JUNEAU ICEFIELD RESEARCH PROGRAM MASS BALANCE PROGRAM 1946-2011 The Juneau Icefield Research Program (JIRP) is the longest ongoing program of its kind in North America, facilitating arctic and alpine education and expeditionary training in the fields of climate science, glaciology and glacial geology. JIRP continues to provide valuable ground truth for climate change studies.
The Juneau Icefield straddles the Coast Range of Southeast Alaska, with an exceptionally temperate maritime climate on the west side grading into a subpolar climate to the east. This makes it a prime candidate for field investigations to assess the impact of climate change. JIRP has examined the mass balance of the Juneau Icefield since 1946 with principal efforts focused on Lemon Creek Glacier and Taku Glacier. This database initiated in 1946 by Maynard M. Miller reports the annual mass balance of Lemon Creek and Taku Glacier to the World Glacier Monitoring Service (WGMS). Crucial to the survival of a glacier is its mass balance, the difference between accumulation and ablation (melting and sublimation). Climate change may cause variations in both temperature and snowfall, causing changes in mass balance. It is the most sensitive climate indicator of a glacier. Matt Beedle in his thesis and Cristiciello et al, (2010) note that Taku and Lemon Creek Glaciers are strongly influenced by ablation season temperature and suggest that both glaciers are influenced more by climatic parameters (temperature and precipitation) at the glacier terminus than in the accumulation zone. The warmer winter temperatures observed since 1977 result in an increase in the ratio of liquid to solid precipitation.
Mass balance is just like your bank account with accumulation being the deposits and ablation being withdraws. A glacier with a sustained negative balance is out of equilibrium and will retreat. A glacier with a sustained positive balance is out of equilibrium and will advance. If a glacier still has a sustained negative balance after a period of significant retreat the glacier is likely in disequilibrium and will not survive, as it has no significant annual accumulation area. This is why we measure mass balance, below is a description of how mass balance is
measured on the Juneau Icefield and the results of this long term study. It is crucial that we continue to maintain the invaluable climate-glacier data set and provide it to the WGMS.
How does JIRP Measure Mass Balance: JIRP has relied on applying consistent mass balance methods at standard measurement sites (Pelto and Miller, 1990; Miller and Pelto, 1999; Pelto, 2011). The key annual measurements are: 1) Test pits at fixed locations on Taku Glacier and Lemon Creek Glacier ranging in elevation from 950 m to 1800 m directly measuring the snow water equivalent (SWE) through the entire snowpack profile. 2) Ablation measurements at survey stakes along survey profiles, along with repeat height measurements of the stakes. 3) Ablation adjustments to snowpack observations from observations of the TSL (transient snow line; snowline at the time of observation) and ELA( equilibrium line altitude: snowline at the end of the melt season. 4) LaChapelle (1956) first noted that the density of the snowpack is consistent and fixed 540-570 kgm-3 after early July on the Taku Glacier, this has been observed by many other detailed studies since.
Snowpits The standard element used by JIRP is the snowpit. This is dug down to the previous summer surface, identified by a dirty horizon or ice horizon and density discontinuity. This layer is well developed during each ablation season on the Lemon Creek Glacier (LaChapelle, 1954). The most important and reliable measure is the depth of the snowpack retained from the previous winter. In mass balance the key element is the snow water equivalent of the snowpack. This is the amount of water that would be yielded if the snowpack was melted. The snow water equivalent is the product of depth and snow density. The density of the snowpack is quite variable during the winter and spring, but by July the mean density is consistent from year to year and location to location on the Juneau Icefield. LaChapelle (1954) noted that one of the first characteristics apparent upon examination of the snowpit profiles is the remarkable uniformity of firn density in a vertical profile, and in distribution over the glacier, and with time during the summer.
