
calib
rate
d D
ep
th (
cm)
Day (July 2014)
0
2
4
6
8
10
12
14
08 09 10 11 12 13 14 15 16 17 18 19 20 21
De
pth
(cm
)
Day (July 2014)
Depth Manual Measurements
73
As depth and discharge are associated, a discharge rating curve was created using
the manual measurements obtained using the velocity-area method. Figure 5.27 clearly
illustrates a linear relationship between depth and discharge for the supraglacial stream.
The correlation between the depth and discharge was high with an R2 value of 0.95. With
this high correlation value and the equation on Figure 5.27, the depth measurements from
the top depth sensor could be transformed into discharge (Figure 5.27).
Figure 5.27. Relationship between measured discharge and measured depth.
Discharge could be calculated from July 8 and 21 (Figure 5.28). Data had to be
omitted or linearly infilled on three separate occasions, either because the depth sensor
was temporarily removed from the stream, or because the depth sensor was experiencing
a rapid change in pressure from freezing or thawing. The supraglacial stream discharge
ranged between 0.01-0.1 m3/s. During July 8-13, a daily cyclical pattern can be seen in
the discharge. This pattern is much less obvious from July 14-21. Evidently, a significant
change and reduction in discharge occurred between the first and second week of the
study.
y = 0.8711x - 0.0155R² = 0.9579
0.00
0.02
0.04
0.06
0.08
0.10
0.04 0.06 0.08 0.1 0.12 0.14
Dis
char
ge (
m3/s
)
Depth (m)
74
Figure 5.28. Discharge at the top depth sensor from July 8-21.
5.3.9 Froude number
The calculation of the Froude number (Fr) is an important aspect of the dynamics
of the supraglacial stream as it allows for the flow to be classified as subcritical (<1) or
supercritical (>1). The Froude number can be calculated as follows:
𝐹𝑟 = 𝑈/√𝑔𝐷0 [Eq. 5.16]
where U is the velocity (m/s), g is gravity (m/s2), and D0 is stream depth (m) (Parker,
1975; Karlstrom et al., 2013).
The Fr was calculated using the manually collected velocity and depth
measurements. As such, it should be noted that the Fr calculations are based on
instantaneous moments in time at only one location within the stream. Fr was calculated
for the top of the Reach 2 (at the top depth sensor location) and results are presented in
Table 5.6.
At the time when the measurements were made, the stream flow was always
supercritical as the Fr was constantly above 1. The average Fr during the entire study
0
0.02
0.04
0.06
0.08
0.1
08 09 10 11 12 13 14 15 16 17 18 19 20 21
Dis
char
ge (
m3 /
s)
Day (July 2014)
Discharge Linear Infill
75
period was 2.81 at the top of the Reach 2. The maximum Fr occurred on July 11, with
3.63 and the minimum was 2.17, on July 9.
Table 5.6. Froude number calculated for the top of Reach 2 from July 8-22.
Date / Time Top of Reach 2
July 8 15:15 2.20
July 9 12:07 2.17
July 9 14:30 3.11
July 9 16:05 2.89
July 11 12:08 3.38
July 11 14:45 3.63
July 13 17:40 3.32
July 14 15:15 3.22
July 16 12:44 2.57
July 17 13:40 2.66
July 17 17:43 2.22
July 19 16:17 2.42
July 20 20:08 2.73
July 22 10:15 2.86
5.3.10 Reynolds number
The Reynolds number is another important stream characteristic as it describes
whether the flow was laminar (<500) or turbulent (>2000). The Reynolds number (Re)
calculation for open channel flow is as follows:
Re = 𝑅ℎ x 𝑈
𝑣 [Eq. 5.17]
where U is the velocity (m/s), v is the kinematic viscosity (m2/s), and Rh is the hydraulic
radius (m). The Rh is the ratio of surface to its perimeter:
Rh = A / P [Eq. 5.18]
76
where A is the cross-sectional area (m2) and P is the wetted perimeter (m) (Camporeale
and Ridolfi, 2012). The cross-sectional area and perimeter are dependent on the channel
shape, which in this case was roughly U-shaped.
As seen in Table 5.7, the Re values are all much greater than 2000, suggesting
very turbulent flow conditions. During the study period, the average at the top of Reach 2
was 88x103, while the maximum was 14x104 on July 9, and the minimum was 56x103 on
July 13.
Table 5.7. Reynolds number calculated for the top of Reach 2 from July 8-22.
Date / Time Top of Reach 2
July 8 15:15 80x103
July 9 16:05 14x104
July 11 12:08 79x103
July 11 14:45 82x103
July 13 17:40 56x103
July 14 15:15 69x103
July 16 12:44 62x103
July 17 13:40 10x103
July 17 17:43 73x103
July 19 16:17 11x104
July 20 20:08 77x103
July 22 10:15 76x103
5.3.11 Pulsating
In the middle of the afternoon on July 16, fellow researchers observed an
intriguing phenomenon occurring at the location where the supraglacial stream flowed off
the front of the glacier, in the form of a waterfall. At this time, it was noted that the
77
waterfall was pulsating, where the discharge went from a maximum to 0, at regular
intervals of approximately 17 s.
A few hours after the original sighting of the pulsating phenomenon (July 16) the
discharge was manually measured at the top and bottom of Reach 2. The measurement at
the bottom of Reach 2 were taken at 17:30 and the measurement at the top of Reach 2
were taken at 17:48; 20 min apart. As seen in Figure 5.29, the average discharge at the
bottom of Reach 2 was 0.052 m3/s; with a maximum of 0.076 m3/s, and a minimum of
0.005 m3/s. Figure 5.30 shows that the average discharge at the top of Reach 2 was 0.041
m3/s, had a maximum of 0.048 m3/s, and a minimum of 0.036 m3/s. Interestingly, a rapid
change in discharge or pulsating was observed at the bottom of Reach 2, but not at the top
of Reach 2. At the bottom of Reach 2, the pulsating flow seems to be occurring at a
somewhat regular rate, with a sudden decrease occurring every 12-15 s.
Figure 5.29. Manual discharge taken at bottom of Reach 2 on July 16 at 17:30.
0
0.02
0.04
0.06
0.08
0 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80
Dis
char
ge (
m3/s
)
Time (s)
78
Figure 5.30. Manual discharge taken at the top of Reach 2 on July 16 at 17:48.
On July 18, a video of the pulsating waterfall was taken at 17:05 (Figure 5.31).
The discharge was recorded on a scale ranging from high to 0 every second for the
duration of the 70 s clip. A definite trend can be seen, where the discharge would be
extremely high for approximately 12 s and be low for 10-18 s.
Figure 5.31. Range in discharge estimated of the waterfall from a video taken on July 18
at 17:05.
On July 22, discharge was examined at the bottom of Reach 2 (Figure 5.32), in
between Reach 3 and 4 (at the stream bend) (Figure 5.33), and at the base of the waterfall
(Figure 5.34). The discharge recorded at the bottom of Reach 2 was taken at 10:35 by
0
0.01
0.02
0.03
0.04
0.05
0 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75 80
Dis
char
ge (
m3/s
)
Time (s)
0
1
2
3
0 5 10 15 20 25 30 35 40 45 50 55 60 65 70
Dis
char
ge
Time (s)
High
Low
Zero
Medium
79
manually recording velocity for 60 s. The discharge was extremely variable; ranging
between 0.077 m3/s and 0.041 m3/s; however, no trend was seen. Five minutes after the
measurements at the bottom of Reach 2, at 10:40, a video was taken in between Reach 3
and 4, which was 130 m farther downstream. Examination of the video shows that
pulsating was indeed occurring at regular intervals in this location. During the 90 s video,
four pulsating phases were observed. The discharge would be high or medium for 7-10 s
and suddenly drop to 0 for 8-15 s. At 11:20 another video was taken, this time at the
waterfall (230 m downstream from the bottom of Reach 2). In this 70 s video only two
pulsating phases were recorded. The discharge flowed for roughly 18 s followed by 10-
15+ s of 0 flow. The period between pulsating and high discharge clearly became
amplified with increased distance downstream.
Figure 5.32. Manual discharge taken at bottom of Reach 2 on July 22 at 10:35.
0
0.02
0.04
0.06
0.08
0 5 10 15 20 25 30 35 40 45 50 55 60
Dis
char
ge (
m3 /
s)
Time (s)
80
Figure 5.33. Discharge determined from a video taken between Reach 3 and 4 on July 22
at 10:40.
Figure 5.34. Discharge estimated at the waterfall from a video taken July 22 at 11:20.
Time-lapse imagery collected during the study shows the occurrence of the
pulsating phenomenon. However, due to limited visibility, only times with 0 discharge,
pulsating could be recorded. The imagery shows pulsating began on July 14 and took
place daily for the duration of the study (Figure 5.35). From July 14-17, the pulsating
often occurred between 22:00 and 5:00. On July 18, 19, 22 shorter erratic episodes were
observed during the day. No time-lapse imagery exists for part of July 20 and 21.
0
1
2
3
0 5 10 15 20 25 30 35 40 45 50 55 60 65 70 75
Dis
char
ge
Time (s)
0
1
2
3
0 5 10 15 20 25 30 35 40 45 50 55 60 65 70
Dis
char
ge
Time (s)
High
Medium
Low
Zero
High
Medium
Low
Zero
81
Figure 5.35. Pulsating occurrence from July 7-22.
