Processes and rates of ice loss at the terminus of Tasman Glacier, New Zealand

Ž .Global and Planetary Change 22 1999 79–91www.elsevier.comrlocatergloplacha

Processes and rates of ice loss at the terminus of Tasman Glacier,New Zealand

Jennifer Purdie, Blair Fitzharris )

Department of Geography, UniÕersity of Otago, Dunedin, New Zealand

Received 12 August 1997; accepted 19 February 1999

Abstract

The Tasman Glacier is the largest glacier in the Southern Alps of New Zealand. Despite a century of climate warming,the glacier has until recently remained at its ‘‘Little Ice Age’’ terminus position, although there has been substantialdownwasting. The lower glacier is covered by an extensive debris layer, which has redistributed ice losses in both space andtime compared to the conventional glacier response to climate change. Ice is lost at the terminus through melt of bare iceslopes, melt under the debris layer, and calving of ice into a newly established pro-glacial lake. Over the past ten years, theglacier has evolved from an ablation regime of melt under the debris and around sinkholes, to one where calving into thelake is of greater importance. This study investigates the processes and rates of ice loss at the terminus during the summer of1995.

Melt rates of ice are greatest on bare ice slopes, averaging 96 mm dayy1 of ice depth over the summer period. Meltunder the debris averaged 7 mm dayy1, and calving accounted for a specific ice loss of 125 mm dayy1. Overall, calving isthe dominant form of ice loss at the terminus, accounting for 73% of the total. Melt under the debris layer accounted for26% of terminus ablation, and melt on bare ice slopes just 1%, due to the relatively small area covered by these. Ice loss atthe terminus is therefore largely de-coupled from climatic influences, and due mainly to the effect of the pro-glacial lake.However, when water loss from the whole ablation zone of the glacier is considered, calving accounts for just 4% of ice

Ž .loss, with the largest portion of ice loss 80% coming from melt on bare ice upstream of the debris cover.A total amount of 21 million m3 ay1 of water is estimated to be supplied to Tasman River due to ice loss from the lowest

2 3 y1 Ž .4 km of glacier. 135 million m a of water is estimated to be lost below Ball Hut approximately 10 km up-glacier , with20% of this being calved from the terminus. This proportion will increase as the glacier terminus retreats up the valley, butat present rates, it will take over a century for the glacier to retreat to Ball Hut. Water stored as ice and then released fromTasman Glacier is resulting in a mean annual flow into Lake Pukaki of 4.3 m3 sy1. The water released from all glaciers inthe region due to glacier ice loss is 6% of the annual inflow to Lake Pukaki. q 1999 Elsevier Science B.V. All rightsreserved.

Keywords: Tasman Glacier; terminus; debris layer; melt; calving

) Corresponding author. Tel.: q64-3-479-8779; fax: q64-3-479-9037.Ž .E-mail address: [email protected] B. Fitzharris

0921-8181r99r$ - see front matter q 1999 Elsevier Science B.V. All rights reserved.Ž .PII: S0921-8181 99 00027-2

( )J. Purdie, B. FitzharrisrGlobal and Planetary Change 22 1999 79–9180

1. Introduction

New Zealand has over three thousand glaciersŽ .Chinn, 1988 , which have been decreasing in massthis century. The cause of this decrease is believed tobe related to an increase in global mean temperature

Ž .over this period of about 0.58C Salinger, 1982 . Theresponse of glacier terminii to these climatic changes

Ž .are complex Meier and Post, 1987 .The aim of this study is to examine loss of ice at

the terminus of Tasman Glacier, and its effects on anew pro-glacial lake. The processes involved areidentified, and water outflow from loss of ice quanti-fied.

Tasman Glacier is the largest glacier in NewZealand and is located to the east of the Main Divideof the Southern Alps. It is 29 km long, and approxi-

Ž .mately 2 km wide at its terminus Fig. 1 . Thisglacier has been decreasing in mass for most of this

Ž .century Skinner, 1964 , although the lower half ofthe glacier is covered by an extensive debris layer,which results in a marked reduction of surface meltrates. Large temperate valley glaciers like those inNew Zealand often display a pattern of decreasingmass which begins with slow downwasting of de-bris-mantled tongues. This is followed by develop-ment of a pro-glacial lake and rapid retreat of a

Ž .calving terminus Kirkbride, 1993 . Tasman Glacierhas followed this pattern, and in the last 15 years, a

2 Ž .2.5 km lake has formed at its terminus 760 m asl ,which has instigated calving. It has been found thatwhen iceberg calving is the dominant ablation mech-anism, glacier retreat can become largely decoupledfrom climatic influences, initiating independent cy-

Ž .cles of advance and retreat Warren, 1992 . Changesin the terminal position then reflect local topographic

Ž .and bathymetric controls Warren, 1991 . Little isknown about the effect of the new Tasman Glacierterminal lake on rates of ice loss at the terminus.

