CHAPTER
29
New Zealand’s glaciers
Trevor J. Chinn, Jeffrey S. Kargel, Gregory J. Leonard, Umesh K. Haritashya, and
Mark Pleasants
ABSTRACT
New Zealand’s mountains support 3,153 inventoried glaciers, 99.4% of this number (99.9% by
volume) on South Island, and the remaining few on
Mt. Ruapehu, a North Island volcano. Here we (1)
provide a historical, geological, and climatic context for New Zealand’s glaciers; (2) review published knowledge of their current state and recent
dynamics; (3) present a synoptic overview from
ASTER imaging of the glaciers of Mt. Ruapehu
(North Island), including relations to volcanic
activity; use ASTER to examine changes affecting
glaciers of Mt. Aoraki (Mt. Cook, South Island)
and selected areas southward to Milford Sound;
and (4) review limnological, climatic, and debris
load controls on New Zealand’s glacier fluctuations. Half or more of New Zealand’s ice mass
has disappeared since the Little Ice Age (LIA).
New Zealand has some of the world’s highest ice
mass accumulation rates, shortest glacier response
times, and greatest concentrations of glacier debris
discharge. For the smaller glaciers on steep slopes,
especially those in high-precipitation zones and descending into warm climatic zones where ablation is
rapid and response times are short, these small
glaciers are not responding to the end of the LIA,
but rather their observed fluctuations are a response
to decadal climate oscillations and centennial-scale
trends (including atmospheric warming). Decadal-
scale climate changes driving short-term glacier
fluctuations of fast-response glaciers in New Zealand correlate, foremost, to the Antarctic Oscillation
(AAO) and the Southern Oscillation Index (SOI),
and, second, to the El Niño Southern Oscillation
(ENSO). In contrast, the largest low-sloping valley
glaciers have long response times due to their great
thicknesses and insulating debris loads, and their
lengths exhibit no discernible influence from decadal climate oscillations; consequently they are far
out of equilibrium with the long-term warming and
short-term fluctuating climate. Many glacier attributes are interrelated in a web of positive and negative dynamical feedbacks. For example, high ice
discharge (which by itself is associated with short
glacier response times) can remove surficial debris
and allow rapid ablation, thereby further shortening response times. Large glacial lakes are characterized by a separate range of dynamic behavior.
Lake formation and growth are promoted on slowresponse low-gradient glaciers with thick debris
cover, and overdeepened valleys, as well as by
climatic warming. Once the lakes enlarge, coalesce,
and expand past a critical point, rapid calving and a
host of other ablation processes accelerate, commonly beyond control by further climate change.
New Zealand’s Southern Alps climate has a strong
east–west gradient affecting all its climate parameters. However, thus far the dynamical responses of
glaciers of comparable geomorphic types are very
676
New Zealand’s glaciers
Figure 29.1. Fox Glacier, West Coast, New Zealand, descending steeply from the Mt. Cook Massif and penetrating
into temperate rainforest. The glacier has relatively little debris cover and shows the characteristic crevassing and
seracs typical of high-activity high-slope maritime valley glaciers of this part of New Zealand. Also visible is the
deeply incised glacial valley and remnants of a young terraced moraine and glaciofluvial deposits (photo by Kargel,
February 2006).
similar on the east and west sides of the Main Divide of the Southern Alps. Although we observe
substantial climatic and climate change differences
across the Alps, thus far glacier responses appear to
be uniform across the entire mountain range.
29.1
INTRODUCTION
If remoteness is measured by distance from nearest
neighbors, New Zealand’s glaciers are among the
remotest in the world; only the glaciers of East
Africa are more isolated. The nearest other glacier
ice is located 2,300 km south on Balleny Island, a
glacierized Antarctic island. The next closest
glaciers are 5,100 km northwest to New Guinea
(Papua Province, Indonesia) and 7,200 km east to
Patagonia’s glaciers (Chile). New Zealand’s
glaciers, located on a midlatitude island nation,
are also exceptional for their extreme maritime
climate and lack of continentality. Fox Glacier,
for example, is famous for extending into a temperate tree fern rainforest (Fig. 29.1). Although
both Fox Glacier and neighboring Franz Josef
Glacier (Fig. 29.2) have undergone recent years of
advance, overall they, like most other New Zealand
glaciers, have retreated (or thinned) dramatically
since the latter decades of the 19th century.
The Southern Alps glaciers are similar in latitude
and elevation to those of the Northern Patagonia
Icefield in South America (100 of longitude to the
east) and the recently disappeared summit glacier of
Marion Island in the South Indian Ocean (132 of
longitude to the west) (see Section 33.3.7). Thus,
New Zealand’s glaciers fill in a huge geographic
gap in the global distribution of glaciers and include
what are arguably the most extreme maritime glaciers in the world. However, New Zealand’s glaciers
occur across a wide range of elevations and latitudes, span multiple climate zones, (Griffiths and
McSaveney 1983), and are far from homogeneous
in their characteristics and behavior. The locations
of geographic places and glaciers discussed in this
chapter are shown in Fig. 29.3.
Introduction
677
Figure 29.2. Franz Josef Glacier (West Coast, Mt. Cook area), like neighboring Fox Glacier (Fig. 29.1), has a
heavily crevassed serac-riddled valley glacier tongue (A, B), little debris cover (C), and evidence of rapid ice
extension in recent years, rather than of the compressive flow common of many glacier termini. However, the recent
advance has ended. The longer term record is mainly of retreat, as indicated by the scoured sculpted bedrock (D).
The terminus position is close to what it was in 1967 and 1994, but far advanced relative to 1980, according to
comparisons with photos given in Hooker and Fitzharris (1999) (photos by Kargel, February 2006). Figure can also
be viewed in higher resolution as Online Supplement 29.1.
Along with glaciers in many other parts of the
world, the glaciers of New Zealand have retreated
dramatically since the end of the Little Ice Age
(LIA) (Hoelzle et al. 2007, Chinn et al. 2012).
Between about ad 1750 and 1890, persistent retreat
from LIA maxima appears to have begun asynchronously and has proceeded at different rates on different glaciers; this is expected for varied response
times1 resultant from variable glacier geomorphic
and geospatial characteristics. Recession has been
rapid at some glaciers, but others have shown very
little change in length. Dates from moraines indicate that, for many glaciers, the LIA maximum was
reached as early as 1600 (Wardle 1973, Gellatly et
al. 1988). The first widespread retreats may have
begun between 1850 and 1890, which we mark as
the end of New Zealand’s LIA. During the early
1
Response time is defined as thickness/ablation rate in the
ablation zone according to Jóhannesson et al. (1989).
20th century, retreats were widespread but still
fairly minor; this was followed by rapid wasting
of land-terminating glaciers starting in the mid20th century (Gellatly 1985). After a late 20th century period of widespread advance, retreat of some
glaciers resumed and increased in the first decade of
the 21st century. Within the general century-long
recession, some fast-response glaciers made significant resurgences late last century, while most others
have steadily diminished. Lake-terminating glaciers
are retreating faster than ground-terminating
glaciers. The appearance of large proglacial lakes
in the mid to late 20th century was a tipping point,
when the climatically driven downwasting of large
debris-covered valley glaciers lowered to the level of
their meltwater outlet streams. Ice losses increased
dramatically by the addition of iceberg calving to
continuing downwasting and has divorced the
glacier reatreat rates from further climate fluctuations. The glacier records are thus complex, as are
678
New Zealand’s glaciers
Figure 29.3. Locations of the geographic places and
glaciers discussed in this chapter.
their connections to climatic influences and other
controlling factors. Below we (1) provide the geological and climatological contexts for New Zealand’s glaciers, (2) summarize detailed historical
observations of changing glaciers in the Southern
Alps, especially for the Aoraki (Mt. Cook) area, (3)
address those changes in the context of nonclimatic
processes such as debris cover and lake growth, (4)
present remote-sensing case studies extending historical field data from a synoptic perspective, and
(5) discuss climatic controls of glacier fluctuations
in New Zealand.
29.2
REGIONAL CONTEXT
29.2.1 Geologic setting
Because of the great importance—in retarding ablation loss—of rock debris cover on many of New
Zealand’s glaciers, and the high spatial variability
of rock mechanical properties in the archipelago,
we attach greater importance to the geology and
geophysics of New Zealand than we might in some
other regions of the world. A good general geology
reference is Campbell and Hutching (2007).
Whereas a dusting of rock debris on ice can double
ablation rates, a thicker debris blanket—just a few
centimeters—can reduce ablation by an order of
magnitude or more. Thus, spatially and temporally
variable rock debris on any individual glacier can
have greater influences on glacier ablation and mass
balance than shifting climate. Differences in the
amount and lithology of supraglacial debris—
hence, local geologic history—can contribute to
differing behavior of glaciers, but connections
between these aspects have been little investigated.
The New Zealand archipelago is a product of half
a billion years of plate margin interaction, since
the Cambrian. Its ancestry includes a sliver of
Gondwana, which broke away from the megacontinent in the Cretaceous, when New Zealand
received massive deposits of sand from Australia
and Marie Bird Land (now part of Antarctica).