07 08 09 10 11 12 13 14 15 16 17 18 19 20 21 22
Pu
lsat
ing
Eve
nt
Day (July 2014)
No Data
82
CHAPTER SIX: ANALYSIS
Results presented in the previous chapter suggest significant changes occurred
throughout the study period. The analysis examines the factors that affect the individual
changes to the glacier surface characteristics and the evolution of stream characteristics.
6.1 Changes to the Glacier Surface Characteristics
6.1.1 Ablation
During the entire study, the total averaged ice loss across the study area was 57.5
cm measured over 14 days. The average ablation was 4.2 cm/day from July 8-22, 2014.
Using photogrammetry and survey techniques, Whitehead (2013) estimated the ablation
rates on Fountain Glacier. He concluded that from June 17-20, 2009 ice loss was 3.3
cm/day, from June 30-July 6, 2010 ice loss was 5.5 cm/day, and from June 26-29, 2011
ice loss was 4 cm/day. Although the measurements were taken using different techniques,
the overall values correspond quite well.
Typically, net radiation and air temperature are the two main drivers of ablation
melt (Benn & Evans, 1998). As suspected, when the net radiation increased or decreased,
the ablation responded in a similar manner (Figure 6.1). The total daily ablation and max
daily net radiation were compared. Using Pearson’s r correlation coefficient a value of r =
0.65 was calculated. Figure 6.1 shows the ablation does not seem to be strongly impacted
by changes in air temperature. In fact, ablation and air temperature were only moderately
correlated with a value of r = 0.40. Dozier (1974) suggests that there is less ablation on
rainy days; therefore, ablation and relative humidity were plotted together in Figure 6.3.
Importantly, during high relative humidity, the ablation dropped anywhere from a quarter
to half. The ablation and relative humidity had a strong negative correlation of r = -0.68.
83
Figure 6.1. Relationship between ablation and net radiation.
Figure 6.2. Relationship between ablation and air temperature.
Figure 6.3. Relationship between ablation and relative humidity.
0
50
100
150
200
250
300
09 10 11 12 13 14 15 16 17 18 19 20 21 22 23
-50
50
150
250
350
Ab
lati
on
(m
3 /d
ay)
July (2014)
Ne
t R
adia
tio
n (
W m
2)
Ablation
Net Radiation
09 10 11 12 13 14 15 16 17 18 19 20 21 22
0
50
100
150
200
250
300
09 10 11 12 13 14 15 16 17 18 19 20 21 22
0
2
4
6
8
10
12
Ab
lati
on
(m
3 /d
ay)
July (2014)
Air
Te
mp
era
ture
(oC
)
Ablation Air Temperature
0
50
100
150
200
250
300
09 10 11 12 13 14 15 16 17 18 19 20 21 22
0
20
40
60
80
100
Ab
lati
on
(m
3/d
ay)
Day (July 2014)
Re
lati
ve H
um
idit
y (%
)
Ablation Relative Humidity Rain Event
84
6.1.2 Albedo
Due to the lack of an albedo reference, only relative albedo could be calculated.
As a result, the relative albedo values cannot be compared to real glacier albedo values.
However, daily variations in albedo can be compared to the literature. According to
Karlstrom et al. (2013), meltwater has the ability to influence changes in albedo. It is
believed that variations in the albedo during the first week were a result of varying
amounts of meltwater on the surface during the time of the albedo surveys.
Figure 6.4 shows the only meteorological variable that was correlated to albedo;
which was relative humidity at r = 0.44. On July 14, the relative humidity rapidly
increased and the albedo decreased. On July 16, the albedo values increased and
remained rather high in comparison to the first week. As Hock (2005) suggests, rain adds
water to the surface, which in turn decreases the albedo. This is clearly evident. Another
impact on the albedo is the cryoconite coverage. Redistribution of cryoconite by surface
washing during rain events is a phenomenon that has been widely documented (Brock et
al., 2006). In conclusion, multiple rain events washed the cryoconite off the surface and
increased the albedo.
Figure 6.4. Relationship between relative albedo and relative humidity.
01 02 03 04 05 06 07 08 09 10 11 12 13 14 15 160.5
0.55
0.6
0.65
0.7
0.75
0.8
07 08 09 10 11 12 13 14 15 16 17 18 19 20 21 22
0
20
40
60
80
100
Re
lati
ve A
lbe
do
Day (July 2014)
Re
lati
ve H
um
idit
y (%
)Albedo Relative Humidity Rain Event
85
6.1.3 Roughness
Information gleaned from images, the aerodynamic roughness survey, and direct
observation show that the glacier surface was extremely rough and littered with
cryoconite holes during the first week of the study. During this time, the holes grew
slightly each day, but overall the change in the glacier surface was small. In the middle of
the study, a significant change occurred in parallel with the multiple rain events. As more
rainfall occurred, the smoother the glacier surface became. Knighton (1972) suggested
that rain can cause the surface to melt and in turn causes a change in the permeability of
the ice surface. The rain essentially washes away the original weathering crust leaving
behind a non-permeable layer of ice (Figure 6.5). As a result, overland flow of water is
enhanced.
Figure 6.5. Depiction of glacier surface roughness change from rainfall.
The aerodynamic roughness length parameter was calculated to be 1x10-4 m on
July 22 and had a higher value of 6.7x10-4 m on July 23. Brock et al. (2006) has compiled
the aerodynamic roughness length results from a number of studies for high-latitude
glaciers and ice sheets. The July 22 value of 1x10-4 m corresponds with aerodynamic
roughness recorded in Greenland and a polythermal glacier in Sweden with smooth
86
surface ice. Values of 6x10-4-7x10-4 m have been seen for general polar glacier ice, which
is similar to the estimated value for July 23.
Research has shown that rain and low cloud inhibit the development of the
weathering crust (Stenborg, 1969). On July 22, the net radiation increased, the relative
humidity decreased, and the rain ceased (Figure 6.6). This explains the difference
between the values obtained from the two aerodynamic roughness surveys. On July 22,
the glacier surface was smooth; however, on July 23, the glacier surface had begun to
return to its previous 'sunny weather' roughness state and was closer to polar glacier ice
values for roughness.
Figure 6.6. Relationship between net radiation and relative humidity.
6.1.4 Changes to the Surface Characteristics and Effects on Supraglacial Stream
During the first week of the study, July 7-14, the variables that influenced the
glacier and supraglacial stream remained fairly constant. The net radiation and air
temperature varied on a daily cycle and the relative humidity remained relatively similar.
Ablation during the first week was high and generated large amounts of meltwater that
entered the supraglacial stream. The albedo varied depending on the amount of water that
0
20
40
60
80
100
-50
50
150
250
350
15 16 17 18 19 20 21 22 23
Re
lati
ve H
um
idit
y (%
)
Ne
t R
adia
tio
n (
W m
2)
Day (July 2014)
Net Radiation Rain Event Relative Humidity
87
was on the glacier surface; however, it remained quite low as the surface was covered in
dust. The surface was composed of a thick layer of weathering crust which was littered
with cryoconite holes. Water traveling at the surface had to travel through the permeable
weathering crust layer.
On July 14, the meteorological parameters caused a change in the glacier surface
characteristics. The relative humidity rapidly increased and the net radiation decreased.
The multiple rainfall events and cloudy conditions during the entire second week of the
study (July 14-22) caused the ablation amount to decrease and led to a significant
reduction in the meltwater available to enter the supraglacial stream. The glacier surface
albedo increased as the rainfall began to flush away the cryoconite. Eventually, much of
the cryoconite was washed off the surface, into the supraglacial stream. Lastly, the rain
also melted away the weathering crust, causing the glacier surface to become smooth and
therefore caused greater overland flow of water.
6.2 Evolution of Stream Characteristics
6.2.1 Stream Temperature
According to Isenko et al. (2005), field observations often show supraglacial
stream temperatures range between 0-0.4oC. During the first week, the temperature
ranged from -0.3-0.4oC. This is reasonably close to other observed temperatures, as the
accuracy of the depth sensor was ±0.05°C. During the second week of the study, the
temperatures were much higher than the normal supraglacial temperature range.
The two main meteorological variables generally responsible for changes in
stream temperature are net radiation and air temperature (Isenko et al., 2005). Figure 6.7
and Figure 6.8 show that during the first week of the study. On a diurnal basis, stream
88
temperature did indeed fluctuate with both net radiation and air temperature, with a
positive correlation of r = 0.62 and r = 0.49, respectively. Relative humidity had a
moderate negative correlation, with a value of r = -0.45; as when relative humidity
decreased, stream temperature increased (Figure 6.9).
Figure 6.7. Relationship between stream temperature and net radiation.
Figure 6.8. Relationship between stream temperature and air temperature.
-50
50
150
250
350
-1
0
1
2
3
07 08 09 10 11 12 13 14 15 16 17 18 19 20
Ne
t R
adia
tio
n (
W m
2)
Stre
am T
em
pe
ratu
re (
oC
)
Day (July 2014)
Stream Temp Net Radiation
0
2
4
6
8
10
12
-1
0
1
2
3
07 08 09 10 11 12 13 14 15 16 17 18 19 20A
ir T
em
pe
ratu
re (
oC
)
Stre
am T
em
pe
ratu
re (
oC
)
Day (July 2014)
Stream Temp Air Temp
89
Figure 6.9. Relationship between stream temperature and relative humidity.