Ž .Studies by Kirkbride 1995a in the mid-1980s foundunusually low surface velocities near the terminus of

Žthe glacier. More recent studies Kirkbride and War-.ren, 1997 show that these velocities have almost

doubled since the onset of calving at the terminus.Few studies of calving glacier fronts are able todirectly measure calving rates at glacier fronts due totheir inherent instability and inaccessibility. Warren

Ž .and Glasser 1992 studied the calving response ofSouth Greenland glaciers to climate change. Theyfound that, although historical information exists onthese glaciers, ice depths, calving fluxes and speeds,and the bathymetry of the fjords and lakes near theice fronts remain unknown due to the difficulty ofaccess. Tasman Glacier has a relatively accessiblecalving terminus, enabling direct measurements ofterminal retreat to be undertaken.

Ice loss on Tasman Glacier in the last 100 yearshas been mainly through downwasting, rather than

Ž .terminal retreat. Skinner 1964 estimated a verticalloss of 82 m of ice at Ball Hut for the first half of

Ž .this century. Hochstein et al. 1995 infer a down-wasting rate ranging between 0.3 and 1.2 m ay1. Atthe same time, the ice velocity of Tasman Glacier

Ž .has slowed. Brodrick 1894 reported a surface ve-y1 Ž .locity near Ball Hut of 167 m a . Anderton 1975

reported a velocity of 88 m ay1 in the same place.A field study of ice melt on Tasman Glacier by

Ž .Kirkbride 1989 shows that it varies with seasonand altitude. Ablation on bare ice reached a maxi-mum of 117 mm dayy1 in summer near Ball Hut.Loss of ice at this site was found to be 17 m yeary1,and 10 m yeary1 further up the glacier at De laBeche corner. Ice melt under the debris layer onTasman Glacier was small in the mid-1980s, at0.6–1.3 m ay1. Exposed ice wall backwasting in the

y1 Ždebris-covered area equalled 20 m a Kirkbride,.1989 . The debris mantle has a pronounced effect on

ice melt under the debris. Debris enhances melt onŽice when its thickness is less than 30 mm Mattson et

.al., 1993 . The insulation effect increases with in-creasing debris thickness up to a thickness of 2 m, at

Žwhich point ablation practically ceases Bozhinskiy.et al., 1986 . The debris mantle on Tasman Glacier

Ž .ranges from 0 to 3 m in thickness Kirkbride, 1989 ,so it has greatly reduced melting in the terminusarea. Current ice melt rates are unknown, as arerelative ice losses due to different processes of massloss.

The glacier tongue has slowly downwasted to thepoint where englacial conduits have become ex-posed. These formed sinkholes which spread intosmall lakes, and gradually coalesced into one large

Ž .lake Kirkbride, 1993 , informally named ‘‘TasmanLake’’. This covered an area of 1.95 km2 in April

( )J. Purdie, B. FitzharrisrGlobal and Planetary Change 22 1999 79–91 81

Ž .Fig. 1. Schematic map of location of Tasman Glacier and the terminus area inset from Kirkbride, 1993 .

1993, when its depth was approximately 120 mŽ .Hochstein et al., 1995 . Calving of icebergs into this

lake has greatly accelerated ablation and retreat ofthe terminus.


2. Methodology

Three main types of ice loss are examined:Ø melt of bare ice slopesØ melt under the debris layerØ calving of icebergs into the lake.

In addition to this, a sub-aqueous cliff profilesurvey was undertaken around the cliff edge to try toascertain the depth of ice that was being lost tocalving, and the resulting profile of the sub-aqueousice.

Measurements were undertaken between mid-ŽJanuary and mid-March 1995 with some being re-

.peated in February 1996 on the lowest 2.5 km ofTasman Glacier, covering an area of 4 km2.