Many of the rocks of the Southern Alps were
formed by an intense episode of metamorphism
and anatectic melting that may have peaked in
the Late Cretaceous around 68 million years ago
(Chamberlain et al. 1995). Modern New Zealand
has developed only over the past 15 million years,
and especially the last 5 million, as the fast-moving
Indo-Australian Plate has moved northeastward,
impinging against and overriding the Pacific Plate
in North Island, and slipping past it in South
Island. Consequences have included a continuing
sequence of thrust faulting, strike-slip faulting, rock
metamorphism, rapid uplift, and—in the North
Island—volcanism. Rapid buildup of the archipelago has been opposed by rapid denudation
caused by fluvial and glacial erosion (Chamberlain
et al. 1995, Furlong 2007). Coates and Cox (2002)
summarized much of the Late Tertiary development
of the New Zealand land mass from what had
been a mainly oceanic environment, rising up along
a transpressive plate margin at the edge of Gondwanaland: ‘‘As the plates began to collide, the New
Zealand crust came under pressure and the Alpine
Fault was formed. Underlying Haast schists were
still well below the surface along the line of the
fault. Between 15 and 5 million years ago Gondwana rocks of the Australian Plate were carried
north along the Alpine Fault and brought alongside
Torlesse rocks. Chlorite-grade schist came to the
surface. As the crust thickened under pressure,
new areas of land rose above sea level and the
Southern Alps were born. From 5 million years
ago to the present day, the rate of uplift accelerated,
squeezing garnet-oligoclase schist up to the surface.’’
Some 1.5–2 million years of glacial processes on
the South Island have been partly controlled by the
Regional context 679
three main types of underlying bedrock resultant
from the tectonic and geologic history: granite
and related intrusives and granite gneiss and other
high-grade metamorphic rocks of Fiordland; the
Haast schists of Otago and Westland; and the densely jointed greywacke rock of the eastern Alps.
Each rock type has responded differently to glaciation, produced differing mountain hypsometry, and
differing feedbacks on glacial climates.
The glacierized and formerly glaciated parts of
Fiordland (including Milford Sound/Mt. Tutoko
area) are underlain by massive granite gneisses
which intrinsically have low erosion rates and produce little glacial debris cover, minor interglacial
valley collapse, and minimal valley infilling. The
topography has developed mainly by subglacial
bed erosion occurring over the entire Pleistocene,
with removal of the pre-Pleistocene landscape;
headwardly propagating bed erosion due to glacier
sliding appears to be primarily responsible (Shuster
et al. 2011). Consequently, the terrain is dominated
by a spectacular set of glaciated fiords, similar to
those of Alaska, Norway, and Greenland, where
similar rocks occur and similar glaciomarine histories have taken place.
Farther north and to the west, in the schist zones,
rock strength and resistance to erosion by glacial,
fluvial, and freeze–thaw processes is reduced compared with the more resistant rocks of Fiordland.
This lower rock strength, combined with high precipitation rates and the steepness of the rapidly
uplifting Main Divide rocks, conspires to supply
a high rate of debris onto the glaciers. This debris
has infilled fiords and lake basins and left behind
massive lateral moraine walls bounding the outwash plains of each main glacier and river. Glacial
lakes exist only in distributary embayments outside
the main river valleys.
To the drier east in Otago, lower rainfall amounts
have decreased erosion rates; here, roche moutonnée features, and large Pleistocene proglacial
lakes—not yet infilled– are common. Glaciers on
the eastern greywacke terrain north of the Otago
schist terrain, have to cope with the highest debris
supplies. The friable densely jointed greywacke
readily crumbles and avalanches onto the glaciers
(Whitehouse 1983), and they are redeposited in the
lowlands, forming the Canterbury Plains (which
are a coalescing series of Pleistocene outwash
fans), leaving a testament to the massive amount
of material removed from the Southern Alps.
Glacial–interglacial epochs resulted in a thick
sequence of proximal–distal facies transitions (from
gravel to sand, silt, and peat) of the Canterbury
Plains and areas offshore (Schaefer et al. 2006,
Almond et al. 2007). Indeed, Coates and Cox
(2002) illustrate that there has been erosion of 20
km of greywacke rock from the spine of the central
Southern Alps since uplift began. The more recent
glacial epochs have left an impressive sequence of
moraines in the Tasman Valley and elsewhere
(Almond et al. 2007).
To the south of South Island, subduction speeds
and tectonic uplift rates taper off, and none of the
Subantarctic Islands along that plate margin
reaches above the equilibrium line altitude (ELA);
thus, there are neither modern glaciers, nor any
evidence of Pleistocene ice activity on these young
islands.
On North Island, the direction of subduction is
reversed from the direction of transpressional forces
and subduction that are uplifting the Southern
Alps, and a very different tectonic environment
exists. The only two mountains reaching above
the ELA on North Island are Mt. Ruapehu, a stratovolcano in the Taupo Volcanic Zone (TVZ), and
the young dormant Mt. Taranaki to the west.
Mt. Ruapehu hosts some small glaciers, whereas
Mt. Taranaki, although reaching ELA heights,
has no glaciers, but Keys (1991) reports that Mt.
Taranaki’s summit crater sports perennial snowfields.
29.2.2 Climatic context and
glacier overview
New Zealand has no marked dry season and also
has comparatively mild seasonal temperature fluctuations. The Southern Alps of New Zealand lie
athwart the prevailing westerly weather systems
and generate a strong west–east orographic precipitation gradient with an associated steep eastward rise of glacier equilibrium line altitudes.
Extreme maritime glaciers occur west of the Main
Divide, with rock glaciers and glaciers indicative of
a comparatively dry climate lying to the east. However, New Zealand glaciers are mainly high-activity
maritime types with precipitation at or well above
3 m yr1 (Chinn 1989). The West Coast glaciers are
high-activity high-throughput (i.e., high-precipitation input and high-water/ice output), maritime
types (Meier 1961). A small percentage shift in mass
input or output can result in a large absolute shift in
mass balance. Consequently, these glaciers, when
680
New Zealand’s glaciers
considered alongside most others in the world of
similar sizes, have rapid response times to climatic
perturbations.
On geologic timescales, precipitation, erosion,
and uplift rates are coupled in New Zealand. Precipitation varies widely, from about 3 m yr1 at the
West Coast shoreline and increasing to 10–15 m
yr1 only a few kilometers west of the Main Divide.
From the Main Divide eastward, precipitation
diminishes due to the föhn effect, becoming <1 m
yr1 over the eastern foothills.
As is true of most glacierized parts of the world,
mean annual and mean summer temperature are
key variable parameters responsible for forcing
glacier responses for some New Zealand glaciers.
For example, Anderson et al. (2010) found that a
1 K change in temperature has an effect on Brewster
Glacier’s mass balance equivalent to a 50% change
in precipitation. Wind also can transport and redeposit snow, moving it from peaks and slopes onto
high-altitude accumulation fields that otherwise
might not accumulate snow. Clouds also are important in the energy balance of glaciers, especially
during dry periods, when the type and coverage
of clouds can affect ablation rates (Hock 2005,
Braithwaite 2009).
The combination of winds and positive temperatures are key in the transfer of sensible heat, and so
mean windiness is another key climate variable in
partial control of glacier melt rates, in addition to
global or regional warming. The most important
climate and weather-sensitive variable parameter
controlling melt rates is direct solar insolation at
the surface; although mean annual or summertime
radiant flux at the top of the atmosphere varies little
over century timescales, mean cloud and haze
coverage during the prime melt season may shift
rapidly, thus affecting solar insolation at the surface
and altering mass balance (Hock 2005, Braithwaite
2009, and see Chapter 2 of this book by Bishop et
al.).
In sum, climate variables other than temperature
can affect glacier mass balances as much as or even
more than the direct influence of rising global temperatures; of course, these climate variables normally are interlinked. These weather influences on
New Zealand’s glaciers are well known in some
other areas of the world and, when integrated with
longer term climate, may be important controls on
glacier mass balance. A key difference is that soot
and dust from industrial sources, wildfires, and
deserts—which can be crucially important in affecting snow and ice ablation in parts of the Himalaya
and elsewhere are not important in New Zealand.
New Zealand, like almost all the rest of the
world, has experienced significant climatic warming, as well as decadal oscillations in mean temperatures since the 19th century. Fig. 29.4 shows the
overall average temperature record for the whole
archipelago, the overall warming conditions subsequent to about 1900 following late 19th century
cooling, and decadal oscillations.
Figure 29.4. New Zealand annual temperature anomaly since the 1850s, shown as the difference in annual
average minus the average for the period 1971–2000 (source: NIWA).
New Zealand’s historical glacier dynamics 681
However, as we explore in more detail in Section
29.5.2, climate change is not homogeneous in New
Zealand, though the abundance of mass balance,
glacier length, and snowline elevation data for the
country does not yet reveal an east–west dichotomy
of responses (Clare et al. 2002). The present climate
of New Zealand largely results from orographic
influences on both prevailing and exceptional storm
systems. Prevailing baroclinic wave patterns on
South Island are mainly low-pressure systems
moving eastwardly from the Tasman Sea. Under
El Niño conditions, the low-pressure centers of
these systems tend to pass New Zealand to the
south of South Island, thus bringing in classic
‘‘norwester’’ storms. Under La Niña conditions
the lows often pass near the latitude of Otago, so
that storms arrive from different directions depending on where in the South island the lows hit. Consequently, prevailing wind directions and storm
tracks vary across New Zealand depending on both
location, and the phase of the ENSO2 cycle. Averaged over the whole ENSO cycle, the West Coast
bears the brunt of these storms; the föhn effect
causes a drier east side of the Main Divide. However, this storm pattern can reverse during some
La Niña periods. El Niño also brings in tropical
storms to North Island, and their remains can cause
exceptional melting of glaciers on both sides of the
Main Divide of South Island.