The most important aspect of the snowpit is accurate identification of the depth. Typically our snowpits hit the previous summer firn surface, not blue ice. Particularly on Lemon Creek
Glacier and below 1100 meters on the Taku Glacier, the previous summer surface does contain sufficient dust-dirt to alter the albedo. The summer surface is also a laterally continuous surface that was refrozen several times and often has depth hoar just above it. The depth hoar is a low density layer. Hence, the laterally continuous ice layer, with depth hoar above it is another feature to be used in identifying the previous yearâ€™s firn surface. It is important in identification of this layer that the lowest section of snowpit being large enough to see if the layer is dirty or laterally continuous. Snow depth can be verified in shallower snowpacks, less than 3.5 meters using probing or crevasse stratigraphy. Measurements of retained accumulation in the snowpits are completed during late July and August and are adjusted to end of the balance year values based on the variations of the TSL, observed ablation and the measured balance gradient (Pelto and Miller, 1990; Miller and Pelto, 1999). Figure 1: Map of snowpit locations on the Juneau Icefield
For each of the 17 locations where snow pits are utilized on Taku Glacier and 4 locations on Lemon Creek Glacier, a hand held GPS is used to locate the approximate location of the study area. Once onsite the southern wall of the snow pit is marked off in order to prevent contaminating any density measurements that will be taken; the south wall of the pit is used for taking the density measurements in order to mitigate any error that may come from ablation caused by direct sunlight on the snow pit wall. The south pit wall is marked off before the pit is even started (Figure 2 and 3).
Figure 2. Chris McNeil marking off the southern wall upper Matthes snow pit. Photo: Jackson Beall
The area to the north, approximately 3 X 3 meters is dug down in a sequence of four steps like a spiral stair case. The deepest level is where the previous summerâ€™s ablation horizon will be determined.
Figure 3. Above the pit being dug notice the stairs at the bottom of the image, and the area staked off at the top, where the measurements of density are made. John Wros excavating snow pit 2010. Photo by: Jackson Beall
The Snowpit will be dug down until sufficient evidence of the previous summer’s ablation horizon is present. This is determined by any combination of the following: The presence of a dirty layer, a layer of depth hoar, density discontinuity, color change, crystal size change, and or the presence of a large continuous ice layer just about the previous year’s ablation horizon. Snow pits are always dug at least 50cm into the previous yours layer in order to ensure the continuity of the layer. Once the snow pit is dug down into the previous year’s layer the southern wall of the snow pit is shaved back to expose a flat clean face from the top of the snow pit down into the previous year’s layer, this face will be used to take density measurements of the snow pack (Figure 4 and 5).
Figure 4. Kent Redell taking snow core 2010. Photo by:Jackson Beall
Figure 5. Pascaline Bourgain taking snow core sample 2011. Photo by: Kent Walters
Using a 500 cc snow corer, samples are taken every ten centimeters down the vertical profile of the snow pit into the previous yearâ€™s layer (Figure 6). Each sample is blown into a plastic bag and then weighed on a digital scale. The digital scale is setup on a tarp on one of the steps in the pit(Figure 7). Subtracting the weight of the bag from the sample, the weight and depth of each sample is recorded. The quotient of the weight and the volume of the snow yields the density. The final step of the snow pit is recording all ice lenses present in the vertical profile of the snow pack (Figure 6). Ice lenses are horizontal layer of ice formed when water pecolates through the snow pack until it hits a denser, colder layer of snow then spreads out laterally and refreezes. The depth, thickness, and continuity of the ice lenses are recorded. Due to the small size of the ice lenses the density of the ice must be assumed at 0.9 g/cm3.
Figure 6. Chris McNeil measuring the total depth of 2011 snow accumulation, and noting ice layer position upper Llewelyn snow pit in front of C-25. Photo: Kent Walters
Figure 7. Weighing snow core samples on the Lemon Creek Glacier. 2011. Photo: Kent Walters
Calculating the Snow Water Equivalent requires subtracting the total thickness of the Ice lenses from the total depth of the snow pit (from surface to previous years ablation horizon). The product of the total snow depth (minus ice lenses) and the mean density of the snow is your SWE. Finally adding the product of the total thickness of ice and the assumed density of the ice to the SWE equals the total amount water accumulated at that location.
Probing and Crevasse Stratigraphy: The snowpits are just point measurements of snowpack amidst the vast expanse of the icefield. How representative are the snowpits? To address the error resulting from extrapolations in 1984, 1998, 2004, 2005 and 2010, JIRP measured the mass balance at an additional 100-500 points with probing transects in the accumulation area to better determine the distribution of accumulation around the test pit locations. Measurements were taken along profiles at 100- 250 m intervals. The standard deviation for measurements sites within 3 km, with less than a 100 m elevation change, was +0.09 m w.e. (water equivalent); this indicates the consistency of mass balance around the snowpit sites.