Using the maximum daily values of the following: net radiation, air temperature,
relative humidity, and stream temperature over the course of the entire study, the
correlations were analyzed. Air temperature did not have a significant correlation because
the stream temperature often hit a maximum of 0.4oC despite changes in the air
temperature. The net radiation had a strong negative correlation of r = -0.70; for when net
radiation decreased, stream temperature increased. Lastly, the relative humidity had a
correlation of r = 0.54 because when the rain events occurred during the second week, the
stream temperature also increased. It is assumed that during all the rain events the stream
temperature increased.
On July 17-18, the stream temperature remained at a constant high (without the
heat of rain) of 0.4°C. This was evidently due to the spike in net radiation on July 17 and
the elevated air temperature the second half of July 17 and on July 18.
6.2.2 Streambed Erosion
In a study by Marston (1983), it was noted that a temperature of 0.005-0.01oC can
account for incision between 3.8-5.8 cm/day, while observed erosion rates generally
40
50
60
70
80
90
100
-1-0.5
00.5
11.5
22.5
33.5
07 08 09 10 11 12 13 14 15 16 17 18 19 20
Re
lati
ve H
um
idit
y (%
)
Stre
am T
em
pe
ratu
re (
oC
)
Day (July 2014)
Stream Temp Rain Event Relitive Humidity
90
range from 4-8 cm/day. During the first week of this study, the stream temperature was
between -0.3-0.4oC and the erosion was between expected values with 4-6.8 cm/day.
Figure 6.10 shows relatively high net radiation and Figure 6.11 shows that relatively low
relative humidity occurred during the first week. Normally, solar radiation penetrates
through the flowing water, causing melt of the channel walls and sensible heat exchange
with the air, resulting in warming of the water. This explains the normal rates of erosion.
During the second week of the study, the streambed erosion rates were extremely
high and ranged between 9.4 -10.5 cm/day. Also during this time, the net radiation
decreased and the relative humidity increased. Although normally erosion occurs because
of melt from net radiation, it is obvious that increased rates of erosion occurred during
rainfall. Karlstrom et al. (2013), suggests that rainwater can increase the temperature of
the stream and thermally erode the streambed.
Figure 6.10. Relationship between streambed erosion and net radiation.
08 09 10 11 12 13 14 15 16 17 18 19 20 21 22
0
2
4
6
8
10
12
08 09 10 11 12 13 14 15 16 17 18 19 20 21 22 23
-50
50
150
250
350
Stre
amb
ed
Ero
sio
n (
cm/d
ay)
Day (July 2014)
Ne
t R
adia
tio
n (
W m
2)
Streambed Erosion Net Radiation
91
Figure 6.11. Relationship between streambed erosion and relative humidity.
6.2.3 Discharge
On a diurnal time scale, the discharge is strongly influenced by the net radiation
and is affected by the daily weather; specifically cloud cover. Dozier (1974), suggested
that there is often a 2-4 hour lag between peak daily discharge and peak downward
radiation. Figure 6.12 displays the relationship between discharge and net radiation.
Evidently, the two variables are related, but the lag time appears to be rather inconsistent.
On July 8, the maximum daily net radiation occurred around 14:00, with the maximum
daily discharge actually occurring an hour before. On July 9, both the peak discharge and
peak net radiation occurred around 13:00. On all other days, the maximum daily net
radiation peaked between 11:00-13:00, 30 mins-4 hours before the maximum daily
discharge. Discharge and relative humidity were also plotted together, which can be seen
in Figure 6.13. It can be seen that when the relative humidity increases, the discharge
decreases.
Given the change in lag, a correlation value is difficult to determine on an hourly
time scale. Using the maximum daily: discharge, net radiation, and relative humidity, net
08 09 10 11 12 13 14 15 16 17 18 19 20 21 22
0
2
4
6
8
10
12
08 09 10 11 12 13 14 15 16 17 18 19 20 21
0
20
40
60
80
100St
ream
be
d E
rosi
on
(cm
/day
)
Day (July 2014)
Re
lati
ve H
um
idit
y (%
)
Streambed Erosion Rain Event Relative Humidity
92
radiation was found to have a correlation of r = 0.49 and relative humidity had a
correlation of r = -0.49 with discharge. Undoubtedly, the change in net radiation and
relative humidity/ rainfall events had a large impact on the overall discharge.
Figure 6.12. Relationship between discharge and net radiation.
Figure 6.13. Relationship between discharge and relative humidity.
6.2.4 Step-pool Sequence
The July 13 stream profile had a total of 5 step-pools; however, there were a total
of 26 step-pools in the stream a week later. The rapid development within the streambed
was evidently connected to the change that occurred in the meteorological parameters at
the same time.
-50
0
50
100
150
200
250
300
350
0
0.02
0.04
0.06
0.08
0.1
08 09 10 11 12 13 14 15 16 17 18 19 20 21
Ne
t R
adia
tio
n (
W m
2 )
Dis
char
ge (
m3 /
s)
Day (July 2014)
Discharge Net Radiation
0
20
40
60
80
100
0
0.02
0.04
0.06
0.08
0.1
08 09 10 11 12 13 14 15 16 17 18 19 20 21R
ela
tive
Hu
mid
ity
(%)
Dis
char
ge (
m3 /
s)
Day (July 2014)
Discharge Rain Event Relative Humidity
93
6.2.5 Pulsating
Although pulsating has been observed in supraglacial streams in the past, no
previous connections between meteorological parameters have been identified. An
analysis of the collected pulsating data shows a link between two sources of
meteorological data. First and foremost, as seen in Figure 6.14, the pulsating events
clearly began in response to an increase in the relative humidity and the rain events.
Using maximum daily relative humidity and daily occurrence of pulsating, the correlation
was found to be strong with a value of r = 0.75. Although the days the rain events and
pulsating occurred were correlated, the exact timing of the pulsating and rainfall events
were not.
Figure 6.14. Relationship between pulsating events and relative humidity.
As noted in the results section, from July 14-17, the pulsating often occurred
between 22:00 and 5:00 and on July 18, 19, and 22 shorter erratic episodes were observed
during the daytime. As the timing of the pulsating seems to be related to time of day, net
radiation and pulsating events were plotted in Figure 6.15 and a correlation value was
calculated. As seen in the figure, pulsating very clearly occurs during low net radiation
40
50
60
70
80
90
100
07 08 09 10 11 12 13 14 15 16 17 18 19 20 21 22 23
1
1.1
1.2
1.3
1.4
1.5
1.6
1.7
1.8
1.9
2
Day (July 2014)
Re
lati
ve H
um
dit
y (%
)
Pulse Event No Data Rain Event Relitive Humidity
94
and the correlation value confirmations that net radiation and pulsating had a strong
negative correlation of r = -0.79. Although net radiation is the meteorological parameter
connected to the timing of the pulsating events, the true influence was stream discharge.
As discussed above, discharge alternated with net radiation.
Figure 6.15. Relationship between pulsating event and net radiation.
Figure 6.16 illustrates the relationship between discharge and pulsating, which
had a correlation value of r = -0.49. The correlation between discharge and pulsating was
slightly less than the correlation between net radiation and pulsating, presumably due to
the missing discharge data.
Figure 6.16. Relationship between pulsating event and discharge.
-50
50
150
250
350
14 15 16 17 18 19 20 21 22 23
-0.5
1.5
3.5
5.5
7.5
9.5
Day (July 2014)
Ne
t R
adia
tio
n (
W m
2 )
Pulsing Event No Data Net Radiation
0
0.5
1
1.5
2
2.5
0
0.02
0.04
0.06
0.08
14 15 16 17 18 19 20 21
Dis
char
ge (
m3 /
s)
Day (July 2014)
Discharge Pulsating Event No Data
95
6.2.6 Summary of Changes and Stream Characteristic Connections
In summary, the stream was influenced by the rainfall and cloudy conditions. The
stream temperature increased from the warm rainwater that was added to the stream.
Under the influence of the rainfall and increased stream temperature, the rate of
streambed erosion rapidly increased. Discharge decreased as the net radiation was
blocked by the rain clouds. In addition, the streambed morphology was significantly
altered as step-pools rapidly formed during the week of rain. Lastly, pulsating very
clearly began in parallel with the beginning of the rainfall events.
The increase in streambed erosion within Reach 2 was not only impacted by
meteorological parameters, but also affected by a combination of thermal and mechanical
influences that also altered during the second week. According to Moore (1991), there is
often a close relationship between air temperature and rain temperature. This assumption
corresponds to direct observations during the study. Given that the air (rain) temperature
was between 4-8oC and the stream was originally just around 0oC, the rain would have
transferred heat to the stream. The mixing of warm rain water into the supraglacial stream
would have been enhanced by the raindrops and turbulent flow (Moore, 1991).
As noted in the surface characteristic section, the cryoconite on the glacier surface
was washed into the stream. According to Isenko et al. (2005), stream sediments can also
influence the water temperature. The albedo of the glacier increased as a result of the
sediment being washed off of it. The introduction of the sediment into the stream
facilitated additional thermal and mechanical erosion. This erosion takes place along the
boundary between water and ice, as suspended sediment is transported downstream and
smashed into the ice walls by the turbulent flow (Ferguson, 1973; Knighton, 1985).