2.1. Melt of bare ice

Movement of ice near the lake edge causes bareice slopes to be exposed within the debris-coveredlower glacier. As these bare ice slopes melt back, thedebris falls off the top and piles up at the bottom,and so eventually, the slope is buried again. Six bareice areas are examined in this study. These representa variety of slope angles and aspects. Ablation stakesmeasured ice wall retreat, which approximated melton bare ice slopes. The stakes were 1-m lengths of21 mm diameter white PVC plumbing pipe, markedwith a scale. A hand-held ice auger was used to drillholes perpendicular to the ice surface, to a depth of 1m. Ablation stakes were then inserted into the holes.The stakes were monitored every second day todetermine the rate at which the slope was retreating.

Ž .The ‘‘straight-edge method’’ Adams, 1966 wasadapted for use to measure slope recession. A PVCdisc, 30 cm in diameter, was placed around theexposed stake, and laid against the ice surface. Thesurface level was then read off the stake where thedisc met the stake. At least four stakes were beingmonitored at any one time on various slopes. A totalof 29 stakes were inserted over the study period.

2.2. Melt under the debris layer

Melt under the debris layer was estimated usingŽ .the method of McSaveney 1975 . This involves

measurement of the temperature gradient within thedebris layer. The conduction and advection of heatthrough the debris layer can then be determined, andthe melt of the ice underneath estimated.

At the site chosen, the debris was excavated to theice underneath. The debris was 1.1 m deep, and thedebris ice interface quite abrupt. Six thermocoupleswere placed at regular intervals throughout the debrislayer. These were inserted several centimetres intothe debris adjacent to the excavated hole, in anattempt to place them in ‘‘undisturbed’’ debris. Theywere then linked to a Campbell 21X micrologger,and the debris replaced. Temperatures were loggedevery 30 s, and averaged every hour, for the periodmid Jan–mid March 1995.

2.2.1. Melt caused by conductionMelting of ice under the debris mantle by conduc-

tive heat transfer has been estimated using the knownlinear relationship between thermal gradient, conduc-

Ž .tivity, and melt rate Kirkbride 1989 . Following theŽ .method used by McSaveney 1975 and subsequently

Ž .Kirkbride 1989 , the following assumptions areadopted:Ø the debris layer approximates to a thin planar

sheet,Ø heat flow is in response to a temperature gradient

normal to the plane of the sheet,Ø the debris cover may be treated as a uniform

aggregate of isotropic thermal conductivity,Ø ablation of ice is due to heat flow across the

ice–debris interface, therefore it is the thermalconductivity and gradient close to this interfacewhich are important.

Ž .McSaveney 1975 suggests the thermal conduc-tivity of a debris mantle ranges from 1.5 to 2.0 W

y2 y1 Ž .m 8C . Kirkbride 1989 calculates the thermalconductivity of the debris mantle to be 1.8 W my2

8Cy1. This is the value used in this study. Thisconductivity is the weighted mean of the thermalconductivities of the rock, air, water and occasion-ally, ice components of the debris mantle and is an

Ž .estimate only. Melting caused by conduction Q isc

calculated using the temperature gradient in the lower25% of the debris layer.

2.2.2. Melt caused by adÕectionHeat is also transferred through the debris mantle

by the flow of water and air. The advection depth ofthe debris mantle is based on the detection of air

Ž .motion within this debris layer. Kirkbride 1989 hassuggested that wind ‘‘pumping’’ would only affectair in the uppermost 10 to 30 cm of densely packed


debris with small particles in the interstices. Largeboulders in the mantle will allow deeper penetrationof advective heat. Air advection is considered to beof negligible importance at this site.

Water advection is more important, however, as ithas a high heat capacity, and is capable of significantheat transfer. An assumption made in the followingcalculations is that, where debris cover is thick, thetemperature of percolating rain water will equilibri-ate with that of the debris. Water temperature istherefore, assumed to be equal to the mean tempera-ture in the lower 25% of the debris layer. The

Ž .amount of ice melted by advective heat transfer Qa

can then be calculated as:

tc mw wQ sa Qw

Žwhere: Q samount of ice melted by advection mmay1 .day ; tswater temperature in lower 25% of de-

Ž . Žbris layer 8C ; c sspecific heat of water 4.186=w3 y1 y1.10 J kg K ; m smass of water fallen on thew

Ž . Ž y1 .glacier rainfall mm day ; Q s latent heat ofwŽ y1 .fusion of ice 0.334 MJ kg .