New Zealand’s multidecadal climate patterns are
teleconnected not only to ENSO, but to the whole
Pacific Basin, and particularly to the Antarctic
Oscillation (Section 29.6). However, more generally
we may surmise that any long-term oceanographic
or atmospheric climate change that affects the
Pacific Basin’s oceanic and atmospheric dynamics
(periodicity, intensity, or spatial characteristics of
their climatic effects) will also affect the prevailing
storm behavior on South Island, and consequently
its glaciers as well. Local orographic control of
climate and the kinetics of glacier responses to
climate changes add further complexity, as is widely
recognized (e.g., Jóhannesson et al. 1989, Leclerq
and Oerlemans 2011). A key concept is that of
response time, which has been defined various ways
to refer to differing dynamical responses of glaciers
to changing climate conditions. Jóhannesson et al.
(1989), for instance, defined response time as the
thickness of ice in the ablation zone divided by
the ablation rate near the terminus. It is a crude
methodology, but it works to describe roughly the
number of years of lag time between a stepwise
climate perturbation and a signal seen in terminus
position fluctuation. Data on New Zealand glaciers
indicates a wide range of response times varying
from 5–8 years to over 100 years.
Since about 1850, the nation’s glaciers have lost
nearly half of their 100 km 3 of ice, estimated to have
existed at that time (Ruddell, 1995). There have
been periods of positive mass balances recorded
during the last three decades of the 20th century
when many fast-response glaciers readvanced.
However, most of these glaciers are now undergoing renewed shrinkage.
The Mt. Cook area, which is featured in case
studies below, exemplifies an area containing relatively long response–type glaciers, which tend to be
large (many >10 km 2 ) and are heavily moraine
mantled. However, many of these glaciers have surpassed a tipping point and are now characterized by
widespread proglacial lake incursion and subsequent rapid retreat. Judged by the glacier extent
time series of large glaciers, the LIA maximum
occurred between 1850 and 1890 (not synchronously). However, these maxima reflect climatic
conditions that occurred a few decades to a century
earlier.
In addition to long-term trending and decadally
oscillating climate changes (Section 29.6), two other
major phenomena control the dwindling ice (and
occasionally some regrowth) of New Zealand’s
glaciers. They include debris cover, discussed in
Section 29.5, and the formation of large supraglacial and proglacial lakes, discussed in Section
29.4.
29.3
NEW ZEALAND’S HISTORICAL
GLACIER DYNAMICS
29.3.1 Early historical observations
2
ENSO ¼ El Niño Southern Oscillation refers to a
multiannual, quasiperiodic alternation of warm and
cool surface waters in the tropical eastern Pacific, and
associated cycles in air pressure anomalies there.
However, the phenomenon has global manifestations.
Warm phases are referred to as El Niño, and cool
phases as La Niña.
Fluctuation in the frontal positions of New Zealand
glaciers has been directly monitored since they
were first visited around 1860 in the Godley and
Havelock Valleys east of Mt. Cook (Fig. 29.3), as
recorded by Ackland (1892) and Kerr and Owens
(2008). On his visit to upper Godley River glaciers
682
New Zealand’s glaciers
in 1862, geologist Julius von Haast observed that
Godley Glacier and Classen Glacier were advancing, with the ice riding over vegetation (Haast
1879, Sealy 1892); we now know this to be a familiar
behavior from that period in many parts of the
world, broadly attributed to the effects of the Little
Ice Age. These glaciers then began an uninterrupted
retreat after Haast’s visit, which has continued to
the present day. Similarly, valley glaciers at the
head of the Havelock branch of the Rangitata River
were at their maximum extents when pioneer runholder (sheep rancher) Ackland first visited them in
1866 (Ackland 1892). Ackland made a number of
subsequent visits between 1867 and 1880 and commented on the spectacular collapse of these glaciers
in the decades following his first visit (Haast 1864,
Ackland 1892). Adjacent to the Rangitata Valley
glaciers, the large Lyell and Ramsay Glaciers of
the Rakaia River headwaters were apparently at
their maxima when first visited by Haast in 1862
(Haast 1879, Gage 1951) and have also retreated
continuously since that time.
By contrast, ample historical evidence shows that
glaciers in the Mt. Cook region, particularly the
large debris-covered glaciers, must have had longer
response times to reach equilibrium with climate
perturbations, thereby delaying their attainment
of LIA maxima until about a half century later
than the fast-response glaciers. Mueller Glacier
fluctuated slightly about its maximum position until
around 1905, whereas the Tasman, Hooker, and
Murchison
Glaciers
had
long-maintained
unchanged areas and stationary fronts, but with
progressively lowering profiles (thinning) commencing in the 1890s (Brodrick 1889, 1905,
1906a, b, Harper 1896, 1934, Gellatly 1985).
Brodrick (1905) and Skinner (1964) have provided
valuable surveys of this surface lowering or ‘‘downwasting’’; however, the dynamics have since
changed with the initiation of large supraglacial
and terminal lake development in the mid-20th
century (Gellatly 1985).
29.3.2 Franz Josef Glacier’s long
historical record
New Zealand’s longest running and most detailed
observations of glacier frontal position have been
recorded at Franz Josef Glacier (Fig. 29.5). Due to
its steep longitudinal profile and extraordinarily
high snow accumulation rates, Franz Josef Glacier
has a high mass throughput, short ice residence
time, and fast-response time, and thus is exceptionally sensitive to decadal-scale climate changes (e.g.,
Bell 1910, Speight 1914, Suggate 1952, Sara 1968).
Therefore the terminus fluctuation records of Franz
Josef Glacier are well suited to provide a decadalscale proxy summary of climate changes. Literature
on the adjacent Franz Josef and Fox Glaciers
mainly describes retreat since the 1890s, and mor-
Figure 29.5. Cumulative length fluctuations of Franz Josef Glacier.
New Zealand’s historical glacier dynamics 683
aines and tree-ring dates indicate that these glaciers
attained a LIA maximum in 1750 (McKinzey et al.
2004). Interestingly, a detailed record of terminus
fluctuations constructed from historic photographs
for nearby Stocking Glacier (Salinger et al. 1983
and more recent unpublished data evaluated by
the first author) located to the east of the Main
Divide closely matches the fluctuations that
occurred on the west-side Franz Josef Glacier and
shows that there is no difference in responses of
similar-type glaciers between the east and west sides
of the Main Divide.
Franz Josef Glacier flows northwestwardly down
the steep west side of the Main Divide. Its adjacent
twin, Fox Glacier in the next valley, has almost
identical shape and behavior, but a more limited
fluctuation record. Franz Josef Glacier’s century
of dramatic retreat was interrupted by small readvances in 1948 and 1967, before a climate shift
brought on the more substantial advance of the
1980s–1990s. The two small readvances were only
apparent at the most responsive glaciers, such as
Franz Josef, and went unnoticed as minor thickening of tributaries in the majority of other glaciers,
especially in the large slow-response glaciers. The
latest advances commenced at Franz Josef Glacier
in late 1983; similar advances have occurred at
other fast-response glaciers in New Zealand. The
most recent advance phase of Franz Josef Glacier
involved a broad increase in surface ice flow speeds
and considerable thickening in addition to terminus
advance (Herman et al. 2011).
There is strong conformity between the trends of
Fox, Franz Josef, and Stocking Glaciers. Terminus
response times for Franz Josef Glacier have been
variously estimated at 5 years (Suggate 1952); 4–8
years (Soons 1971); 5–7 years (Hessell 1983, Hooker
1995); and 4–5 years (Tyson et al. 1997). At nearby
Stocking Glacier, Salinger et al. (1983) also found a
5 to 7-year terminus response time. Fox Glacier
responds almost as rapidly, though perhaps a year
longer. Length fluctuations are in accordance with
the ELA and response time simulations of Franz
Josef Glacier by Woo and Fitzharris (1992). Hence,
year-to-year climate variations are not manifested
in Franz Josef Glacier’s responses, but decadal ones
are. Furthermore, Franz Josef Glacier’s centurylong record of retreat and episodic readvances are
not a result of the end of the LIA. Attempts to
relate the old historical and more recent detailed
Franz Josef Glacier fluctuations to local climate
records of precipitation and temperature have been
made with varying success (Suggate 1950, Hessell
1983, Salinger et al. 1983, Gellatly and Norton
1984).
The nearby Tasman Glacier has a much longer
reaction time, owing to its great length and thickness, gentler valley slopes, and slightly lower precipitation regime. However, the presence of the
proglacial Tasman Lake places this glacier in a
completely different response regime than that of
Franz Josef Glacier. A record of shifting ELAs for
Tasman Glacier is available from 1959 to the present and was reported for part of this time by Chinn
(1995). This record shows a series of low ELAs from
1974 to 1977; this is not much different from the
positive balance period estimated from the response
times found for Franz Josef Glacier. Positive
balances still dominated in 1977, the year that
snowline surveys commenced (Section 29.3.3).
The year 1976 1 has been selected as the commencement time of the trend of positive mass
balances that reversed the late 20th century general
recession.
The glaciers of Mt. Cook—mainly valley
glaciers—show a great range of dynamical behavior. The ice velocities of Franz Josef and Fox
Glaciers, measured by ASTER image analysis
during two summers (2002 and 2006), are roughly
a factor of 5 greater than those of Tasman Glacier
(Herman et al. 2011), as one would expect from
glacier mean gradients as well as mean precipitation
rates in accumulation zones.