96
Therefore, during the second week, there was more available energy to melt/erode the
channel perimeter.
97
CHAPTER SEVEN: DISCUSSION
Chapter seven is an examination of the connections between the changes in surface
characteristic and stream characteristic. The development of the step-pool sequence and
the pulsating phenomena will also be discussed. Finally, a discussion of uncertainty is
presented.
7.1 Surface Characteristic and Stream Characteristic Connections
7.1.1 Watershed Runoff Contributions
The stream runoff contributions for Reach 2 is derived from three sources: snow
or ice melt from the glacier surface, melt from the boundary of the channel itself, and
rainfall during the ablation season (Dozier, 1974; Marston, 1983). Figure 7.1 shows the
breakdown of the three daily sources of stream runoff within the Reach 2 watershed. The
melt from ablation was by far the highest of the three sources with 94.5% of the total
watershed water balance, the streambed erosion accounted for 4%, while the rainfall only
made up 1.5%. Although there were multiple rainfall events, the total amount of rain was
fairly low. Unlike alluvial rivers where rainfall would be a major contributor, the rainfall
would have to be excedingly high, much closer to the daily amount of ablation, to
become important for the total stream runoff.
98
Figure 7.1. Influences of Reach 2 watershed stream runoff contributions.
Figure 7.2 is a comparison of the discharge and calculated Reach 2 watershed
stream runoff. However, as seen in the figure the Reach 2 watershed runoff and stream
discharge increase correspond. The R2 correlation value is 0.38. As additional water
from further upstream was already present in the stream when discharge was calculated a
direct comparison is not possible.
Figure 7.2. Comparison of Reach 2 watershed stream runoff and stream discharge.
0
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7.1.2 Stream Incision
Supraglacial streams exist in areas where the stream erosion is greater than the
rate of glacier ablation, causing vertical incision into the glacier (Knighton, 1981;
Marston, 1983). Stream incision (SI) (cm/day) was calculated as follows:
SI = ADA - SE [Eq. 7.1]
where ADA is the average measured ablation of the glacier surface across the watershed
(cm/day) and SE is the stream erosion (cm/day). Note prior to the calculation both ADA
and SE were transferred into negative values to represent ice loss.
Marston's work (1983) in the Juneau Icefield suggests that with erosion rates
ranging from 4-8 cm/day, that the glacier surface ablation rates would then range from 1-
4 cm/day. Figure 7.3 shows that from July 9-13 the stream erosion was only slightly
greater than the ablation, with an average stream incision of -1 cm. From July 14-22, the
stream erosion was much greater than the ablation, leading to an average stream incision
of -6.4 cm.
Figure 7.3. Stream incision.
-12
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100
Figure 7.4 illustrates the conceptual transformation of the stream incision before
and after the rainfall events. From July 9-13, the glacier surface melted a total of 22 cm
and the stream eroded 30 cm. During this time, the stream erosion was only 8 cm greater
than the surface melt. From July 13-22, the total ablation was 23 cm, while the stream
erosion was 97 cm. Remarkably, during the week of rainfall the stream incised a total of
74 cm into the glacier surface.
Figure 7.4. Cross section of stream watershed illustrating the streambed incision before
the rainfall (July 9-13) and after multiple rainfall events (July 13-22).
Images of the stream were taken in approximately the same location on July 9, 16,
and 22 (Figure 7.5). In the image taken on July 9, the water level of the stream was only
slightly lower than the bank. The stream channel in the July 16 image appeared fairly
similar to that taken on July 9. The water level in relation to the bank height only looks
slightly deeper. The image taken on July 22 shows the change after the rainfall. A deep
and narrow stream channel had formed, and the water level was much below the surface
of the glacier.
101
Figure 7.5. Images of the stream incision all within 3 m of the bottom depth sensor;
where A) shows the stream fairly level with the bank; B) shows a slight incision of the
stream; and C) shows significant incision as the water level is much below the surface.
During the rainfall, the stream incision was much greater; this caused an enhanced
development and deepening of the stream. Importantly, a feedback exists where, as the
stream incises deeper into the glacier surface, each meander will sweep past a given point
at a slightly lower level (Moody et al., 2003). As the stream propagates farther up-glacier,
this increases the watershed area, and further facilitates the growth of the stream (Gabler
et al., 1999). In the long term, the stream maybe influenced by climatic changes. If the
climate shifts to more humid conditions or rainfall increases, the erosional impact can
cause channels to become narrower and deeper (Waugh, 2000).
7.2 Development of Step-pool Sequence and Pulsating Phenomenon
As previously discussed, step-pools have been observed in mountainous, bedrock,
arid, and supraglacial streams, and are a common element of fluvial environments
(Knighton, 1998). Although these formations occur in a number of environments, the
mechanics allowing them to form are quite different. Step-pools in alluvial rivers develop
during extreme floods; however, discharge decreased during the formation of step-pools
102
in this supraglacial stream. As supraglacial streams have strongly diurnal discharge
patterns (Marston, 1983), the environment is adapted to large and rapid changes in
discharge. It is hypothesised that for step-pools to form in a supraglacial stream as a
result of high discharge, rainfall amounts would have to far exceed the normal discharge
fluctuations that occur. This did not occur in the studied stream.
Vatne and Refsnes (2003) assumed that step-pools also form in glacier meltwater
streams during high discharge. This is because frictional heat is related to the volume of
water in the stream (Knighton, 1972), so at high discharge step-pools could form from the
added energy in the form of heat. In this case, results and analysis show that discharge
decreased; nevertheless, the assumption that step-pools are formed from the added energy
of heat could be viable.
7.2.1 Formation of the Step-pools
Knighton (1981) noted that pools tend to form immediately downstream of
meander bends. It has been suggested that as water travels around a stream meander, a
helical flow pattern develops causing a pool to be eroded directly after the stream
meander. An erosional pool is formed directly underneath a step from the added energy
in the waterfall, and a promontory (or lip) forms directly downstream from the pool due
to the dispersion of energy within the pool. The lip forms a reverse bed slope within the
stream profile, and thus forces water to travel against gravity to travel farther downstream
(Vatne & Refsnes, 2003).
The calculated Reynolds Numbers show that flow conditions were always
turbulent and that helical flow was observed in the stream throughout the study.
Nevertheless, only 5 step-pools were observed in Reach 2 during the first week of the
103
study. Although overall discharge was significantly less during the second week of the
study, local increases in discharge occurred from the pulsating phenomenon. Figure 7.6A
shows the very beginning of a pulse and the helical flow around the stream meander bend
can be seen.
Figure 7.6. Helical flow step-pool formation; where A) is the beginning of a pulse with
helical flow occurring around the stream bend; B) was taken a few seconds later during
the middle of a pulse. Increased local discharge can be seen and the water crashes down
directly below the stream bend; C) was again taken a few seconds later, in between
pulses events. It can be seen where the helical flow created the pool and subsequent lip.
Images were taken on July 22 at 10:40 at the major stream bend.
Figure 7.6B was taken a few seconds later during the middle of a pulse. The
discharge is very high and the water is crashing down directly below the meander bend.
The fluctuation in discharge, as well as the added thermal energy from the stream
temperature increase, and mechanical energy from the suspended sediment are the
hypothesized mechanisms that allowed for the helical flow to rapidly form the pools and
104
subsequent lips within the streambed (Figure 7.6C). The images were taken at the major
stream bend. The slope is extremely steep and this part of the channel was very sinuous.
Figure 7.7 shows an aerial view of the stream with the locations of the 26 pools
that formed in Reach 2 during July 13-24. The image shows only about half the pools are
located immediately downstream of a stream meander bend. Dozier (1974) suggested that
structural elements such as shear planes normal to the stream result in knickpoints or
steps. Figure 7.7 does show a clear correlation with some of the shear planes (or in this
case transverse fractures). However, the presence of structural elements on their own
does not explain the rapid development of the step-pools.
Figure 7.7. Step-pool locations on aerial view of the stream showing the correlation of
pools with stream meander bends and transverse fractures.
A second mechanism for step-pool formation in relation to the glacier structure
and in connection with the multiple rainfall events has been hypothesized. On July 20, the
transverse fractures on the north side of Fountain Glacier formed a hydrological
connection with the bed and sediment-laden water flowed from the fractures (Hambrey,
105
pers. comm.). This hydrological event continued on and off from July 20 until July 22.
Figure 7.8A displays two of the locations where this event took place. Figure 7.8B shows
the emergence of water from one of the transverse fractures. It is hypothesized that as
rainfall accelerated the recharge of the subglacial hydrological system, the hydrostatic
pressure within the glacier increased. This resulted in artesian flow at the surface and
lubrication of the shear plane fractures.
Figure 7.8. Image of a hydrological event on taken on July 20 where A) shows two
locations that sediment-laden water was traveling from the base of the glacier through
the transverse fractures onto the surface; and B) shows is close up image of the water
flowing from a transverse fracture.