The rainfall is measured at the terminus of Tas-man Glacier.

Ž .The total sub-debris melt Q is the sum of meltt

caused by conductive and advective heat transfersŽ . Žthrough the debris layer Kirkbride, 1989 Q sQt c

.qQ . To calculate the annual melt, the length ofa

the ablation season is needed. This is given byŽ .Kirkbride 1989 as being 250 days for Tasman

Glacier, and supported by observations from K.Ž .Wakelin, of Glacier Explorers pers. commun., 1995 .

2.3. CalÕing into the lake

The terminus of Tasman Glacier is a steep icecliff at the lake edge. This cliff shows approximately20–25 m of bare ice above the water’s edge.

Ž .Hochstein et al. 1995 found that the total glacierthickness is approximately 160 m, including theportion under the water. The terminal ice cliff has aright angle at its upper edge where it meets the nearhorizontal plane of the glacier surface. Ice is lostthrough melting of the cliffs, both above and belowthe water, and by calving of icebergs from the face.Warm water at the lake surface cuts into the ice, sodeep horizontal fissures or ‘‘thermal notches’’ grow

back from the edge of the cliff. These undercut thecliff until large icebergs dramatically calve off. Thesnout of the glacier may also be floating, whichenables much deeper fissures to develop as thestresses of a grounded body of ice are released.

In order to quantify rates of cliff recession, sixtransects at right angles to the cliff top were estab-lished from the cliff edge back up the glacier. Theywere marked with paint at 2-m intervals. Thesetransects were monitored at two weekly intervalsthroughout the summer. The rate at which the cliffreceded due to melting and calving into the lake wasthen calculated.

2.4. Sub-aqueous cliff profile measurements

Gravity surveys undertaken by Hochstein et al.Ž .1995 show the terminal ice cliff as being vertical,along the actively calving ice cliff, to the lake floor120 m below. To determine whether any shelf of icewas left when icebergs calved at the water surface,depth measurements were undertaken around the cliffedge. Measurements were taken using a line loweredfrom a boat at several locations around the calvingice cliff face in February 1996. In December 1997, aline was lowered in several places from the cliff topto ascertain the water depth very close to the cliff.

3. Results

3.1. Melt of bare ice

Fig. 2 traces the melt on the six bare ice slopesmeasured over the ablation season. The average re-cession of ice slopes due to surface melting was 96mm dayy1 of ice depth. At this rate, over a 250 dayablation season, bare ice slopes would recede 24 m.The maximum recorded melt was 175 mm dayy1 ofice depth, and the minimum recorded melt 30 mmdayy1. Bare ice slope melt, therefore, accounts forbetween 7.5 m and 44 m in an ablation season. Themost melt for the season was measured on an eastfacing slope, closely followed by that on north facingslopes. South facing slopes had the least meltrecorded for the season.


Fig. 2. Average daily melt on bare ice slopes with different slope aspects.

This study area on the terminus of Tasman Glacierinvolves an area of debris covered ice of 4 km2.

Ž .Kirkbride 1989 estimated bare ice slopes to ac-count for 5% of the glacier surface at the terminus.Using this figure, and an ice density of 900 kg my3,it is estimated that an average of 864 m3 dayy1 of

Žwater is lost through melt on ice slopes a maximum3 y1.estimate is 1575 m day . Over one ablation sea-

son, an average of 216 000 m3 of water would meltfrom bare ice slopes at the terminus.

3.2. Melt under the debris

Large temperature fluctuations at the surface ofthe debris mantle are dampened by the debris, so

strong diurnal fluctuations do not reach the ice sur-face. Fig. 3 shows the attenuation and lag of temper-ature fluctuations through the debris layer. A strongtemperature gradient existed through the debris layerover the entire ablation season. The surface tempera-ture ranged between 3.3 and 34.18C, and the temper-ature at the debris–ice interface stayed more con-stant, within the narrow range of y0.4–1.28C.

Melting by conductive heat transfer has beenestimated using the relationship between thermal

Žgradient, conductivity, and melt rate Kirkbride,.1989 . Only one site was able to be measured on this

study. This site had a debris thickness of 1.1 m. Asthe debris ranges from 0 to 3 m in thickness over the


Fig. 3. Example of temperature through the debris layer for Julian days 39 to 45r1995.