29.3.3 Proxy mass balance from the
Snowlines Program and aerial
photography
A detailed inspection of glacier ELA responses to
climate changes has been undertaken using three
decades of photos taken on annually repeated aerial
photography flights (ELA as defined by Meier and
Post 1962). These surveys, known as the Snowlines
Program, have been completed each March over the
Southern Alps from 1977 and are ongoing (Willsman et al. 2008) to record end-of-summer snowlines
as an inexpensive surrogate for recording mass balance fluctuations (Chinn 1995). This series of
oblique aerial photos, in addition to recording
end-of-summer snowline positions, has also documented many other features like snout positions,
glacial lakes, etc. Annual photographic surveys
have continued at a select set of 50 ‘‘index’’ glaciers
(Chinn 1995) distributed along east–west transects
throughout the Southern Alps.
684
New Zealand’s glaciers
Figure 29.6. Changes in specific mass balance (volume per unit area) calculated ultimately from end-of-summer
snowlines and inferred ELA shifts observed in the Southern Alps from 1977 to 2008 for small to medium-size
glaciers (excluding those with long response times). Also shown are 95% confidence limits based on the specific
mass balance of 50 individual index glaciers. Estimates for 1989/1990 are from observations of only two index
glaciers, and for one index glacier from 1990/1991.
The general trend of glacier recession over the
last 100 years has included decadal oscillations as
observed in several decades of index glacier results
obtained since 1977. These oscillations include
some widespread excursions into positive mass
balances since 1980, and most recently and strikingly in the 1990s, when all index glaciers had positive balances during some years (Willsman et al.
2008). All index glaciers have recently reverted back
to negative balances throughout the Alps. These
mass balance oscillations were expressed in terms
of average ELA shifts and then recalculated as
specific mass balances; the average specific mass
balance over the sample of Southern Alps glaciers
is shown in Fig. 29.6, demonstrating more positive
balances than negative ones in recent decades. Estimated volume change of the entire suite of index
glaciers is shown in Fig. 29.7.
Figure 29.7. Cumulative volume change in glaciers of the Southern Alps computed from mean annual departures
from the ELAs of all measured glaciers for the entire period of these surveys (data conversions use data reported by
Willsman et al. 2008, Glacier Snowline Survey, NIWA).
New Zealand’s historical glacier dynamics 685
In addition to detailed assessments of 50 index
glaciers, a larger set of 78 glaciers was monitored
via oblique air photography to record frontal variations; these were then correlated regarding response
times to climate shifts. We assume that similar
though not identical timings in the reversals of sign
of glacier length variations (retreat shifting to
advance, and vice versa) represent delayed reactions
to the same climate shifts. Differences in response
time then can be discerned from the differing dates
of these length change reversals. Reference to Franz
Josef Glacier’s record then gives absolute response
times. Results from this study indicate that all cirque glaciers appear to have reestablished equilibrium, following the end of the LIA, by the 1990s.
Therefore their volume response times to climate
change since the end of the LIA has taken less than
100 years. New Zealand’s mountain glaciers include
steep fast-responding glaciers, which do so within 5
to 20 years after a climatic perturbation. In contrast, most valley glaciers exhibit a slow dampened
response characteristic of large (thick) low-gradient
glaciers. Hence, the response times of valley glaciers
are significantly longer than those of mountain
glaciers. Notably the subset of valley glaciers that
are both heavily debris covered and strongly
affected by lake development, which we highlight
in a case study below (Section 29.4.3).
29.3.4 Glacier responses since the end of
the LIA
LIA terminus altitudes and positions, and subsequent recession distances, have been measured
using nadir view air photos of 127 glaciers in the
Southern Alps, by comparing LIA moraine positions (assumed to date from 1850; Hoelzle et al.
2007) with the 1978 frontal positions of glaciers
extracted from the New Zealand Glacier Inventory
(Chinn 2001). LIA glacier maxima were assessed
from the positions of their remnant moraines which
can be distinguished from younger fresh moraines
which typically lack vegetation and have sharpcrested forms. The majority of main Neoglacial
advances over the past 5,000 years were similar in
extent to LIA advances, and therefore moraines
from both periods are commonly nested into single
massive structures with the innermost ridge most
likely representing a feature from the later LIA
event.
Mean length reductions measured for the 127
glaciers (Chinn 1996) vary by glacier type (Fig.
29.8) but are roughly proportional to original
glacier lengths (not shown). During the period from
1890 to 1978, New Zealand glaciers shortened by an
average of 38%, with length changes showing large
variability in individual response, ranging from 0 to
6.6 km of retreat, with a mean recession rate of 13.3
m yr1 (Chinn 1996). Glaciers with proglacial
lakes do not show greater losses than other valley
glaciers; however, they do show greater temporal
and glacier-to-glacier variability in the amount of
retreat. Mean retreat rates range from 7.8 m yr1
for cirque glaciers to 17.7 m yr1 for valley glaciers.
Cirque and mountain glaciers have typically lost
nearly half their LIA lengths, whereas valley
glaciers have typically lost only a quarter of their
former lengths.
Area change is a much more significant indicator
of climate response than length change. A limited
number of area loss measurements supplied for 25
glaciers selected by glaciologist-turned-pastor
A. Ruddell (pers. commun. 1995) over the same
period (1890 to 1978) show that nearly all glacier
types have dwindled by an average 26%, with the
greatest loss of over 30% incurred by smaller mountain glaciers. Six of the largest glaciers, including
Murchison, Tasman, Hooker, Mueller, La Perouse
and Balfour, had little area loss up to 1978, and up
until that time had responded dominantly by surface lowering. Since that time the limnological
attack cycle, documented below in some detail
(Section 29.4.3.4), has dominated the control of
area losses of some of these large glaciers. Hoelzle
et al. (2007) calculated a 61% volume loss and 51%
area loss of ice for a different sample of glaciers in
the Southern Alps from the LIA maximum (ca.
1850) to the 1978 inventory date. Associated with
this retreat, glacierized area has diminished by
about 51%, making the 1990s’ extent of glacier
ice probably less than at any time during the preceding 5,000 years.
More recently, Chinn et al. (2012) have measured
ice volume loss for the entire Southern Alps
between 1976 to 2008 using annual snowline ELA
data coupled with topographic measurements of 12
of the large slow-response glaciers. They found that
the ice volume of the Southern Alps decreased
(water equivalent) from 54.5 km 3 in 1976 to 46.1
km 3 in 2008 (Chinn et al. 2012). This equates to a
rate of 0.3 km 3 yr1 over the last three decades. Of
this loss 71% was accounted for by the 12 large
slow-response glaciers. These overall rapid losses
have prevailed despite significant positive mass
balances and advances of many glaciers during
the last part of the 20th century (Chinn 1999; Fig.
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New Zealand’s glaciers
Figure 29.8. A century of mean length changes up until 1978 of various categories of glaciers (Chinn 1996).
29.8). Retreat has continued, for the most part, this
century as well.
Chinn (1996) derived an upward post-LIA ELA
shift of 84 m for cirque glaciers up until the late
1970s. This is equivalent to a warming of 0.6 C.
In a study of past and present glaciers of the
Waimakariri Basin, Chinn (1975) identified an
ELA rise of 200 m, equivalent to 1.4 C warming
since the end of the LIA. Salinger (1979) shows that
measured temperature over the previous century
had warmed by 1.0 C, with most of the increase
occurring since the 1950s. Thus, the temperature
record and ELA shifts are in rough accord.
Glaciers with proglacial lakes have been treated
separately because of their unique behaviors. The
influences of climate change, lake/calving dynamics,
and debris emplacement are difficult to deconvolve
for this family of glaciers, though the fact that many
glaciers are behaving similarly argues for an underlying climatic trigger in initiating the lake growth/
calving regime. The significance and processes
involved in the growth of ice contact lakes has been
addressed by Kirkbride (1993, 1995a), Warren and
Kirkbride (1998), Purdie and Fitzharris (1999),
Röhl, (2006), but still there is little understanding
of what happens below the waterline.
With proglacial lakes, the terminus elevation
remains constant regardless of ice volume loss or
ELA rise until the glacier retreats and withdraws
upvalley away from the lake (or advances through
and beyond the lake). Using additional data from
1978 to 1995, which includes the main period of
lake development, the retreat rates for glaciers with
terminus lakes were investigated. The mean retreat
rate during lake expansion is 50 m per year compared with 12 m per year before lake development.
The fastest recession, over 100 m per year, occurred
on Classen Glacier, whose lake was also one of the
earliest to form in the 1950s; this glacier appears to
have steadily retreated since then.
Reports from Harper (1934) indicate that
Douglas Glacier had anomalously slow lake development until accelerated growth of the lake began
in the 1970s. Mueller Glacier and Hooker Glacier
each now have terminal lakes, younger than
Tasman Lake, that have grown linearly during
the ASTER era. On each glacier, the coalescence
of small ponds and thermokarstic sinkholes has
produced single lakes that now control glacier
retreat. The position of the ice front of Tasman
Glacier had remained constant through historical
time from the 1890s until 1974 when ‘‘thermokarst’’
ponds joined to form the first proglacial lake. This
has increased dramatically in size, accelerated by
the diversion of Murchison River into Tasman
Lake as a result of a storm in January 1994. Our
observations, reported below, indicate an accelerating retreat of Tasman Glacier and accelerating
expansion of Tasman Lake. The lake must continue
to expand.