Retained accumulation thickness has been observed at up to 300 points in a single summer season on Lemon Creek Glacier (1998) and 450 measurements on Taku Glacier (1998). This number of measurement can only be completed via probing. Probing can penetrate ice layers within the past winters snowpack, because under the layer is more weak snow. The previous summer surface cannot be penetrated, because the entire layer was melted and refrozen many times, raising its density and cohesion.
Annual layers in the walls of crevasses are often quite obvious. It is similar to reading tree ring width for climate analysis. Crevasse stratigraphy provides a mean to view the two dimensional nature of the annual layer, versus a point measurement that is yielded by either probing or a snowpit. Only vertically walled crevasses can be used for these observations. The key to identification of the annual layer in crevasses is the lateral continuity of the ice layer, no other feature will be continuous. The crevasse often opens up more at this layer. Crevasse stratigraphy is not a standard method used on the Taku Glacier, but has been used since the beginning of the program for assessing snow depth in specific regions of the glacier, particularly
where snow depths are large and probing cannot be used to validate the snowpits. The high ice between Camp 8 and Camp 18 is one such region.
Probing transects for determination of ablation rate In 1984, 1998 (Mauri Pelto), 2004, 2005 (Matt Beedle) and 2010 (Chris McNeil), measured the mass balance along transects from near the TSL at 900 m to 1150 m in late July on Taku Glacier using probing at a horizontal interval of 200 meter (Figure 8).Three measurements made within 25 m were averaged to determine the snowpack depth at each probing location. SWE is then determined from the snowpack depth and mean snow density observed in the snowpits along the probing transect in three locations. This directly identifies the mass balance Figure 8. Location of Probing profile on Taku Glacier
gradient at this elevation for late July
(Figure 9). The balance gradient
In late July probing transects have been utilized to
determined from probing above the TSL
determine the balance gradient in more detail in the
ranges from 3.3-3.8 mm m-1, with a mean
vicinity of the ELA. The balance gradient is the
of 3.5 mm m-1. The balance gradient has
change in mass balance with elevation.
been consistent on Taku Glacier for each year observed regardless of the respective mass balance.
The SWE at the two test pits near this profile DGTP1 and TKGTP5 provide another direct measure of the balance gradient from the TSL to the snowpits at 1000 m in late July (Figure 10). The Taku Glacier balance gradient from the TSL to the snowpits at 1000 m from 1998 to 2009
ranged from 2.8 to 3.7 mm m-1, with a mean of 3.3 mm m-1. The gradient is slightly lower than the mean 3.5 mm m-1 observed from probing above the TSL. This is why the snowpits closest to the TSL are the most important.
Figure 9: Juneau Icefield balance gradient determined from probing in various years. Not the similar gradient in this elevation range of the Taku Glacier in July near the ELA.
Figure 10. Balance gradient on Taku Glacier between the snowpits and the transient snowline, based on snowpit data 1998-2010, except 2003 when the TSL was nearly at the snowpits.
The TSL is identified in Landsat and MODIS imagery by overlaying the images on the USGS DEM. For years with multiple images, the rate of rise of the TSL is determined. This rate of rise is only calculated for periods of longer than 15 days. For example in 2006 the TSL was
identified in five Landsat images on Taku Glacier.
The TSL in 2006 rose from 370 m on May
26, to 575 m on June 10, 730 m on July 5, 800 m on July 29, and finally 980 m on Sept. 15 (Figure 5). Ablation at the TSL is the product of the observed balance gradient and the TSL rate of rise. The TSL rate of rise was 3.6 m/day from 7/5/2006 to 9/14/2006 (Figure 11). In 2004, the TSL rose from 850 m on July 15 to 950 m on August 16 to 1030 m on Sept. 1 (Figure 12). The balance gradient from snowpit data was 3.0 mm m-1 and from probing was 3.5 mm m-1. For 2006 the mean ablation rate at the TSL from July 5 to Sept. 14 was 10.8 mm/day from the snowpit balance gradient and 12.6 mm m-1 from the probing balance gradient.