Ž .glacier surface Kirkbride, 1989 , an assumption ismade that this site is an average debris thickness andan average ablation rate for the terminus area. An

average melt rate of 6.7 mm dayy1 of specific icedepth was calculated for the site. Over the ablationseason, this would mean a loss of 1680 mm of ice at


this site. The minimum melt rate recorded was 3.3mm dayy1, and the maximum 10.8 mm dayy1. Asdebris covers 95% of the 4 km2 area studied, it canbe calculated that an average volume of 21 443 m3

dayy1 of water is lost through melt under the debrisby conductive heat transfer.

Melting of ice under the debris mantle by advec-tive heat transfer is calculated as 0.1 mm dayy1 ofspecific ice depth. Considering that debris covers95% of the 4 km2 terminus area studied, it can becalculated that an average volume of 283 m3 dayy1

of water is lost by heat transfer of water. Melt underthe debris by advective heat transfer therefore hasthe least effect on glacier ablation. Melt under the

Ž .debris was due mostly to conduction 99% , and onlyŽ .slightly due to advection 1% .

3.3. CalÕing into the lake

3.3.1. Measured calÕingCalving occurs in two main ways. Firstly, by the

cutting action of the warmer surface water horizon-tally into the ice, and the calving of the block aboveit. Secondly, by the formation of deeper fissuresparallel to and upstream of the cliff face, which leadsto fracturing below the waterline.

Thermal notches were estimated to extend backinto the ice cliff for up to 10 m. When this occurs,the weight of the ice block above caused the ice tofracture down from above, several metres back fromthe cliff edge. A fissure appeared and graduallyexpanded over a period of weeks until block collapseoccurred. This form of block collapse is assumedonly to occur when the lake surface is ice-free, aseason of approximately 250 days.

The other form of calving that occurs at theterminus of Tasman Glacier is caused by the releaseof confining pressure within the ice as it meets thelake water, and is presumed to occur year round. Asthe ice mass expands, crevasses appear in the iceparallel to the glacier front, and columnar slab calv-ing occurs. When this is coupled with a buoyant icetongue adjacent to a terminal lake, friction from theglacier bed is lessened, and movement in the buoyanttongue further enhances expansion and calving.

The length of actively calving cliff around thelake edge is approximately 4600 m. The top of the

cliff receded between 4.5 and 7.5 m over the 2months of the field season. This equates to an aver-age daily recession of 0.08 m dayy1 of ice. How-ever, the rate of recession varied along the cliff face.

An average of 41 339 m3 dayy1 of water was lostthrough calving. The maximum recorded figure was62 260 m3 dayy1. Over the 1994–95 year, an esti-mated 14 734 086 m3 of water was lost throughcalving of ice.

3.3.2. An alternatiÕe estimate of ice calÕing ratesThe shoreline has changed since last recorded in

Ž .1993 by Hochstein et al. 1995 . By examining aerialphotographs taken in 1995, the area of ice lost at theedge of the lake has been calculated. The volume ofice lost is estimated as 22 097 850 m3 yeary1 ofwater equivalent. This equates to a cliff recession of125 mm dayy1 of ice, or 41.3 m ay1.

This change in shoreline, assessed over 2 years,indicates an ice loss that is higher than that measuredby cliff top transects during the field season. It wasobserved over the study period that the six markedtransects did not seem to record any large calvingevents, although they were observed elsewhere alongthe cliff face. It is assumed, therefore, that calvingcalculated by the change in shoreline is the morerealistic estimate for long-term ice loss from theterminal ice cliffs.

3.4. Total ice loss at the terminus

The total measured ablation for the study period isthe sum of the ice melt on bare ice slopes, the melt

Table 1Relative ice loss at the terminus of the Tasman Glacier

Form of ablation Average specific Total averageloss of ice at daily lossmeasurement of water oversite terminus area

y1 3 y1Ž . Ž .mm day m day

Melt on bare ice slopes 96 860Melt under the debris 7 21700CalvingŽ .a measured recession 83 41300Ž .b change in shoreline 125 62600

Ž .Total a 188 63900Ž .Total b 228 85200


under the debris, and the ice that calved into thelake. This equals a figure of 228 mm dayy1 of iceloss during the ablation season. The total amount ofwater lost over the study area in the ablation seasonequals 85 202 m3 dayy1, or about 1 m3 sy1 of riverflow.