Remote-sensing case studies
Godley Valley glaciers have shown dramatic
retreat. The three tributaries which were confluent
as the single Godley Glacier in Haast’s time (1862)
have since separated into Maud, Grey, and Godley
Glaciers. Godley Glacier separated around 1950,
and withdrew up a narrow valley between a moraine and the mountainside, collapsing rapidly prior
to the formation of Godley Lake. The maximum
retreat of 6.6 km for Godley Glacier does not
entirely reflect response to climate shifts, as this
retreat includes much of the retreat of Maud Glacier and Grey Glacier. Maud Glacier and Grey
Glacier parted in 1990 at the head of a 2 km long
lake (Kirkbride and Warren 1997). Maud carries
the debris of at least two known rock avalanches
from Mt. Fletcher, and recession has ceased at its
spectacular ice cliffs.
Wilkinson Glacier, a reconstituted glacier fed by
icefalls from Mt. Evans, had no lake until about
1980. The lake now appears to have reached maximum growth; only an ice cone remains of the
glacier that once filled the depression beneath
Bracken Snowfield.
Ivory Glacier, a cirque glacier, was a glaciological
research basin from 1968 to 1975. This glacier has
shown a consistent slowly increasing rate of retreat;
a pond formed in 1953, which grew to a small lake
by 1966 when calving later created an ice-cliffed
front. Between 1999 and 2000, Ivory Glacier withdrew from the lake and remains as a small ice apron
at the base of the steep headwall. This small cirque
glacier was retreating when many other bigger ones
in New Zealand were expanding. Once lake growth
687
was initiated, the glacier quickly succumbed to
calving and retreat.
Lyell Glacier and Ramsay Glacier, both debriscovered valley glaciers located in the upper Rakaia
River valley, each show separate periods of
accelerated retreat combined with lake expansion.
Ramsay Lake is likely to double its present size,
while Lyell Glacier has been grounded behind its
lake within which a delta is growing.
29.4
REMOTE-SENSING CASE STUDIES
29.4.1 ASTER observations of
Mt. Ruapehu, North Island
North Island’s Mt. Ruapehu (Tongariro National
Park, a UNESCO mixed cultural–natural World
Heritage site) has been among the world’s more
frequently active volcanoes throughout the Holocene, including in recent decades. The mountain’s
flanks and adjacent surroundings have been shaped
by lava and pyroclastic deposition, lahars, glacial
erosion, and deposition of large Pleistocene moraines (McArthur and Shepherd 1990). Significant
volcanic eruptions and/or lahars, several of them
documented in great detail (Manville et al. 2000,
Kilgour et al. 2010), have occurred roughly every
other year for the past century.
There are 18 inventoried ice bodies, totalling 5.06
km 2 , on Ruapehu (Fig. 29.9). Of these, 7 are named
glaciers, though we define 9 including some informal names in Fig. 29.9B; the largest glacier is 0.866
Figure 29.9. ASTER VNIR images of Mt. Ruapehu with its summit lake and glaciers, showing mid-spring snow
cover (left) and midsummer retreated snowfields (right). Figure can also be viewed in higher resolution as Online
Supplement 29.2.
688
New Zealand’s glaciers
km 2 in area. These glaciers are notable mainly
because they are geographic and climatological outliers in the global glacier record; they are outliers
even for New Zealand. The glaciers are sustained by
an intense precipitation and extreme maritime
climatic regime, with about 5 m (water equivalent)
of rain and snow per year. At ASTER image resolution, changes in the extent of Ruapehu’s glaciers
are substantially obscured by other surface changes,
such as lahar formation and shifting late-season
snow patches. However, some of these changes
are notable and interesting when viewed in ASTER
data.
During the Pleistocene glaciations, a crater ice
cap and outlet glaciers formed over Mt. Ruapehu
covering 140 km 2 (McArthur and Shepherd
1990). Today’s (2013) glaciers, covering just 0.6%
of the Pleistocene ice area, are inset into the
volcanic terrain, within summit craters, ash gullies,
lahar channels and levees, and shallow valleys
bounded by lava flows. Although glacial erosion
is prodigious, cirques are poorly developed due to
competition from active volcanism. The larger
glaciers on and around the summit are partly controlled by climate variations, but their responses are
also perturbed by geothermal heat and blanketing
by pyroclastics.
Mt. Ruapehu was the site of early systematic
glaciology studies in New Zealand, stimulated by
a disastrous lahar emanating from the crater lake
on December 24, 1953 (Odell 1955; described
further below). Krenek (1958, 1959) set up the first
New Zealand mass balance studies on Whakapapa
Glacier. Krenek stated, ‘‘The summers of 1955 and
1956 were unusually warm and there was accelerated wasting and retreat of the glaciers, especially
Whakapapa glacier.’’ From 1941 to 1954 the glacier
retreated 120 m. A program to measure the Whakapapa Glacier included recording temperatures and
glacier frontal positions from 1957 to 1961 when
annual temperatures were high; and photographic
records, established with standards set up by
glaciologist A.J. Heine, showed a steady decline
in ice area (Heine 1962). Comparison of the ASTER
images of Ruapehu with the inventory survey made
from 1985 photos made by the national park staff
(Chinn 2001) shows some wasting of ice since that
time. However, quantitative assessments are made
difficult by lingering early-season snow in the
ASTER imagery.
Volcanic ash deposited on Whakapapa Glacier in
1945 was exposed during the mid-20th century
warming interval, and spectacular ice-cored ‘‘dirt-
cones’’ developed beneath the ash (Krenek 1958).
Similar cones are common in the ablation zones of
Alaskan glaciers wherever silt, sand, or fine shaley
pebbles form a thin supraglacial debris layer and
boulders and cobbles are few. Variations of fine
sediment thickness from a millimeter to a centimeter are sufficient to drive the differential ablation
needed to form these cones.
Keys (1988) undertook a major unpublished
study on the glaciers of Ruapehu. He reported that
from 1961 to 1988, the mean terminus retreat of
nine glaciers then existing on Ruapehu was 240 m
(about 9 m yr1 on average). Most of the glaciers
thinned by anywhere from 5 to 30 m (0.2 to 1.1 m
yr1 on average) in the same period.
A small lake contact glacier (informally designated ‘‘Unnamed Glacier’’ in Keys 1988), which
is actually the northern part of Crater Basin
Glacier, actually thickened in the same period by
about 30 m (Keys 1988), thus suggesting that
unique geothermal/limnological dynamics had previously taken the glacier out of equilibrium with the
climate (thinning or retreating it far more than
climate alone would cause), and it was then returning toward equilibrium after the geothermal or
limnological interaction subsided. In fact, the crater
lake outburst flood and debris flow of 1953, which
caused the Tangiwai Railway disaster (BOI, 1954),
lowered the crater lake level by 8 m, and that might
be what reduced the geothermal interaction of the
lake with the glacier and allowed Unnamed Glacier
to thicken. The southern part of the same Crater
Basin Glacier, having less contact with the lake,
thinned by 90 m in the same period. Keys (1988)
considered the two parts of Crater Basin Glacier to
be so dynamically disconnected that the northern
part deserved its own name; hence the informal
name assigned by Keys (1988; Fig. 29.9B).
Whakapapa Glacier, over 1 km long in 1955, is
said to have disappeared entirely due to a rapid
retreat between 1955 and the 1970s (Williams
n.d.). However, our ASTER time series suggests
that the area sometimes may rebuild into short-term
perennially stable snowfields. The valley floor has
been exposed by the retreat of this ice mass or
perennial snowfield, likely from a process known
as parallel downwasting. This occurs when the
glacier surface is close to the same gradient as its
bed, and is therefore nearly uniformly thick. Downwasting first lowers the glacier surface without
much retreat, but when the ice thins toward the
bed, there is a massive retreat in a very short time
interval as the bed is exposed across a wide surface
Remote-sensing case studies
689
Figure 29.10. Young lahar deposit imaged just days after formation. Figure can also be viewed in higher resolution
as Online Supplement 29.3.
area. This type of behavior may be repeated by
some other glaciers on Ruapehu in the future. However, ice of Summit Plateau Glacier was up to 130 m
thick as of 1988, so it cannot disappear entirely
anytime soon, barring a major violent eruption that
could catastrophically remove the ice.
A lake fills the summit crater, which is partly
dammed by glacier ice, making it an unstable
feature that intermittently drains, sometimes catastrophically producing lahars (Figs. 29.9, 29.10).
The lake is geothermally heated and rarely freezes.
Trunk (2005) presented a 4-year history of ASTER
thermal observations of Ruapehu’s summit lake.
Trunk’s thesis contains 25 images indicating lake
temperatures fluctuating between about 18 and
52 C. The geothermally elevated temperatures are
seasonally modulated.
The crater lake, like quite a few others in active
stratovolcano/subduction zone settings, is acidic. A
pool of molten sulphur occurs on the lake floor, and
slicks of buoyant sulphur spherules frequently
occur on the lake surface. The second author of this
contribution (J.S.K.) has observed and analyzed
similar spherules in the hydrothermal setting of
acidic pools in the Yellowstone caldera (Kargel et
al. 1999). The water column of Ruapehu’s Crater
Lake is frequently highly turbid with both suspended sulphur-bearing lake sediment and glacial
flour. Subsequently the color of the lake has been
reported to undergo wide variations from gray to
green to turquoise. In Fig. 29.11 we document some
of the lake’s color variations in an ASTER (false-
color) image time series, including changes in the
lake’s suspended or floating material. Following the
lahars of 2007, the lake appears to have partly
cleared itself of the floating or high-level content
of coarse-grained suspended sediment, leaving just
fine suspended glacial flour, thus producing the
characteristic VNIR standard false-color composite
bright blue (real color probably is similar to the
false color).