Figure 13 is a comparison of MODIS and Landsat imagery from the same date in 2009 indicating the relative accuracy of MODIS is comparable, but not quite as good. For Taku Glacier there are sixteen periods since 2004 where the TSL was observed at dates separated by more than 14 days. The TSL rise ranged from 3.1 to 4.4 m/day. Mean rise of the TSL for 16 periodâ€™s averages 3.7 m/day during the July-September period, for the elevation range between 750 - 1100 m. At shorter intervals the elevation difference is too small to avoid large errors in balance gradient assessment. As evidenced in Table 2 earlier in the summer and or at lower elevations we lack the field data to corroborate the balance gradient, and the rate of TSL rise is markedly faster. For the five periods for elevations from 450-800 m the mean rate of rise was 7.9 m/day.
In 2010, the TSL rose from 750 m on July 29, to 810 m on August 5, 875 m on August 14, 925 m on August, 30, and 975 m on September 20 . The mean observed probing balance gradient was 3.3 mm m-1 and TSL rise was 3.7 m/day, yielding an ablation rate of 12.2 mm/day Taku Glacier from July 28 to September 20, 2010.
All of these rates are in water equivalent units.
Ablation of 12.2 mm/day with a density of 0.54 is equal to 2.25 cm of snow depth loss per day.
Figure 11. TSL identification on Taku Glacier in 2006 Landsat image from 9/14/2006. A=5/26/2006, B=7/5/2006, C=7/28/2006, D=9/14/2006.
Figure 12. The TSL change during the summer of 2004.
Figure 13. Side by MODIS and Landsat 7 image from 7/29/2009 is provided with the TSL noted with red dots on each image.
Mean daily ablation at the TSL can be determined from the rate of rise of the TSL and the balance gradient. Given the mean observed balance gradient near the TSL when it is between 750 and 1000 m on Taku Glacier is 3.5 mm m-1 from probing and 3.0 mm m-1 from snowpits, and daily TSL rise is 3.7 m/day, then the computed mean daily ablation is 11-13 mm water equivalent for the July-September period in the vicinity of the ELA. This illustrates the usefulness of the shift in position of the TSL for assessing ablation.
Variations in the TSL through the course of the melt season identify ablation in the vicinity of the TSL and near the end of the ablation season in the vicinity of the ELA (Miller and Pelto, 1999). This application of the TSL to glacier mass balance depends on the calibration of the TSL to ablation for specific regions on specific glaciers. The consistency of the rate of rise of TSL on Taku Glacier from 750-1050 m over periods of several weeks or more suggests it can be reliably used in mass balance assessment. The key is to make sure to get snow depth measurements using probing closer to the snowline than the lowest snowpits on Lemon Creek and Taku Glacier. Or if time allows complete a snowpit closer to the July snowline.
Ablation has been observed during the field season is observed at survey stakes along survey lines where repeat surveys are completed and through migration of the TSL (Pelto and Miller, 1990; Pelto et al., 2008). Ablation stakes, driven into the firn in the accumulation zone record
the ablation of the remaining firnpack in the accumulation zone between the snowpit accumulation measurements in July and the end of the ablation season in early September. This provides an essential measure to adjust the July accumulation thickness snowpit measurements to the end of the ablation season. The maximum number of such ablation stakes used during a single season was 200 in 1967. During the several years where more than 30 ablation stakes were emplaced, it is apparent that ablation rates above 900 m are nearly constant on the Lemon Creek Glacier. Below 900 m ablation rates increases with decreasing surface elevation.
In the ablation zone, the mass balance curve is adjusted based on the ELA and on measurements of ablation in nine different years from 1950-1997. The resulting ablation peaks at 12 m at the terminus (Pelto and Miller, 1990). Independent examination of ablation at the terminus (Motyka and Echelmeyer, 2003), has identified ablation rates at the terminus of 12-14 m during two slightly warmer than usual ablation seasons 2003 and 2004.