Melt on bare ice slopes is an order of magnitudehigher than that under the debris. However, becausebare ice slopes only constitute about 5% of theterminus area, the total amount of water released bytheir melting is small in relation to melt under the

Ž .debris Table 1 . Calving of ice into the lake is theŽ .predominant form 73% of total ice loss occurring

at the terminus of Tasman Glacier. Melt under thedebris accounts for 26%, and melt on bare ice slopesjust 1%. These figures relate to ice loss at the

Žterminus only defined for the purposes of this study2 .as the lower 4 km of the glacier . Calving accounts

for a maximum of 4% of ice loss from the wholeablation zone, and melt of bare ice above the debrismantles zone accounts for approximately 80% of iceloss.

Fig. 4. Results of sub-aqueous cliff profile measurements at 10locations along the terminal ice face, Tasman Glacier.

3.5. Cliff profile measurements

Both surveys undertaken showed the ice cliff wasvertical in some places, and had a shelf or gentlysloping sub-aqueous ice apron in others, as can beseen in Fig. 4. It seems that the ice platform, leftwhen ice calves above the thermal notch at thewaterline, may leave an ice shelf that melts backfairly rapidly. Further study would clarify this point.

4. Discussion

4.1. Ice melt at the glacier surface

The lake exerts an influence on other ablationprocesses occurring at the terminus. It plays a signif-icant role in enhancing bare ice melt several hundredmetres back from the lake edge. The formation of thelake and the onset of calving has caused an accelera-tion of ice at the terminus, and the replacement ofcompressional flow with a tensional regime. Thismovement of ice towards the lake edge exposes freshbare ice slopes amongst the supra-glacial debris.

The bare ice slopes are covered with a fine layerof mud, which is deposited by debris sliding downthe ice surface. This layer on ‘‘bare ice’’ slopes inthe terminus area of Tasman Glacier varied consider-ably on each slope, but ranged between 0 and 20mm. Bare ice slopes with a thin layer of debris are

Ž .estimated by Oke 1987 to have an albedo of ap-proximately 0.25. This is lower than that of cleanice, therefore producing a higher melt rate. Mattson

Ž .et al. 1993 found that melt on an ice slope in-creased with increasing debris thickness, up to 10mm, mainly due to albedo effects. Melt decreasedfor any thicker cover due to insulation of the sub-strate ice. Similar effects were noted by DrewryŽ . Ž .1972 , and Clark et al 1994 .

4.2. The effect of the debris layer

On Tasman Glacier, where the debris layer rangesfrom 0–3 m in thickness, the insulating effect on theunderlying ice is demonstrated to be strong. Thedebris of the lower Tasman Glacier suppressed abla-


tion on substrate ice by 93%. This finding agreeswith other studies, such as Nakawo and YoungŽ . Ž .1981 , Bozhinskiy et al. 1986 and Mattson et al.Ž .1993 , who also found that debris suppresses abla-tion rates on underlying ice and, when several metresthick, can isolate the underlying ice from the thermalinfluence of the atmosphere almost completely. Mc-

Ž .Saveney 1975 noted that debris reduced ablation byŽ .80%, and Kirkbride 1989 by 89%. Consequently, it

is clear that the debris layer on Tasman Glacier hasbeen the major contributing factor in the lack ofretreat of the glacier terminus, despite about 100years of climate warming.

This study shows that melt under the debris ismostly due to conduction of heat through the debris

Ž .layer 99% . Melt due to advection was only 1% ofthe total. This compares well with McSaveneyŽ .1975 , who noted that conduction supplied 94% ofthe heat causing melt on the Sherman glacier inAlaska, and advection only 6% of the heat. Simi-

Ž .larly, Kirkbride 1989 found that conduction re-sulted in 95% of melt under the debris of TasmanGlacier, and advection 5%.