Native sulphur and associated iron sulphide
(pyrite or more commonly marcasite) and volcanic
gases are tracers of volcanic and geothermal activity. The volatiles also drive the pyroclastic and
much of the lahar activity of the volcano (Lecointre
et al. 2004). Additionally, the volatiles at Ruapehu
and similar volcanic/geothermal features are also of
interest for the role they play in biologically
mediated redox cycling of sulphur and iron, and
for insights regarding the potential production
and existence of extraterrestrial sulphur (Kargel et
al. 1999) and, speculatively, for possible geothermally hosted extraterrestrial biological systems
on planetary bodies such as Europa (Kargel et al.
2000). Therefore, the presence of glaciers with
geothermal and geochemical systems on the mountain are of special interest.
Ruapehu is truly a land of fire, ice, and brimstone. Fig. 29.12 documents the effects of pyroclastic and lahar eruptions in March and
September 2007 and the subsequent deposition of
yellowish and greenish material on the rocks and
snow/ice around the crater. The September 2007
690
New Zealand’s glaciers
Figure 29.11. ASTER time series of Mt. Ruapehu’s Crater Lake. (A) Time series prior to color normalization (dates
shown in panel C). (B) ASTER images after color normalization. (C) Crater Lake, where color has been normalized
(discussed in text), thereby allowing direct visual color comparisons. Note the growth of the lake and changes in
turbidity or of floating material, probably suspended and floating sulphur. Figure can also be viewed in higher
resolution as Online Supplement 29.4.
lahar emitted water, sediment, and sulphur. The
yellow and green materials in Fig. 29.12 are
probably sulphur and related volcanic/geothermal
sublimate deposits. This is a supposition, but makes
sense, because volatiles, such as native sulphur and
gypsum, have been observed to form, partly by
exudation of a molten phase in the case of sulphur,
on lahar sediments on Ruapehu following emplacement (Graettinger 2008, Kilgour et al. 2010). The
geothermal sublimate deposits appear to be
unstable, disappearing after several years; this is
consistent with what we see in our Fig. 29.12 time
series. A recent study (Casey 2012) compared the
trace element composition of impurities, due to dust
and aerosol deposition, in snow and ice from
glaciers in Svalbard, southern Norway, Nepal,
and Ruapehu (New Zealand) and not surprisingly
found a volcanic influence in the Ruapehu case; this
study was a landmark in high-sensitivity studies of
trace elements in glaciers and snow, and the
approach promises important advances in the study
of airborne contaminants in the cryosphere.
One lahar from Ruapehu was especially significant. In 1945, a volcanic eruption of Ruapehu
began a deadly chain of events (BOI, 2001 online
report on the Tangiwai Railway disaster). On
Christmas Eve, 1953, during an official visit of the
Queen, a massive lahar, emanating from those 1945
volcanic deposits wiped out a railway bridge and a
passenger train, killing 151 people. The glacial/
volcanic lake played a role in four ways to cause
the disaster: (1) lava, ice, and water interacted and
helped produce the fragmental volcanic debris that
later collapsed to cause the lahar; (2) ice and fragmental debris blocked the drainage and dammed
the lake; (3) ice and water added mass to the lahar;
Remote-sensing case studies
691
Figure 29.12. Summit of Mt. Ruapehu is shown here with color saturation and contrast enhancements of ASTER
VNIR images. Colors were processed band by band using dark-object subtraction and then controlling the brightpixel saturation point such that nonglacier areas distant from the summit have the same tone and color without
saturating much snow. Color was then uniformly saturated and contrast was uniformly but nonlinearly stretched in
the upper row; the bottom row is enlarged to highlight the fresh pyroclastic and lahar deposits (January 9, 2008
image, dark blue) and yellowish and greenish deposits in the February 12, 2009 scene. The pyroclastic/lahar and
yellowish/green deposits occur on both the ice-free and glacier and snow surfaces; they also appear to have formed,
then disappeared in some patches. A cloud (bright red) obscures the lower right of the January 9, 2008 scene, top
row. Figure can also be viewed in higher resolution as Online Supplement 29.5.
and (4) ice and debris yielded suddenly, thus causing a catastrophic glacier lake outburst flood and
lahar. The volcano has continued with small pyroclastic or lahar eruptions, at a rate of roughly one
per year. On March 18, 2007, a week before ASTER
acquired an excellent image, a lahar of 1.3 million
m 3 (four times the size of the 1953 lahar) was
unleashed (Fig. 29.10), but no deaths and relatively
little damage was incurred because of improved
infrastructure and a warning system that functioned
as designed.
29.4.2 ASTER observations of small
glaciers in the Southern Alps
29.4.2.1
Glaciers in the Mt. Tutoko area
(Fiordland’s Darran Range)
The cumulative imprint of Pleistocene glaciation
(including the effects of a large ice cap), the end
of the Pleistocene Ice Age, and the effects of Holo-
cene warming are readily evident in the glaciated
landscape of the Fiordland Region. The spectacular
glacial geomorphology presents its finest examples
in the Milford Sound and Mt. Tutoko areas (Fig.
29.13), where deep U-shaped valleys, fiord valleys,
hanging valleys, tarn lake basins, and now-empty
cirques record the former more extensive distribution of thick ice, and vertical valley walls leading
directly into the ocean. An ice cap free of central
nunataks likely covered the ranges south of the
Darran Range which have accordant peak heights
and are moderately ‘‘smoothed’’ without any
prominent horn peaks. In this region, only a few
peaks are high enough to penetrate the present
ELA, and there exist only scattered glacierets which
dot the peaks to the south end of Fiordland. By
contrast the higher jagged peaks of the northern
Darran Range suggest that these peaks rose above
any regional ice cap. The Darrans rise well above
the current ELA and host numerous steepmountain glaciers. Unlike the Pleistocene glaciers
692
New Zealand’s glaciers
Figure 29.13. ASTER images of the Mt. Tutoko/Milford Sound area (A) and changes in Donne Glacier observed
over a 3-year period. Note the small glacial lakes, including recent expansion of the main lake. Figure can also be
viewed in higher resolution as Online Supplement 29.6.
and the modern glaciers of the Mt. Cook area,
which together constructed the great outwash
gravel plains of Canterbury and Otago, all glacially
transported debris loads in Fiordland have been
carried out to sea by the immense outlet glaciers.
In Fig. 29.13 we highlight Donne Glacier in the
Darran Mountains, to the east of Mt. Tutoko
(Milford Sound’s highest peak). Topographic maps
of the area, including a 1:50,000-scale map based on
1986 photos, show only small supraglacial ponds
near the terminus of the glacier. A 33-year sequence
of snowline photos (from the Snowline Program)
record the following terminus activity: a ‘‘dry’’ terminus in 1977 and 1981; the first small pond
appears in 1987; larger ponds and increased debris
cover in 1995, and a similar state in 1998; some
shrinkage of the terminus lake by 2000 and then
regrowth in 2003. In 2009 the photos show total
collapse of the lower trunk and terminus, and in
2010 many icebergs are present.
ASTER imagery from 2002 shows a large rock
basin lake. Comparison of 2005 and 2002 images
shows coalescence of the ponds and retreat of the
terminus, such that Donne Glacier had established
a lake-calving type of terminus. The 2005 image
shows a shortened glacier and enlarged lake, and
evidence of a major calving event as indicated by
the icebergs in the lake. Many other small glaciers in
the area, however, show little sign of any major
changes in the same ASTER image pair. The
growth of supraglacial ponds into large proglacial
lakes, such as the lake development at Donne
Glacier, is a typical response for a low-gradient
glacier tongue that rises sharply up a steep névé
to the snowline. Although this is a spectacular process and is largely responsible for rapid loss of ice
volume in New Zealand (because of its rapid effects
on very large glaciers), the percentage of glaciers
affected by this process is small.
29.4.2.2
Brewster Glacier
This is a benchmark glacier; it is 2.5 km 2 , nearly
debris-free, and represented by a long history of
detailed monitoring and a wide variety of scholarly
research (e.g., Zemp et al. 2011). Brewster Glacier is
worth using as a case study because of the detailed
field information available and because its recent
behavior stands in contrast to other regional
glaciers, such as Donne and Tasman, which have
exhibited dramatic retreat dynamics. The glacier
appears to be far out of equilibrium with its climate.
In recent years Brewster Glacier has shown wide
fluctuations in ELA, in some years extending nearly
to the top of the glacier (Anderson et al. 2010) and
in 2007/2008 occurring above the top (Zemp et al.
2011). The two most recently reported mass balance
years showed large negative balances, 1.65 and
0.83 m yr1 for, respectively, 2007/2008 and
2008/2009 (Zemp et al. 2011). Notwithstanding this
disequilibrium, it has not significantly retreated
since Chinn’s first aerial photos were taken in
1967. Hence, it shows very subdued response with
very little change in ASTER images except for interannual shifting of the snowline.
ASTER image change assessment was carried out
by this chapter’s authors using a simple example,
with the purpose of outlining the basic technique
involved (Fig. 29.14). The method involves selecting
an image pair acquired on near-anniversary dates,
subtraction of one image from the other, and rescaling the resultant DN values to positive integers (see
Remote-sensing case studies
693
Figure 29.14. A pair of ASTER 321 RGB false-color composite images (shown at different scales—top and
bottom rows), spanning some three years of changes at Brewster Glacier, and their respective image differencing
results (right panels). The methodology used for mage differencing is described briefly in the main text, and in detail
in Chapter 4. Late-season snowlines and interannual change in snowlines are clearly evident. Figure can also be
viewed in higher resolution as Online Supplement 29.7.