It is believed that as more rainfall occurred, water traveled to the base of the
glacier, and caused the pieziometric surface to rise. It is hypothesized that the transverse
fractures became lubricated and caused slip between the fractures (Figure 7.9). At the
locations where the transverse fractures crossed the stream, a step formed, which in turn
also created an erosional plunge pool (Figure 7.10). The image was taken on July 22
106
within Reach 2. It should be noted that no meander bend occurred just before the step-
pool and at this location the slope was moderate.
Figure 7.9. Structural step-pool formation from glacial slip events.
Figure 7.10. Transverse fracture formed step-pool. Image was taken on July 22 within
Reach 2.
Upon further review of the two step-pool formation mechanisms, the
environments (again seen in Figure 7.6 and Figure 7.10) appear to be rather different.
Within the stream, the step-pools formed from both transverse fractures and helical flow.
107
The transverse fractures step-pools formed in straight areas, while the helical flow step-
pools occurred in locations with a meander bend. Very few transverse fractures were
located in Reach 1 and they occurred irregularly in Reaches 2-4. A greater number of
meander bend step-pools were located in the lower half of Reach 2 in comparison to the
top, and although Reach 3 and 4 were not surveyed, the majority of the step-pools within
the stream were actually observed within this region. The difference in the number of
step-pools formed by helical flow can be attributed to the change in slope and sinuosity.
The lower half of Reach 2 had a greater slope by 2%, presumably leading to a greater
velocity, and a 0.02 higher sinuosity value (more meander bends). It is not surprising
then, that Reach 4 had the most step-pools as the slope was 25% greater and the sinuosity
was 0.08 higher than the lower half of Reach 2.
As illustrated in the results section, the pool depth and step height increased
between July 22 and 23. The data suggests that the extra day of rainfall allowed for
further pool evolution. It is also possible that helical flow was the cause of larger steps
and deeper pools due to the more sinuous conditions in the lower half of Reach 2. In fact,
both of these reasons likely explain the difference in step height and pool depth.
7.2.2 Pulsating Phenomena
Despite the step-pool formation mechanism, the step is often convex in formation
and characteristically is a narrower and shallower section within the channel. At the
junction between a step convexity and pool, the flow immediately changes from
supercritical to subcritical, causing the formation of a hydraulic jump from the step
(Knighton, 1981). As changes in flow behaviour occur within these sequences, there is
108
also a local change in the Froude number with the value decreasing between the straight
and meandering reaches.
Roll-waves form in locations where flow is unstable due to very high velocity
and/or very steep gradients. When the velocity of the wave is greater than that of the
surface flow adjacent to the channel bed and sides, a breaking wave forms (Carver et al.,
1994). Supraglacial channels are characterized by increased flow velocity largely due to
the restraints of the cross‐sectional area (Knighton, 1981). Roll-waves are strongly three
dimensional in nature (Carver et al., 1994) and the liquid-solid interface in supraglacial
streams allows for a free surface flow and facilitates water to travel in a three
dimensional oscillation (Camporeale & Ridolfi, 2012). The combination of rapid
increases in velocity and the liquid-solid interface of supraglacial streams allows for a
good environment for roll-waves to form.
Carver et al. (1994) suggests that in a supraglacial stream, a Froude number
greater than 2 would allow for the formation of a roll-wave. Results show that although
discharge decreased the second week of the study, the velocity was still high, and the
Froude number was always greater than 2. When instability occurs, the formation of a
series of roll-waves can breakdown into traveling waves or pulsating flow (Carver et al.,
1994). It is hypothesised that this occurred within the studied stream; where a roll-wave
would form upstream and travel downstream as a traveling wave.
Figure 7.11 depicts the transformation of the streambed and change in flow
patterns on July 13 (the day before the rain began), and July 23 (after it had been raining
for a week). Figure 7.11B-E illustrates the pulsating phenomena in relation to the step-
pools. In Figure 7.11A, the water in the stream flowed normally as no step-pools existed
109
within the channel. With the rapid formation of the step-pools, the flow pattern changed.
The water would enter a pool (Figure 7.11B), slush backwards (Figure 7.11C), and then
forwards (Figure 7.11D). Importantly, if the water was slushing forwards as another
traveling wave entered the pool, constructive interference would occur (Figure 7.11E). If
the timing was not synchronized destructive inference would occur, and no pulsating
transpired. As seen in Figure 7.11F, the traveling wave (or beginning of the pulse) had a
turbulent rounded front with a smooth tail section. During the pulse, the discharge was
high, but in between pulses, the channel was either completely empty or only minimal
amounts of water traveled through. After that the pulse abated, the next wave would
come along in somewhat regular intervals.
Figure 7.11. Pulsating phenomena where A) shows the normal stream conditions; B-D)
shows the water sloshing back and forth within the step-pool; E) shows the constructive
interference; and F) shows the traveling waves.
110
As discussed, pulsating flow was observed and was known to begin within Reach
2. This is because very few, if any, step-pools formed in Reach 1. Within Reach 2, the
slope and sinuosity increased, and the transverse fractures began, allowing for the
formation of step-pools. Van der Meer (2004) believed that the pulsating phenomena he
witnessed began at a slight widening of the streambed. In this case, the traveling wave
was probably initiated in one of the top step-pools, but the initial location was never
identified.
The pulsating phenomenon within supraglacial streams has only been noted in the
literature on a few occasions. Carver et al. (1994) observed pulsating where a discharge
of 0 was measured for a 6-7 s period, and Knighton (1981) witnessed a well-defined
wave period of 9.1 s, but stated that the pools within the streambed would lengthen the
passage of waves and successive waves would amalgamate in the downstream sections.
Lastly, van der Meer (2004) documented a pulsating period of 8-12 s. He noted that
within the streambed, numerous transverse ribs caused the emergence of a sinusoidal
wave that increased in amplitude farther downstream.
Results from this study did show a somewhat sinusoidal wave pattern; however,
Figure 7.12. Depiction of the wave as the amplitude and period increased downstream.
111
the wave grew in size from the amalgamation of the waves ( Figure 7.12). As the
pulsation traveled farther downstream, the period between waves and the amplitude of
the wave increased. This was potentially caused by the pools, which acted like temporary
stores. The farther downstream the more pools existed the more water could be stored.
Unlike previous studies, the actual change in period and amplitude changes
downstream were determined. On July 22, within a 230 m stream section, the discharge
at the bottom depth sensor had small fluctuations with no 0 discharge observed. At the
stream bend (approximately halfway between the bottom depth sensor and waterfall), the
discharge was high for 7-10 s, then suddenly dropped to 0 for 8-15 s, to the discharge at
the waterfall flowing for roughly 18 s followed by 10-15+ s of 0 flow. Fundamentally,
within a 60 s timeframe and 100 m section, the amplitude and period had increased to the
point that 1 less full pulsation cycle was observed downstream. The step-pools were
believed to evolve over the course of the study. As the values were taken on the last day
of the study, they represent a snap-shot in time for the most developed step-pool stage
observed.
The exact development of the step-pools was not monitored regularly and as such
the evolutionary time-scale of the step-pools in connection to rainfall amounts is not fully
understood. Nonetheless, the pulsating phenomenon may give some insight into the step-
pool evolution. Pulsating often occurred between 22:00 and 5:00 from July 14-17, but on
July 18, 19, and 22 shorter erratic episodes were observed during the day. It can be
suggested that pools would have deepened during the study, specifically on July 18, 19,
and 22. This is believed as during the day more discharge occurred, yet pulsating was
occurring during this time.
112
7.3 Discussion of Uncertainty
The discussion of uncertainty includes: meteorology, ablation, albedo, roughness,
stream temperature, discharge, stream position/step-pool sequence/streambed erosion,
and pulsating. This section is important as it both discusses and quantifies some of the
uncertainty associated with the presented results.
7.3.1 Meteorology
The first potential uncertainty in meteorological data was the collection of air
temperature, relative humidity, and net radiation. While there are four AWS on Bylot
Island, data from only Bylot-1 and Bylot-3 were appropriate for analysis in this study.
Bylot-1 is located at the base of Fountain Glacier and is the closest weather station to the
field site. Unfortunately, data from Bylot-1 was only retrieved between July 15 and July
24. Nevertheless, it can be argued that Bylot-3 provides a better representation of the
conditions on Fountain Glacier, due to its position on glacier ice.
The second meteorological data source was a precipitation/evaporation pan
temporarily installed in the centre of the study area, on the surface of Fountain Glacier.
The depth was recorded to the nearest mm, resulting in an uncertainty of 0.5 mm. On July
13 and 23 evaporation data was collected and evaporation was also estimated using
Ohno’s equation; to the nearest mm the values are the same. On all the days when the net
water balance was negative, the error in evaporation was assumed to be 0.5 mm or less
(Figure 7.13).
Although the majority of the precipitation values were obtained from Pond Inlet,
which was 35 km away. The combination of manually recorded precipitation values
from: the gauge, time-lapse images determining the time of rainfall events, and the
113
change in relative humidity correspond quite well (except on July 21 when it was noted
that we were being rained on and Pond Inlet looked wonderfully sunny). From July 14-16
both rain/evaporation data on Fountain Glacier and rain data were obtained from Pond
Inlet. The measured rain/evaporation was 9 mm. Pond Inlet received 11.4 mm of rain, but
when the 3 mm of evaporation was subtracted, the rainfall calculated for Fountain Glacier
was approximately 9.4 mm. This results in 0.4 mm of uncertainty. On the days the net
water balance was positive, the uncertainty was suspected to be 1 mm or less (Figure
7.13).