4.3. CalÕing

The terminal ice cliff retreated through calving, ata rate of over 30 m ay1. This is small in comparison

Žwith studies on 12 Alaskan glaciers Brown et al.,.1982 , which had calving rates between 220 and

y1 Ž3700 m a , and a Chilean glacier Warren et al.,. y11995 , which calved at a rate of 4500 m a . Retreat

of Tasman Glacier terminal position, however, hasincreased from 0 to over 30 m ay1 in the last 20years. The ice loss through calving may be enhancedif the terminus is floating. The lake level, whichfluctuates by several metres a day, is clearly ob-served to change on the rocky shoreline at theSouth-western end of the lake. The same changes inlevel, however, are not observed on the calving icecliff. The clearly defined thermal notch does notusually change its intersecting level with the watersurface. On one occasion, however, this has beenobserved to occur. The lake level dropped by severalmetres and a new thermal notch began to be pro-duced at a lower level on the ice face. The authors’viewpoint, then, is that the tip of the glacier tongue

is probably buoyant at times. More field work wouldclarify this point.

4.4. Ice loss at the terminus

Table 2 compares ice melt rates on the lowerTasman Glacier measured in this study with earlierstudies. It can be seen that calving, which was notoccurring at the times of the earlier studies, nowaccounts for the majority of ablation occurring within2 km of the terminus. Otherwise, measured meltrates are generally similar to those of Goldthwait and

Ž . Ž .McKellar 1962 and Kirkbride 1989 , and showthat melt on bare ice is an order of magnitude higherthan that under the debris.

In this study, we measured melt under a debristhickness of 1.1 m. Over the field season, meltaveraged 6.7 mm dayy1 of ice depth. KirkbrideŽ .1989 found melt under a similar debris thickness tobe 7.5 mm dayy1 of ice depth. These figures areclose, given the assumptions and natural variability.Kirkbride measured different debris thicknesses, andcalculated an increased ablation gradient up-glacierdue to decreasing debris thicknesses. He estimatedtotal ice loss from the surface of Tasman Glacierbelow Ball Hut to equal a volume of 113=106 m3

ay1 of water. This figure is the same as that derivedŽ .earlier by Goldthwait and McKellar 1962 for the

whole Tasman Glacier for the 1958r59 ablationseason. They did not measure ablation under thedebris layer, but assumed it to be ‘‘small comparedwith that on white ice’’. The current study alsomeasured ice loss through calving, which was notoccurring at the time of these earlier studies. Calvingproduces a loss of ice of 22=106 m3 ay1 of water

Table 2Comparison of ice loss data in Table 1 with previous studies ofTasman Glacier

This Kirkbride Goldthwait andŽ . Ž .study 1989 McKellar 1962

Melt on bare ice slopes 96 80 Sth facings55y1Ž . Ž .mm day Nth facings110

Melt under the debris 7 4 )

Ž .1.1 m thicknessy1Ž .mm day

y1Ž .Calving mm day 125 ) )

Total 228 84 83


on Tasman Glacier. This figure is 20% of annual iceloss below Ball Hut, and just 4% of ice loss in theablation zone. The total ice loss figure below BallHut including calving losses would therefore equal135=106 m3 ay1 of water.

The amount of ice loss from Tasman Glacier hasŽ .often been debated. Hochstein et al. 1995 consider

the downwasting rate of Tasman Glacier to haveincreased this century, from 0.3–0.5 m ay1 in theperiod 1914–64, to 1.2 m ay1 from 1972–82. Kirk-

Ž .bride 1995a found a marked decrease in ice veloc-ity near the terminus contributes to accelerated thin-

Ž .ning over time. Kirkbride 1995b estimates an in-crease in summer bare ice ablation at Ball Hut since1890 from 63 mm dayy1 to 88 mm dayy1. This is an

Ž .increase of 39% over this period. Ruddell 1995states that the runoff from Tasman Glacier is now

Ž .17% more than in the 1890s. Chinn 1995 however,observes that the mass balance of Tasman Glacierwas negative from the 1950s until 1973, but hasgenerally been positive since, apart from a stronglynegative year in 1990.

A loss of 0.14 km2 ay1 was measured at theterminus of Tasman Glacier solely due to calving.This is confirmed by aerial photos which show that,through expansion of sinkholes and calving, 0.17km2 ay1 of ice has been lost at the terminal lake andsinkholes in the terminus area between 1971 and

Ž .1995. Ruddell 1995 found that this century, thearea of the Tasman Glacier has decreased from 108.3km2 in the period 1860–90, to 99.4 km2 in the1980s. This equates to a loss of 0.09 km2 ay1. Therates found in this study are substantially higher thanthis. This is because the rate of ice loss has acceler-ated in the last 15 years due to the onset of calving.Estimates of ice loss are imprecise over the period1960s to 1993, due to the lack of information on lakedepths over this time. Whatever mass balancechanges are occurring up-glacier, this study hasshown that, at the terminus, rapid retreat is now inprogress, with consequent substantial ice losses.