Section 4.7.2 of this book by Kääb et al. on the
ICESMAP algorithm). The difference image shows
any unchanging portions of the image pair as
neutral gray tones; otherwise, areas that have
become black or dark toned may be a result of
glacier retreat or snow melt exposing bare rock or
soil, or emplacement of a rock landslide; areas that
have become brighter may represent emplacement
of a snow avalanche deposit onto a rock surface, or
glacier advance; areas that are bluer may represent
growth of a lake or a decrease of vegetation pigment
intensity (leaf area index or pigment abundance); or
redder zones may represent drainage of a pond or
growth of vegetation. In this example, changing
snow cover conditions are readily evident, which
is more likely attributed to interannual variability
than to any sort of multiyear trend.
29.4.3 ASTER observations of Mt. Cook
glaciers
29.4.3.1
Overview
Climatic controls on New Zealand’s glaciers are
evident from the state and dynamics of modern
glaciers and from assessment of moraine and other
evidence regarding ancient glaciers. More than a
century of general glacier retreat across New
Zealand is a good basis to support general climatic
warming as the foremost process at work (Denton
and Hendy 1994, Anderson and Mackintosh 2006).
However, nonclimatic controls have also affected
New Zealand’s glacial history and ongoing glacier
dynamics. These processes include (1) landslide
supply of locally abundant supraglacial debris,
and their contributions to glacier insulation and
tendency to promote glacier advances (or to slow
down or stabilize the retreat of glaciers) and moraine formation (Santamaria Tovar et al. 2008), and
(2) runaway lake growth dynamics as a cause of
major retreats. Such processes challenge any interpretations of glacier fluctuation in New Zealand
that are based solely on climate change. Mt. Cook’s
glaciers show these complexities well.
Many of New Zealand’s largest glaciers are
found in the Mt. Cook Massif. Due to the mountain’s recent geologic history of rapid uplift and its
abundant friable rock types, in addition to the high
mass turnover of glaciers in this high-precipitation
regime, Mt. Cook is eroding rapidly almost entirely
694
New Zealand’s glaciers
by rockfall as well as rock avalanche and subsurface
glacial processes. Hence, many of Mt. Cook’s
glaciers are heavily debris covered (see further discussion in Section 29.5). The larger valley glaciers
on the mountain may be generally grouped into
three dynamical types: (1) thin, steep-sloping,
fast-flowing, clean ice to lightly debris-covered
glaciers, such as Fox and Franz Josef—those having short glacier response times—are the best indicators of decadal oscillations of climate; (2) slowflowing, low-gradient, thick, heavily debris-covered
glaciers lacking large glacier lakes—these have very
long response times; (3) slow-flowing glaciers similar to (2) above but having large terminal lakes—
these types, like Tasman Glacier, respond rapidly to
their lake-calving environment, which is controlled
by the presence of the lake more than by shifting
climate (though climatic warming probably helped
to catalyze lake development).
29.4.3.2
ASTER time series of Mt. Cook
glacier changes
ASTER has produced a good time series of highquality images of the Mt. Cook Massif and its
glaciers; two of the images, acquired on nearanniversary dates 7 years apart (2002 and 2009),
are shown in Fig. 29.15. A multispectral difference
image representing changes in that same image pair,
with pixel DN values rescaled to positive values, is
shown in Fig. 29.16A (see Online Supplement 29.9
for a high-resolution version). Fig. 29.16B shows a
highly saturated nonlinear contrast-stretched version of the same data to aid visual discernment of
changes. Because the differencing involves two
images acquired on near-anniversary dates, the
solar illumination geometry is almost unchanged,
and so differential shadowing and photometric
effects due to illumination differences are almost
absent. Changes visible in the difference scene
include: shifts in the transient midsummer snowline;
vegetation growth on lateral moraines; a large rockfall and snow or ice avalanches; glacier flow; and
changes in glacier terminus position and debris
cover. The difference scene also highlights glaciers
and parts of glaciers that were flowing significantly
versus ice that was apparently stagnant over the
7-year period.
About two thirds of the largest Mt. Cook glaciers
have had almost stable termini and margins over
the 7 years represented by the ASTER image pair
(to within uncertainties conservatively estimated as
30 m). Nine other large glaciers in the scene either
advanced or retreated. Land-terminating glaciers
advanced or retreated by less than 100 m; four
out of five of them changed by between 2 and 3
ASTER pixels (30–45 m). Four lake-terminating
glaciers retreated by hundreds of meters each,
except for a nearly detached part of Tasman Glacier
as of 2002 which had retreated more than 2 km by
2009. These retreats may be compared with those
measured from early ASTER era imagery and prior
aerial photos (Kääb 2002); in general, the later
ASTER era imagery analyzed here shows a continuation of retreat as assessed by Kääb (2002).
29.4.3.3
Flow vector mapping of Tasman Glacier
Kääb (2002) and Kääb et al. (2003) first used
ASTER images to map flow velocity vectors on
Tasman Glacier and some of its tributaries; they
found that, as expected, flow speeds diminish
sharply toward the terminus, especially in thermokarstic areas of the tongue. They also found that a
major tributary, Grand Plateau, is the major contributor of ice to the tongue. In fact this ice now
appears to pinch off all ice from the upper glacier
and is probably preventing even more rapid retreat
of the calving front of Tasman Glacier. Herman et
al. (2011) have recently assessed glacier flow speeds
of the entire Mt. Cook Massif using ASTER 16-day
repeat images acquired in the midsummer seasons
of 2002 and 2006 (see Table 29.1 summary). Both
seasons’ data showed, as expected from relative
precipitation on the two orographic sides of the
mountain, that south and east-flowing glaciers are
moving 2 to 8 more slowly than glaciers flowing
west and north from Mt. Cook. Sixteen-day repeat
imagery is ideal for assessing flow speeds in the
faster range of values but cannot be used to discern
stagnant from sluggish ice. These ASTER era
results are comparable with slightly earlier glacier
flow velocity mapping results by Kirkbride (1995b).
The ice velocities of Tasman Glacier and Hooker
Glacier in the New Zealand Southern Alps were
measured between February 17, 2009 and April 4,
2010 using an ASTER image pair. Ice velocities
were computed and displayed by means of precise
orthorectification, co-registration, and subpixel
correlation of raw ASTER L1A data using the
COSI-Corr software module within ENVI, the
remote-sensing imaging software application. The
process, described in detail by Leprince et al.
(2007a, b), Scherler et al. (2008), Herman et al.
(2011), and in Chapter 4 of this book by Kääb et
al., began with manual selection of tie points
Remote-sensing case studies
695
Figure 29.15. Pair of ASTER false-color VNIR 321 images of the Mt. Cook area obtained on near-anniversary
dates 7 years apart. Details shown at right. Figure can also be viewed in higher resolution as Online Supplement
29.8.
between a raw ASTER image and an already
orthorectified image. From these tie points, ground
control points (GCPs) were automatically calculated within COSI-Corr from subpixel correlation
between the master and raw image. In this study, a
shaded version of a digital elevation model (DEM)
covering the study area was used as the first orthorectified master image. The first orthoimage created
was then used as the master image for all other slave
images to be orthorectified. Upon successful gen-
eration of a set of GCPs, mapping matrices that
assign ground coordinates to raw pixel data in
ASTER images were defined from the GCP file
and ancillary data of the image were orthorectified
and resampled. Average misregistration has a standard deviation of ¼ 1.9 m, about one eighth of a
pixel. Mapping matrices were used in the resampling process to define the grid with which
ground coordinates were assigned to the raw image.
Displacement maps were produced from subpixel
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New Zealand’s glaciers
Figure 29.16. Image difference of the near-anniversary pair shown in Fig. 29.15. (A) Difference scene rescaled to
positive pixel values: unchanging pixels are gray; A ¼ advancing; S ¼ stable; R ¼ retreating; R,c ¼ retreating and lake
calving. (B) Saturation and contrast enhancements applied. (C)–(G) Enlarged areas of interest. Figure also available
as Online Supplement 29.9.
correlation of two orthorectified images. In all
cases, ASTER image bands 3N with 15 m resolution were used and correlation was performed to
yield displacement data every 60 m. Nadir-looking
ASTER bands allowed us to avoid potential error
due to change in ice thickness. Some stripes present
as a result of satellite attitude artifacts in the east/
west and north/south bands of displacement maps
were averaged and removed using postprocessing
tools.
As the range covered by measured glacier flow
speeds varies from 20 to 225 m yr1 (Fig. 29.17),
corresponding to displacements of about 22–252 m
(409-day interval), signal to noise is primarily in the
range of 4 to 40, if noise is conservatively taken as
3. Our results are thus highly robust. Surface
velocity vector-mapping results are much as anticipated, very similar to those found by Kääb et al.
(2003), who analyzed an April 2000/April 2001
ASTER image pair.
Remote-sensing case studies
697
Table 29.1. Approximate peak and mean (or typical) flow speeds of Mt.
Cook’s glaciers. a
Glacier name
Approximate peak surface
flow speed
(m day1 )
Approximate mean speed
1.0 b
0.5 (both 2002 and 2006 data)
Murchison
0.5 (2002), 0.6 (2006)
<0.4 c (2002 and 2006)
Hooker
1.5 (2002), 1.5 (2006)
0.5 (2002), 0.7 (2006)
Fox
4.2 (2002), 4.6 (2006)
2.0 (2002), 2.4 (2006)
Franz Josef
2.6 (2002), 3.7 (2006)
1.5 (2002), 1.8 (2006)
Tasman
(m day1 )
a
Our assessments of published results of Herman et al. (2011), who analyzed 16-day ASTER
repeat imagery in midsummer 2002 and 2006.
b
Flow speeds are up to around 2 m day1 in a tributary.
c
The measurement limit appears to be around 0.4 m day1 .