Figure 7.13. Net water balance with 0.5 mm of uncertainty for negative values
(evaporation) and 1 mm of uncertainty for positive values (rainfall).
7.3.2 Ablation
There are a number of issues with the collection and processing of ablation data.
Known errors associated with ablation measurements are from the formation of ablation
hollows surrounding the embedded ablation stake (Konzelmann & Braithwaite, 1995).
On occasion the stakes were found to be tilted within the hole due to the top of the hole
being wider than the bottom. In all instances the stakes were placed vertically before
07 08 09 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24
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114
measurement to ensure consistency. The diameter of the hole varies depending on the
colour, size, and material of the stake. The installed stakes were white in colour to ensure
no further melting feedback from albedo. In addition, the stakes were only 0.6 cm
diameter and composed of wood to mitigate the formation of a hollow and weight of the
stake pushing farther into the ice. Another concern when taking measurements is the
slope of the glacier causing one side of the stake to be longer than the other. To diminish
this error, the distance from the top of the stake to the ice surface was always measured
on the up-glacier side of the stake. Lastly, another concern regarding the ablation
measurements is the in consist time of measurement. Ideally, measurements would have
been taken daily at the same time each day. However, this was not possible. To allow for
accurate comparisons between days, the data was divided into average daily ablation
based on 24 hours in a day. Although the exact uncertainty associated with dividing the
data in this manner has not been quantified, the overall trend; where the ablation was
higher during the first week, and lower during the second week is definite.
7.3.3 Albedo
Methodological and meteorological uncertainties during data collection can be
caused from variations in camera position, lighting, shadowing, and surface aspect
(Irvine-Fynn et al., 2010). Firstly, care was used to ensure pictures were taken at a similar
height and angle. Secondly, in the middle of the July 23 albedo survey, cloud cover
caused a visual change in the images (Figure 7.14). All six images were taken within 5
min and 20 m of each other. During calibration, changes in lighting were taken into
account. As the albedo values remain consistent and cloud cover does not appear to
impact on the albedo values, the calibration was considered successful.
115
Figure 7.14. Albedo values under differing cloud conditions taken during the July 23
albedo survey a few min apart.
The average albedo in the SW, NW, NE, and SE directions was calculated to
determine if there was a statistical difference based on direction (Table 7.1). On July 16,
the average in the NW and NE directions were the highest with 0.72, and the lowest in
the SE and SW direction with 0.71. On July 22, the average was greatest in the NE
direction with 0.78, and lowest in the SW with 0.72. Lastly, on July 23, the NW had the
highest relative albedo of 0.62, while the SW, NE, and SE had the same albedo of 0.60.
On July 16, 22, and 23, the range between the albedo in the four directions was 0.01,
0.07, and 0.02, respectively. No statistical difference was found in the four directions.
Table 7.1. Comparison between relative albedo in various directions for July 16, 22, and
23.
Direction July 16 July 22 July 23
SW 0.71 0.72 0.60
NW 0.72 0.73 0.62
NE 0.72 0.78 0.60
SE 0.71 0.75 0.60
116
The overall uncertainty of the relative albedo measurements can be seen in Figure
7.15. On July 16, 22, and 23, the albedo was estimated using both the images from the
albedo and ablation survey. On July 16, 22 and 23, the difference was 0.03, 0.01, and
0.01, whereas the results from the ablation survey were always higher. This difference is
most likely due to the presence of the ablation stake within the picture and the lack of
crevasses in any of the ablation images. On the days that only the ablation survey photos
were used to determine albedo the uncertainty was believed to be 0.03 or less. On the
days that only 1 or a few images were used to determine the albedo, the uncertainty value
has been doubled to 0.06. Lastly, on the days that albedo was inferred from between
days, the uncertainty is anywhere between the 2 values.
Figure 7.15. Relative albedo uncertainty.
7.3.4 Roughness
A number of studies have been done to determine the optimal length and
measurement interval to use for a glacier micro-topographic survey. Brock et al. (2006)
and Irvine-Fynn et al. (2014) suggested a 5 m long profile with 10 cm intervals captures
the scales of change. On Fountain Glacier, this scale was found to be appropriate as it
07 08 09 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24
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117
captured the glacier variability and small scale changes (Irvine-Fynn, pers. comm.). It
should be noted that as cryoconite holes are 1-5 cm in width, this technique captures
some of the cryoconite holes, but is unable to decipher the size or shape of each
individual hole. The roughness measurements between the ice surface and horizontal line
were taken to the nearest mm, but according to Irvine-Fynn et al. (2014), using this
technique will result in an accuracy of ±2.5 mm. An example of the 2.5 mm uncertainty
in relation to the roughness measurements from July 22 can be seen on Figure 7.16.
Figure 7.16. Roughness measurements for July 22 with 2.5 mm uncertainty.
7.3.5 Stream Temperature
Stream temperature measurements were taken continuously by the depth sensor
and point measurements were taken with the velocity meter. Two of the four point
measurements corresponded perfectly with the depth sensor temperature, while the other
two were within 0.2oC. As portions of the temperature data had to be disregarded due to
the sensor being frozen into the streambed, the velocity meter measurements confirm the
accuracy of the remaining sections. The difference between the depth sensor and velocity
meter is explained by the accuracy, which was ±0.2°C for the Flowatch velocity meter,
and ±0.05°C for the Levelogger Edge depth sensor. Figure 7.17 shows stream
6
8
10
12
14
16
0 100 200 300 400 500
Surf
ace
Ro
ugh
ne
ss (
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Profile Length (cm)
118
temperature, the point measurements, and error bars of 0.05°C (which are so small they
cannot be seen).
Figure 7.17. Stream temperature with manual point measurements and 0.05oC error
bars.
7.3.6 Discharge
As mentioned within the results section, several issues arose in relation to
measuring discharge. Within supraglacial streams, depth measurements must be
calibrated due to the erosion of the streambed. At this point, no measurement method has
been found to eliminate the need for some calibration.
Fortunately, a number of velocity area method measurements were made during
this research which allowed for easy calibration, as well as a near perfect correlation
between depth and discharge measurements. The only negative is that there is a small
section of missing data on July 19 when the sensor was not in the stream.
Given that discharge was calculated using the correlation between measured and
calculated depth measurements. The measured discharge was not used in the calculation
and can be used to calculate the uncertainty. The average uncertainty between the
-0.5
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2.5
3.5
4.5
07 08 09 10 11 12 13 14 15 16 17 18 19 20
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119
measured and calculated discharge was 7x10-4 m3/s, which can be seen as error bars in
Figure 7.18.
Figure 7.18. Discharge with uncertainty shown as 7x10-4 m3/s.
7.3.7 Stream position/Step-pools/Streambed Erosion
Stream position, the step-pool sequence, and some of the streambed erosion
results were determined via the RTK surveys. During the stream survey, the average
PDOP was an ideal 0.8, the average horizontal precision was 0.8 cm, and the average
vertical precision was 1.5 cm. Therefore, the uncertainty associated with the stream
position is ±0.4 cm in the horizontal and ±0.75 cm in the vertical. It is important to note
that the uncertainty is probably slightly higher due to user error associated with holding
the rod perfectly level within the moving stream.
In terms of the uncertainty associated with the step-pools, all of the pool depths
on July 13 were greater than 1 cm, and on July 22 only 1 of the pool depths was recorded
as less than 1 cm. This leaves little uncertainty.
The streambed erosion data collected with the RTK corresponds to data collected
when the depth sensor was installed in the streambed. Using the combined dataset sets
the streambed erosion was divided and estimated per day. Again, although the exact
0
0.02
0.04
0.06
0.08
0.1
08 09 10 11 12 13 14 15 16 17 18 19 20 21
Dis
char
ge (
m3 /
s)
Day (July 2014)
120
uncertainty associated with dividing the data in this manner has not been determined, the
overall trend where the streambed erosion was lower during the first week, and greater
during the second week is certain.
7.3.8 Pulsating
The time-lapse camera was installed facing the waterfall days before the pulsating
began, and was only non-operational for part of July 20-21 due to battery failure. The
only downside to the time-lapse camera imagery was the inability to detect the magnitude
of the discharge fluctuations. What was more important on the images was the
recognition of lack of discharge, which was clear to the nearest hour.
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CHAPTER EIGHT: CONCLUSION
8.1 Summary and Implications
This thesis was designed to gain an understanding of individual meteorological
parameters, surface characteristics, and stream characteristics that influenced the
supraglacial hydrology on Fountain Glacier. The analysis was carried out to establish
which weather conditions influenced changes in the surface characteristics and stream
characteristics. The discussion examined the connections between the surface
characteristics and stream characteristics, and development of the step-pool sequence and
the pulsating phenomenon. In relation to the primary objective of determining the
factor(s) that influenced the evolution of the small supraglacial stream located on
Fountain Glacier, the rainfall, and subsequent cloudy conditions, had the greatest impact.
In reference to the secondary objectives, the cloudy conditions during the entire
second week of the study impacted the surface characteristics. The ablation decreased
and led to a significant reduction in the meltwater available to enter the supraglacial
stream. When the rainfall began, the albedo increased as the cryoconite was washed off
the surface and into the supraglacial stream. Lastly, the rain also melted away the
weathering crust causing the glacier surface to become smooth and caused more rapid
overland flow.