4.5. Release of water stored as ice

The amount of ice being released by TasmanGlacier has been estimated at 135=106 m3 ay1

water equivalent. This equates to a mean annual river

3 y1 Ž 3 y1flow of 4.3 m s 0.7 m s of this is due to.calving . As most ice is released in the summerŽ .months Dec–Feb , the inflow to Lake Pukaki from

the Tasman Glacier ice loss in this period is 8.7 m3

sy1. This is a small contribution to the hydro lakesof the Waitaki catchment.

5. Conclusions

This study sets out to identify and quantify theprocesses and rates of ice loss at the terminus ofTasman Glacier. In the last 10 years, the terminus ofthe glacier has changed from a regime of slowdownwasting under the extensive debris-mantle, toone of calving into the new pro-glacial lake. Theformation of this lake at the terminus and the onsetof calving has heralded a period of more rapid retreatof the glacier front.

Several studies have examined the changes inŽTasman Glacier this century Goldthwait and McKel-

lar, 1962; Skinner, 1964; Kirkbride, 1989, 1993,.1995a,b; Hochstein et al., 1995 . The downwasting

of the glacier and the formation and spreading ofsinkholes has been observed. Studies on other NewZealand glaciers have observed similar patterns of

Ž .downwasting Kirkbride, 1993 . Tasman Glacier hasrecently reached an important threshold with theonset of retreat by calving, and with the relative easeof access to the terminal ice face, this study was ableto closely examine the processes operating, and therelative rates of ice loss occurring at the terminus.

The debris layer on Tasman Glacier suppressessubstrate ablation by 93%. Average melt under thedebris layer was 7 mm dayy1, compared to 96 mmdayy1 on bare ice slopes amongst the debris. Moresinkholes are being formed through englacial roofcollapse upstream of the present lake, and makingeven more rapid retreat likely as these link up to jointhe present terminal lake. The terminal lake causesincreased movement of ice around its perimeter. Thismovement of the ice and overlying debris layercauses bare ice to be exposed in places, and rapidmelting to then occur on this exposed ice.

Other studies of calving glaciers have stressed theimportance of topographic and bathymetric controlswhen calving is the dominant form of ice loss at the

Ž .terminus Warren, 1991; Warren and Glasser, 1992 .


Documented measured calving rates are rare, how-ever, due largely to the inaccessibility of calvingglacier fronts. Calving is the dominant form of iceloss at the terminus, although it still contributes onlyapproximately 4% of ice loss in the whole ablationzone. Calving occurs by two main mechanisms. Thefirst is where warm surface water ‘‘cuts’’ into theadjacent ice cliff, forming a ‘‘thermal notch’’, andeventually resulting in block collapse of the iceabove the notch. The second is deeper calving, causedby the release of confining pressure within the ice asit meets the lake water. This may be enhanced by thetip of the glacier tongue being buoyant. The resultsof this calving is that the ice cliff is retreating at over30 m yeary1 up valley. The sub-aqueous profile ofthe ice cliff resulting from this calving seems to bevertical in places, and with an ice shelf or apron inothers. It appears that any shelf left just below thewater line by calving rapidly melts back to match theabove water profile, but leaving an ice ramp, orapron, at depth.

135=106 m3 ay1 of water is being released byTasman Glacier. Despite its rapid retreat rate, theglacier terminus will still take over a century toretreat to Ball Hut, 10 km up valley from the presentterminus.

The rapid retreat of this glacier after a long periodof slow downwasting shows that, with the onset ofcalving, the response of this glacier has becomelargely decoupled from climate. The effects of thelake topography and bathymetry have become domi-nant in ablation at the terminus. Similar processesare occurring on all the other larger glaciers on theeastern side of the Southern Alps. The loss of waterfrom these glaciers results in extra inflows to riversand hydro-electricity storage lakes downstream.

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Processes and rates of ice loss at the terminus of Tasman Glacier, New Zealand

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