How accurate is the mass balance record? Possible errors in the mass balance record include the sparse density of measurement points (1 per 37 km2), extrapolation to the end of the balance year, infrequent measurements of melting in the ablation zone, and measurements carried out by many different investigators. However, Pelto and Miller (1990), suggest that these sources of error are mitigated by: 1) measuring the same locations at the same time using the same methods each year 2) using nine years of ablation data to extrapolate mass balance in the ablation zone, 3) validation of snow depth variation using probing transects. The principal error is the lack of data from the ablation zone
The Taku Glacier mass balance record has been confirmed by independent observation of glacier surface elevation change using the ongoing laser altimetry by the University of Alaska, Fairbanks (Echelmeyer et al., 1996), indicating a Ba of -0.21 m/a for the 1993-2007 period, compared to the JIRP mean Ba of -0.16 m/a. A comparison of the surface elevation from the 2000 Shuttle Radar Topography Mission and a DEM derived from the 1948 USGS mapping indicates a mean Ba of +0.45 m/a versus the JIRP record of +0.27 m/a for the 1948-2000 period
(Larsen et al., 2007).
Ba (m we)
Juneau Icefield Glacier Annual Balance 2 1.5 1 0.5 0 -0.5 -1 -1.5 -2 -2.5
LC bn Taku bn
Year Figure 14. Annual mass balance of Taku and Lemon Creek Glacier since 1953.
The long term record is can be compared with the change in ice thickness using repeat laser altimetry data and a comparison of this data with the 1948 based USGS maps (Arendt, 2006; Arendt et al., 2002; Larsen et. al., 2007). This was accomplished from a centerline profile providing a whole glacier determination. Surface elevation change is not strictly a measure of mass balance, though it is reported as such (Arendt, 2006). The observed change in Taku Glacier surface elevation was 0.69 ma-1 from 1948-1993 and -0.28 ma-1 from 1993-1997 (Arendt, 2006). The observed mass balance for these periods is from field observations is 0.38 ma-1 for 1948-1993 and -0.60 ma-1 for 1993-1997. The surface record includes the very negative mass balance of 1997, while the laser altimetry does not include this full ablation season of 1997, -1.34 m w.e. The long term observed ice surface elevation changes validates the accuracy of the mass balance record of the Taku Glacier.
Comparison of geodetic surface maps of the glacier from 1957 and 1989 allow determination of glacier surface elevation changes. Airborne surface profiling in 1995, and comparative GPS
leveling transects in 1996-1998 further update surface elevation changes resulting from cumulative mass balance changes. Glacier mean thickness changes from 1957-1989, 1957-1995 and 1957-1998 were -13.2 m, -16.4 m, and –21.7 m respectively. It is of interest that the geodetic interpretations agree fairly well with the trend of sequential balances from ground level stratigraphic measurements. To date, however, the infrequent mapping methods in this study have yielded specific balances averaging between 5 and 11% less that those resulting from our annual on-site glacilogical modeling.. For future studies this can be an important factor. The grond data are, therefore, the ones in which have the most confidence. These show cumulative ice losses of –13.9 m (12.7 m we) from 1957-1989, of –19.0 m (-17.1 m we) from 1957-1995 and –24.4 m (–22.0 m we) from 1957-1998. The mean annual balance of the 46 record is -0.48 m/a and a loss of at least 24.7 m of ice thickness for the full 46 year period from 1953-1998
Taku Glacier Taku Glacier is a temperate, maritime valley glacier in the Coast Mountains of Alaska. With an area of 671 km2, it is the principal outlet glacier of the Juneau Icefield (Fig. 1). The mean ELA during the 1998-2010 period is at 950 m. In the region of the glacier from 750-1150 m surface slopes range from 1.1 to 2.1o. This region on the Taku Glacier encompasses 131 km2 and approximately 20% of the glacier. Taku Glacier is noteworthy for its positive mass balance from 1946-1988, which resulted from the cessation of calving around 1950 (Pelto and Miller, 1990). The positive mass balance resulting from this dynamic change gives the glacier an unusually high AAR (accumulation area ratio: percentage of glacier in accumulation zone at end of hydrologic year) for a non-calving glacier and makes the glacier relatively insensitive to climate change (Miller and Pelto, 1990; Pelto et al, 2008; Criscitiello, et al., 2010). The positive mass balance is continuing to drive its advance (Pelto and Miller, 1990; Post and Motyka, 1995; Pelto et al., 2008), while all other outlet glaciers of the Juneau Icefield are retreating, and during a period when alpine glacier mass balance globally has been dominantly negative (Zemp et al., 2009). On Taku Glacier the annual ELA has risen 60 m from the 19461985 period to the 1986-2008 period. Mass balance during the two periods were +0.40 and +0.04 m/a respectively indicative of the snowline rise and a cessation of the long term thickening of the glacier.