The stream characteristics were also influenced by the rainfall. The discharge
decreased during the second week as ablation was the primary source of water. The
stream temperature increased from the warm rain and the added suspended sediment into
the stream from the surface, this meant more energy was available to melt the channel
122
perimeter. Lastly, the combination of increased stream incision and decreased ablation
caused significant stream incision.
Unlike alluvial streams where discharge is the cause of the step-pool formation,
discharge could not have been a factor as it was significantly less during the time of
formation. Alternatively, step-pool formation is thought to occur from helical flow and
the slip of transverse fractures, both also as a result of rainfall. The pulsating
phenomenon began in parallel to the formation of the step-pools and occurred when there
was constructive interference.
In conclusion, the results of this research support the hypothesis that multiple
successive rainfall events, and subsequent cloudy conditions, impacted the glacier surface
characteristics and stream hydrological characteristics.
8.2 Suggestions for Future Work
There are a number of avenues for future research on the step-pool formation and
pulsating phenomenon within supraglacial streams. In particular, a future hydrological
study on Fountain Glacier is proposed.
In terms of the step-pool formation and pulsating phenomenon, there are still
unanswered questions. Firstly, in the event of step-pool formation, it would be ideal if a
stream profile could be collected for the entire stream. In this case, a large number of the
step-pools formed farther downstream and were not captured in the streambed survey.
Also, as the profile survey was only completed twice, and multiple days of rain occurred
in between, the evolutionary time-scale for the step-pools is not fully understood.
Conducting streambed profiles on a more frequent basis would allow for insight into the
evolution of the step-pools in connection to rainfall amounts. Furthermore, as the study
123
period ended, no information could be obtained on how the stream was affected after the
rapid formation of the step-pools. It would be interesting to observe the changes in the
stream morphology days, weeks, and months after an occurrence of this event.
Secondly, the location at which the pulsating originated within the stream was
never located. Although it is known that the pulsating began somewhere within Reach 2,
it is unknown if the first pulse began in the same location each time or if the location
shifted. Further direct studies would be required to understand this aspect of the pulsating
phenomenon.
The study of the evolution of this small supraglacial stream may have larger
implications, as larger supraglacial streams have formed two uniquely incised canyons
that also exist on the surface of Fountain Glacier. These massive canyons are the most
astonishing natural feature I have ever seen, yet the processes that have caused them to
form are not understood. This has led to my personal proposal for a PhD project.
The research objective will be to understand the factors that influence the
evolution of the small supraglacial stream and the two deeply incised canyons that exist
on the surface of Fountain Glacier. In order to achieve this goal, the objectives will
include: a) the continued study of the small supraglacial stream; b) determining the
processes that have caused such large supraglacial canyons to form on the surface of
Fountain Glacier; and c) modeling the potential growth of the small supraglacial stream
to predict if it will form an additional canyon.
Presently, there are unanswered questions, and a lack of information in the field
of glacial hydrology. This proposed PhD research will offer insight into the unique
processes that cause large canyons to form, and contribute to a greater understanding of
124
both the short and long-term evolution of supraglacial streams. This will further our
knowledge of the interactions that exist between glaciers and hydrology and provide a
clearer link regarding glacier dynamics.
125
References
Anderson, M. G. & McDonnell, J. J. 2005. Encyclopedia of Hydrological Sciences :
Volume 4. Etobicoke, Ontario: John Wiley & Sons, Ltd, 2101-2693pp.
Babar, M. 2005. Hydrogeomorphology: Fundamentals, Applications and Techniques.
New Delhi: New India Pub. Agency, 286pp.
Benn, D. I. & Evans, D. J. A. 2010. Glaciers & glaciation, Second Ed. London: Arnold,
789pp.
Benn, D. I. & Evans, D. J. A. 1998. Glaciers & glaciation. London: Arnold, 802pp.
Björnsson, H., Gjessing, Y., Hamran, S. E., Hagen, J. O., Liestøl, O., Pálsson, F. &
Erlingsson, B. 1996. The thermal regime of sub-polar glaciers mapped by multi-
frequency radio-echo sounding. J. Glaciol. 42(140), 23–32.
Blatter, H. & Hutter, K. 1991. Polythermal conditions in Arctic glaciers. J. Glaciol.
37(126), 261–269.
Brock, B. W., Willis, I. C. & Sharp, M. J. 2006. Measurement and parameterization of
aerodynamic roughness length variations at Haut Glacier d’Arolla, Switzerland. J.
Glaciol. 52(177), 1–17.
Burrough, P. A. & McDonnell, R. A. 1998. Principles of Geographical Information
Systems. Oxford: Oxford University Press, 332pp.
Camporeale, C. & Ridolfi, L. 2012. Ice ripple formation at large Reynolds numbers. J.
Fluid Mech. 694, 225–251. doi:10.1017/jfm.2011.540
Carver, S., Sear, D. & Valentine, E. 1994. An observation of roll waves in a supraglacial
meltwater channel, Harlech Gletscher, East Greenland. J. Geophys. Res. 40(134),
75–78.
Cuffey, K. M. & Paterson, W. S. B. 2010. The physics of glaciers, Fourth Ed. Burlington,
MA: Butterworth-Heinemann/Elsevier, 704pp.
Cuncio, M. 2003. Ice deformation associated with a glacier-dammed lake in Alaska and
the implications for outburst flood hydraulics. M.Sc. Thesis, Portland State
University, Portand, 122pp.
Dackombe, R. & Gardiner, V. 1983. Geomorphological Field Manual. London: Allen
and Unwin.
126
Dingman, S. L. 2009. Fluvial Hydraulics. Oxford: Oxford University Press, 576pp.
Dowdeswell, E. K., Dowdeswell, J. A. & Cawkwell, F. 2007. On the Glaciers of Bylot
Island, Nunavut, Arctic Canada. Arctic, Antarct. Alp. Res. 39(3), 402–411.
doi:10.1657/1523-0430(05-123)[DOWDESWELL]2.0.CO;2
Dozier, J. 1974. Channel Adjustments in Supraglacial Streams. Icef. Ranges Res. Proj.
Sci. Result 4, 198–205.
Edwards, A., Douglas, B., Anesio, A. M., Rassner, S. M., Irvine-Fynn, T. D. L., Sattler,
B. & Griffith, G. W. 2013. A distinctive fungal community inhabiting cryoconite
holes on glaciers in Svalbard. Fungal Ecol. 6(2), 168–176. Elsevier Ltd.
doi:10.1016/j.funeco.2012.11.001
Ferguson, R. I. 1973. Sinuosity of supraglacial streams. Bull. Geol. Soc. Am. 84, 251–
256. doi:10.1130/0016-7606(1973)84<251:SOSS>2.0.CO;2
Flowatch. 2009. Flowatch Manual. Retrieved December 8, 2013, from
http://www.ntechusa.com/products/flowatch.html
Fountain, A. G. & Walder, J. S. 1998. Water flow through temperate glaciers. Rev.
Geophys. 36(3), 299–328. doi:10.1029/97RG03579
Gabler, R. E., Sager, R. J., Wise, D. L. & Petersen, J. F. 1999. Essentials of physical
geography, Sixth Ed. Fort Worth: Saunders College Publishing, 610pp.
Gardner, A. S., Moholdt, G., Wouters, B., Wolken, G. J., Burgess, D. O., Sharp, M. J.,
Cogley, J. G. 2011. Sharply increased mass loss from glaciers and ice caps in the
Canadian Arctic Archipelago. Nature 473, 357–360. Macmillan Publishers Limited.
doi:10.1038/nature10089
Glasser, N. F. & Hubbard, B. 2005. Field Techniques in Glaciology and Glacial
Geomorphology. Chichester, West Sussex, England: John Wiley & Sons, Ltd,
412pp.
Hambrey, M. 2014. Personal communications – email to the author.
Hambrey, M. & Alean, J. 2004. Glaciers, Second Ed. Cambridge, UK: Cambridge
University Press, 368pp.
Hambrey, M. J. 1977. Supraglacial drainage and its relationship to structure, with
particular reference to Charles Rabots Bre, Okstindan, Norway. Geogr. Tidsskr. -
Nor. J. Geogr. 31(2), 69–77. doi:10.1080/00291957708545319
127
Hock, R. 2005. Glacier melt: a review of processes and their modelling. Prog. Phys.
Geogr. 29(3), 362–391. doi:10.1191/0309133305pp453ra
Hubbard, B. & Nienow, P. 1997. Alpine subglacial hydrology. Quat. Sci. Rev. 16(97),
939–955. doi:10.1016/S0277-3791(97)00031-0
Inland Waters Branch. 1969. Glacier Atlas of Canada: Bylot Island Glacier Inventory.
Surveys and Mapping Branch, Department of Energy, Mines and Resources,
Ottawa, Ontario.
Irvine-Fynn, T. D. L. 2014. Personal communications.
Irvine-Fynn, T. D. L. 2004. A non-invasive investigation of polythermal glacial
hydrology: Stagnation Glacier, Byot Island, Nunavut, Canada. M.Sc. Thesis,
University of Calgary, Calgary, Alberta, 364pp.