20 18 16 14 12 10 8 6 4 2 0 1946 1949 1952 1955 1958 1961 1964 1967 1970 1973 1976 1979 1982 1985 1988 1991 1994 1997 2000 2003 2006 2009
Cumulative mass balance (m)
Figure 15. Cumulative mass balance of Taku Glacier since 1946.
Lemon Creek Glacier Lemon Creek Glacier, Alaska was chosen as a representative glacier for the 1958 IGY global glacier network (Figure 1). This choice was based on its sub-arctic latitude location and on the ongoing mass balance program of the Juneau Icefield Research Program (JIRP), that had begun in 1948 (Miller, 1972, p. 4-6; Pelto and Miller, 1990). JIRP has continued annual balance measurements on Lemon Creek Glacier through 2010. Based on 1955-57 vertical aerial photography a 1:10,000 scale map was produced in 1958 for the IGY (Heusser and Marcus, 1964). In 1957 Lemon Creek Glacier was 6.4 km long and had an area of 12.67 km2 (Figure 2)(Heusser and Marcus, 1964). In 1998 the glacier was 5.6 km long and had an area of 11.8 km2 (Figure 3) (Marcus et al., 1995). From the head of the glacier at 1300 m to the mean ELA at 1050-1100 m the glacier flows northward, in the ablation zone the glacier turns westward terminating at 600 m.
The glacier can be divided into four sections: 1) Steep peripheral
northern and western margins draining into the main valley portion of the glacier. 2) A low slope ( 40) upper accumulation zone from 1220 m to 1050 m. 3) A steeper section (60) in the ablation zone as the glacier turns west from 1050-850 m. 4) An icefall (180) leading to the two fingered
termini at 600 m. The maximum thickness exceeding 200 m is 1 km above the icefall (Figure 4)(Miller, 1972). Mean mass balance on Lemon Creek Glacier has gone from -0.3 m/a 19531985 to -0.60 m/a 1986-2010, a 27 meter thinning over the 55 years. This is why what used to be an easy step from Camp 17 onto the glacier has become a steep significant descent.
mass balance m.w.e
Lemon Creek Glacier Cumualtive Mass Balance
-5 -10 -15 -20 -25 2009
Figure 16. Cumulative mass balance of Lemon Creek Glacier since 1953.
ELA-TSL Observations Observations of the TSL and ELA can now be reliably made each year using a combination of the less frequent Landsat Imagery, and the daily MODIS imager. The latter insures an observation within a short period of the end of the ablation season. The TSL observed at that point is approximately the ELA, and based on the observed rate of change of the TSL, the final ELA can be reliably determined. The ELA in turn is fairly good indicator of mass balance. The World Glacier Monitoring Service derives plots of ELA versus annual mass balance each year. The plots generated for the WGMS from JIRP data are below. The fit for Lemon Creek Glacier is notably better (Figure 17 and 18).
1500 ELA= -0.1607bn + 1009.2 RÂ˛ = 0.865
1300 ELA (m)
1000 900 800 700 600 -2500
Net -500 Balance(mm) 500
Figure 17. Relationship of Lemon Creek Glacier annual mass balance and the equilibrium line altitude.
ELA= -0.1668x + 976.12 RÂ˛ = 0.6546
1300 1200 1100 ELA (m)
800 700 600 500 -1500
0 500 1000 Net Balance (mm)
Figure 18. Relationship of Taku Glacier annual mass balance and the equilibrium line altitude.
The rise in the TSL observed on Taku Glacier and Brady Glacier indicates a relatively consistent rate of rise over period of more than two weeks. The product of this rate of rise and the observed
balance gradient yield a mean ablation rate. Realize that this ablation rate is biased toward the latter part of the melt season (Figure 19 and 20).