Irvine-Fynn, T. D. L., Bridge, J. W. & Hodson, A. J. 2010. Rapid quantification of
cryoconite: granule geometry and in situ supraglacial extents, using examples from
Svalbard and Greenland. J. Glaciol. 56(196), 297–308.
Irvine-Fynn, T. D. L., Hodson, A. J., Moorman, B. J., Vatne, G. & Hubbard, A. L. 2011.
Polythermal Glacier Hydrology: A Review. Rev. Geophys. 49(2010), 1–37.
doi:10.1029/2010RG000350
Irvine-Fynn, T. D. L., Sanz-ablanedo, E., Rutter, N., Smith, M. W. & Chandler, J. H.
2014. Instruments and Methods: Measuring glacier surface roughness using plot-
scale, close-range digital photogrammetry. J. Glaciol. 60(223), 957–969.
doi:10.3189/2014JoG14J032
Isenko, E., Naruse, R. & Mavlyudov, B. 2005. Water temperature in englacial and
supraglacial channels: Change along the flow and contribution to ice melting on the
channel wall. Cold Reg. Sci. Technol. 42, 53–62.
doi:10.1016/j.coldregions.2004.12.003
Karlstrom, L., Gajjar, P. & Manga, M. 2013. Meander formation in supraglacial streams.
J. Geophys. Res. 118, 1897–1907. doi:10.1002/jgrf.20135,2013
Klassen, R. A. 1993. Quaternary Geology and Glacial History of Bylot Island, Northwest
Territories. Geological Survey of Canada Memoir, 429, 93pp.
Knighton, A. D. 1972. Meandering Habit of Supraglacial Streams. Geol. Soc. Am. Bull.
83, 201–204.
Knighton, A. D. 1981. Channel form and flow characteristics of supraglacial streams,
Austre Okstindbreen, Norway. Arct. Alp. Res. 13(3), 295–306.
128
Knighton, A. D. 1985. Channel form Adjustment in Supraglacial Streams, Austre
Okstindbreen, Norway. Arct. Alp. Res. 17(4), 451–466.
Knighton, A. D. 1998. Fluvial Forms and Processes: A New Perspective. London:
Arnold, 383pp.
Kojima, K. 1979. Snowmelt mechanism and heat budget. Meteorol. Study Notes 146, 1–
38.
Konzelmann, T. & Braithwaite, R. J. 1995. Variations of Ablation, Albedo and Energy-
Balance at the Margin of the Greenland Ice-Sheet, Kronprins-Christian Land,
Eastern North Greenland. J. Glaciol. 41(137), 174–182.
Kostrzewski, A. & Zwoliñski, Z. 1995. Hydraulic geometry of a supraglacial stream.
Quaest. Geogr 4(special issue), 164–176.
Langford, H. J., Irvine-Fynn, T. D. L., Edwards, A., Banwart, S. A. & Hodson, A. J.
2014. A spatial investigation of the environmental controls over cryoconite
aggregation on Longyearbreen glacier, Svalbard. Biogeosciences Discuss 11, 3423–
3463. doi:10.5194/bgd-11-3423-2014
Larson, G. J. 1977. Internal drainage of stagnant ice: Burroughs Glacier, southeast
Alaska. Inst. Polar Stud. Ohio State Univ. Columbus 65, 1–33.
Lemke, P & Ren, J. 2007. Observations: Changes in Snow, Ice and Frozen Ground. In:
IPCC Fourth Assessment Report (AR4) - Climate Change 2007: The Physical
Science Basis, 338–383.
Leopold, L. B. & Maddock, T. 1953. The Hydraulic Geometrv of Stream Channels and
Some Physiographic Implications. Geol. Surv. Prof. Pap. 252, 1–57.
Marston, R. A. 1983. Supraglacial Stream Dynamics on the Juneau Icefield. Ann. Assoc.
Am. Geogr. 73(4), 597–608.
McCathy, J. J., Canzianai, O. F., Leary, N. A., Dokken, D. J. & White, K. S. 2001.
Climate Change 2001: Impacts, Adaptation, and Vulnerability. Intergov. Panel
Clim. Chang. Cambridge, UK: Cambridge University Press, 1005pp.
Moody, J. A., Meade, R. H. & Jones, D. R. 2003. Lewis and Clark’s observations and
measurements of geomorphology and hydrology, and changes with time. Reston,
VA: U.S. Dept. of the Interior, U.S. Geological Survey: Circular 1246, 110pp.
Moore, R. D. 1991. A numerical simulation of supraglacial heat advection and its
influence on ice melt. J. Glaciol. 37(126), 296–200.
129
Moorman, B. J. 2005. Glacier-permafrost hydrological interconnectivity: Stagnation
Glacier, Bylot Island, Canada. (C. Harris & J. B. Murton, Eds.) Cryospheric Syst.
Glaciers Permafr. 242, 63–74.
Moorman, B. J. & Michel, F. A. 2000a. Glacial hydrological system characterization
using ground-penetrating radar. Hydrol. Process. 14, 2645–2667.
Moorman, B. J. & Michel, F. A. 2000b. The burial of ice in the proglacial environment
on Bylot Island, arctic Canada. Permafr. Periglac. Process. 11, 161–175.
Müller, F. & Keeler, C. M. 1969. Errors in short-term ablation measurements on melting
ice surfaces. J. Glaciol. 8(52), 91–105.
Munro, S. D. 1989. Surface roughness and bulk heat transfer on a glacier: Comparison
with eddy corelation. J. Glaciol. 35(121), 343–348.
Ohno, H., Ohata, T. & Higuchi, K. 1992. The influence of humidity on the ablation of
continental-type glaciers. Ann. Glaciol. 16, 107–114.
Oostrem, G. & Brugman, M. 1991. Glacier Mass-Balance Measurements: A manual for
field and office work, Science Report No. 4. National Hydrology Research Institute,
Environnent Canada, 224pp.
Parker, G. 1975. Meandering of supraglacial melt streams. Water Resour. Res. 11(4),
551–552.
Pellicciotti, F., Brock, B., Strasser, U., Burlando, P., Funk, M. & Corripio, J. 2005. An
enhanced temperature-index glacier melt model including the shortwave radiation
balance: development and testing for Haut Glacier d’Arolla, Switzerland. J. Glaciol.
51(175), 573–587.
Rabus, B. T. & Echelmeyer, K. A. 1997. The flow of a polythermal glacier: McCall
Glacier, Alaska, USA. J. Glaciol. 43, 522–536.
Sanderson, T. J. O. 1978. Thermal stresses near the surface of a glacier. J. Glaciol.
20(83), 257–283.
Sharp, M., Richards, K. S. & Tranter, M. 1998. Introduction. In: Advances in
Hydrological Processes: Glacier Hydrology and Hydrochemistry. Chichester, UK:
John Wiley and Sons, 350pp.
Solinst Leveloggers. 2014. Retrieved February 10, 2014, from
http://www.solinst.com/products/dataloggers-and-telemetry/
130
Stenborg, T. 1969. Studies of the Internal Drainage of Glaciers. Geogr. Annaer. Ser. A,
Phys. Geogr. 51(1/2), 13–41.
Trimble. 2014. Precise Positioning Technology. Retrieved March 20, 2014, from
http://www.trimble.com/unmanned/precise_positioning_technology.aspx
van der Meer, J. J. M. 2004. Spitsbergen push moraines. Dev. Quat. Sci. Amsterdam:
Elsevier Science, 212pp.
Vatne, G. & Refsnes, I. 2003. Channel pattern and geometry of englacial conduits. 6th
Int. Symp. Glacier Caves Karst Polar Reg, Madrid. (A. Eraso & C. Dominguez,
eds.), 181–188.
Vatne, G. 2001. Geometry of englacial water conduits, Austre Brøggerbreen, Svalbard.
Nor. Geogr. Tidsskr. Nor. J. Geogr. 55(2), 85–93. doi:10.1080/713786833
Wainstein, P. A. 2011. The Development and Preservation of an Arctic Proglacial Icing.
PhD. Thesis, University of Calgary, Calgary, Alberta, 179pp.
Wainstein, P. A., Moorman, B. J. & Whitehead, K. 2008. Importance of Glacier-
Permafrost Interactions in the Preservation of a Proglacial Icing: Fountain Glacier,
Bylot Island, Canada. Ninth Int. Conf. Permafr. (D. L. Kane & K. M. Hinkel, eds.),
Institute of Northern Engineering, University of Alaska Fairbanks,1881–1886.
Wainstein, P., Moorman, B. & Whitehead, K. 2014. Glacial conditions that contribute to
the regeneration of Fountain Glacier proglacial icing, Bylot Island, Canada. Hydrol.
Process. 28, 2749–2760. doi:10.1002/hyp.9787
Waugh, D. 2000. Geography: An Integrated Approach, Third. Oxford: Nelson Thornes,
657pp.
Whitehead, K. L. 2013. An integrated approach to determining short-term and long-term
patterns of surface change and flow characteristics for a polythermal arctic glacier.
PhD. Thesis, University of Calgary, Calgary, Alberta, 313pp.
Whitehead, K., Moorman, B. J. & Hugenholtz, C. H. 2013. Brief Communication: Low-
cost, on-demand aerial photogrammetry for glaciological measurement. Cryosph. 7,
1879–1884. doi:10.5194/tc-7-1879-2013