Figure 19. Ablation gradient determined from the shift in the TSL.
Figure 20: Rate of rise of the TSL on Taku and Brady Glacier.
2012 Suggested Goal A key issue that deserves greater attention is the TSL elevation on Lemon Creek Glacier. This elevation should be noted by JIRP in the field on several dates. Further as in 2011 a snowpit at 950-1000 m on Lemon Creek Glacier is critically needed. Currently there is redundancy at around 1200 meters with LGTP 1-3. A pit near the location of the 2010 LGTP 5 would be ideal.
. Tobias Kuchenmeister Prussiking into a snow pit to take density measurements 2011. Large ice lenses visible. Photo: Kent Walters.
References Arendt, A.A., Echelmeyer, K.A., Harrison, W.D., Lingle, C.S., Valentine, V.B., 2002. Rapid wastage of Alaska glaciers and their contribution to rising sea level. Science 297 (5580), 382-86. Criscitiello A., Kelly M. and Tremblay B. 2010. The Response of Taku and Lemon Creek Glaciers to Climate. .Arctic, Antarctic, and Alpine Res., 42(1), 34–44. Hock R., Koostra D. nd Reijmeer C.: Deriving glacier mass balance from accumulation area ratio on Storglaciären, Sweden. In, Glacier Mass Balance Changes and Meltwater Discharge IAHS 318, 163-170, 2007. Heusser, C.E. and Marcus, M.G., 1964: Surface movement, hydrological change and equilibrium flow on Lemon Creek Glacier,Alaska. J. Glaciol., 5(37): 61-75. LaChapelle, E.R., 1954: Snow studies on the Juneau Icefield. Walters, R.A. and Meier, M.F., 1589: Variability of glacier mass American Geographical Society. JIRP Report no. 9. Larsen, C.F., R.J. Motyka, A.A. Arendt, K.A. Echelmeyer, and P.E. Geissler 2007: Glacier changes in southeast Alaska and northwest British Columbia and contribution to sea level rise, J. Geophys. Res., 112, F01007, doi:10.1029/2006JF000586. Echelmeyer, K., M. Nolan, R. Motyka, and D. Trabant. 1995. Ice thickness Measurements of Taku Glacier, Alaska, USA, and their Relevance to it Recent Behavior. Journal of Glaciology, Vol. 41, No. 139. Cambridge, England; p. 541-552. Marcus, M.G., Chambers, F.B., Miller, M.M. and Lang, M., 1995: Recent trends in The Lemon Creek Glacier, Alaska. Phvs. Geography, 16: 150-161 Miller, M.M. 1972: A principles study of factors affecting the hydrological balance of the Lemon Creek Glacier system and adjacent sectors of the Juneau Icefield, SE Alaska, 1965-1969. Institute of Water Research. Michigan State University. and the U.S. Federal Office of Water Resources Research, Department of Interior. Publication No. 33. 1-295. Pelto, M.S. and Miller, M.M., 1990: Mass balanceMiller, M.M. and M.S. Pelto. 1999. Mass Balance measurements on the Lemon Creek Glacier, Juneau Icefield, AK 1953-1998. Geogr. Ann. 81A, 671-681. Motyka, R. J., and Echelmeyer, K. A., 2003: Taku Glacier (Alaska, U.S.A.) on the move again: active deformation of proglacial sediments. Journal of Glaciology, 49: 164. Pelto, M. and M.M. Miller. 1990. Mass Balance of the Taku Glacier, Alaska from 1946to 1986. Northwest Science. Vol. 64, No.3, p. 121-130. Pelto M.S. , Miller M.M., Adema G.W., Beedle M.J., McGee S.R., Sprenke K.F., and Lang M. 2008: The Equilibrium Flow and Mass Balance of the Taku Glacier, Alaska 1950-2006. The Cryosphere, 2, 147-157. Pelto, M.S., 2011: Utility of late summer transient snowline migration rate on Taku Glacier, Alaska. The Cryosphere, 5, 1127–1133. www.the-cryosphere.net/5/1127/2011/