As expected there is a tendency toward higher
flow speeds in more steeply sloping parts of the
glaciers, such as the icefalls of Hochstetter Glacier;
the dependence of flow speed on surface gradient is
slightly steeper than linear, but less steep than the
square of the surface gradient. Tasman Glacier also
follows the usual glacier pattern in which surface
flow speeds tend to slow dramatically towards the
terminus. The same also applies where maximum
flow speeds occur roughly along the glacier centerline, indicating flow dominated by internal ice
deformation; dominance by basal sliding would
generate plug flow, which is not seen. As Kääb et
al. (2003) also found, discharge of upper Tasman
Glacier ice has fallen to the extent that within the
last few decades the entire upper Tasman ice flow
has been pinched off by the Hochstetter ice stream
such that it now contributes very little to the lower
glacier. Formation of a glacial lake in the depression between Hochstetter Glacier and upper
Tasman Glacier ice is predicted.
We measured flow speeds near the terminus of
approximately 30–40 m yr1 . Calving of exposed ice
faces is the main process involved in the massive
increase in ice loss when a proglacial lake forms,
and Tasman Glacier is no exception. Calving losses
along the ice front must average more than 30 m
yr1 in order to cause net retreat of the calving
front. High flow speeds on the lower glacier (as
obtained by Kääb et al. 2003 and confirmed here)
are significantly greater than those measured in the
1990s, when flow speeds as low as 1.3–13 m yr1
were measured (Kirkbride and Warren 1999). This
finding supports the suggestion (above) that the
lake has expanded far enough into the glacier trunk
to significantly increase surface gradients and
trigger a drawdown increase in ice discharge. This
process appears to be catalyzing further rapid
glacier retreat.
The velocities presented here may be compared
with manual displacement determinations from
available data. A large thermokarstic part of
Tasman Glacier’s tongue, which shattered in
2006, did not flow measurably (certainly less than
30 m) during the preceding 6 years (flow <6 m yr1
or 1.3 cm day1 ). Another thermokarstic section of
the tongue, just upglacier from 2010’s calving front,
flowed measurably but only about 200 m in the
10 preceding years (mean rate of about 20 m yr1
or 5 cm day1 ). Furthermore, debris deposited by
the December 1991 rock avalanche from Mt.
Cook traveled down the Hochstetter icefall and
completely crossed the trunk of Tasman Glacier
(McSaveney et al. 1992, McSaveney 2002, Almond
et al. 2007). In the 22 years since that event, the
differential ablation ridges that developed within
the debris deposit have deformed into spectacular
parabolic arcs (visible in satellite imagery available
in Google Earth). Displacement by 2,300 m over 19
years subsequent to the avalanche along Tasman
Glacier’s centerline gives a mean flow speed during
that period of about 120 m yr1 ; our results in the
same area, but for the 2009/2010 period, indicate
flow speeds diminishing from about 130 to 90 m
yr1 (averaging around 100 m yr1 ), thus indicating
similar (possibly slightly slower) flow speeds in this
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New Zealand’s glaciers
Figure 29.17. Surface flow of Tasman Glacier and Hooker Glacier assessed from a pair of ASTER images acquired
on February 17, 2009 and April 10, 2010. (A) ASTER VNIR 321 RGB of Tasman Glacier (February 17, 2009 image)
for reference. Yellow box shows approximate outline of area shown in (B). (B) Surface flow vector field of Tasman
Glacier. (C) Map of surface flow displacements over the 2009/2010 measurement period; indicated values of
surface flow speeds are meters of displacement divided by 1.12 years. The area shown is slightly larger than that of
(A). Apparent flow speeds in Tasman Lake are meaningless. Figure can also be viewed in higher resolution as Online
Supplement 29.10.
Remote-sensing case studies
sector to those indicated by our flow vector mapping. Our new flow vector mapping results are presented with higher directional and speed resolution
than those of Kääb et al. (2003), but the flow patterns and range of speeds found are also similar.
Our new results, which have a high signal-tonoise ratio, complement those of Kääb et al.
(2003). Despite the continued spectacular growth
of Tasman Lake and retreat of Tasman Glacier’s
terminus, no other major change was detected in
surface velocity behavior in the decade between the
2000/2001 period assessed by Kääb et al. (2003),
and the 2009/2010 period assessed here. The
699
ongoing retreat of Tasman Glacier does not imply
stagnation at the tongue but rather accelerated
calving interactions. The only major sector of
Tasman Glacier where stagnation seems to be the
case is in the mid-section just above the confluence
with Hochstetter Glacier.
29.4.3.4
Measured lake growth and future lake
growth scenarios
Tasman Lake (Fig. 29.18) is the archetypal Mt.
Cook area lake. Four major Mt. Cook area proglacial lakes are undergoing similar growth (Figs.
Figure 29.18. Tasman Glacier and Tasman Lake viewed from the air and the lake surface. The highly degraded
tongue is evident. Though the contiguous mass of the tongue is still active with glacier ice still flowing, the detached
ice masses in panels (C)–(G) are inactive rapidly degrading debris-covered thermokarstic blocks of ice. These
masses shattered into many icebergs and bergy bits within months of these photos being taken (photos by Kargel,
early February 2006). Figure can also be viewed in higher resolution as Online Supplement 29.11.
700
New Zealand’s glaciers
Figure 29.19. ASTER 10-year time series of Tasman Lake, Hooker Lake, and Mueller Lake—water
index ¼ (AST1 AST3)/(AST1 þ AST3). See Online Supplement 29.12 for a high-resolution time series.
29.16, 29.18). We will later take a brief look at
Hooker Lake and Mueller Lake, but first we examine in detail the growth of Tasman Lake. Kirkbride
and Warren (1999) documented the early growth of
lakes on the terminus of Tasman Glacier and very
effectively predicted the further runaway growth
that subsequently occurred there (Hochstein et al.
1995, Dykes et al. 2010).
A time series of ASTER images of Tasman,
Mueller, and Hooker Lakes and their glaciers is
shown in Fig. 29.19; their area growth curves are
graphed in Fig. 29.20A. Tasman Lake’s growth
curve shows an abrupt lake area increase in 2006.
The second author (J.S.K.) visited the lake during
this event, noted the glacier’s complex perimeter,
and observed the presence of numerous stranded
fractured blocks of ice (Fig. 29.18); this suggested
at the time that the glacier tongue would likely
undergo a period of further rapid breakup and
rapid growth of its lake. Indeed, shortly after his
visit, the tongue shattered, as shown in the ASTER
image time series (Fig. 29.19; see Online Supplement 29.12 for a higher resolution image). However, the lake has subsequently returned to the
Remote-sensing case studies
701
Figure 29.20. Growth histories of Mt. Cook’s glacier lakes and their possible futures. (A) Tasman Lake growth
history from measurements by G. Leonard (this chapter) and Dykes et al. (2010). (B) Tasman Lake area and
perimeter (measurements from ASTER imagery by G. Leonard). (C) Second and third-order polynomial projections
and an exponential growth curve, doubling every 9.5 years, fitted to Tasman Lake area data (a composite of both
Dykes et al. 2010 and Leonard’s work). (D) Growth records for three glacial lakes. Figure can also be viewed in
higher resolution as Online Supplement 29.13.
growth curve it displayed before disintegration of
its tongue (Fig. 29.20B). It is evident, however, that
Tasman Lake’s growth rate is highly nonlinear over
longer time intervals.
Second and third-order polynomials fit the
Tasman Lake area time series very well, especially
since 1980. For purposes of future predictions,
polynomials are far from ideal, not least because
they lack any obvious physical basis. Nevertheless,
Figure 29.20C shows two polynomial projections.
Alternatively, the data can be fit with an exponential growth curve based on lake area doubling every
9.5 years (7.2% annual expansion). Though the
exponential 7.2% annual growth curve fits the time
series less well than the polynomials, it does have
some physical basis. One proposed idea is that the
lake absorbs solar radiation and transmits it to the
ice; thus, the larger the lake area, the more energy is
absorbed and transmitted, and more ice is melted.
Kääb and Haeberli (2001) produced an analytical
model that describes lake growth as a quadratic
function at an earlier stage of growth.
Once Tasman Lake has grown beyond 6,500 m in
length, its calving rate may increase dramatically as
the heat-gathering capacity of the lake increases in
proportion to its area and the calving front retreats
into the deepest part of the lake basin (Fig. 29.21).
As the lake depth near the calving front increases
and the glacier decreases in thickness, the glacier
will likely undergo another rapid breakup in a process that research geologist Bruce Molnia has
termed ‘‘disarticulation’’, in which glaciers undergo
flotation and rapid disintegration along crevasses
and other weaknesses (Molnia 2007). This process
appears to be happening at Tasman Glacier and
was probably the cause of the summer 2006
breakup (Fig. 20.19, 2006 to 2007 panels, best
shown in Online Supplement 29.12).
Lake growth and calving retreat of Tasman
Glacier will likely continue until the length of the
lake approaches 16 km and the glacier calving margin approaches the base of the Hochstetter icefall
(the dominant ice supply) disconnecting the glacier
terminus from the lake and bringing an end to
702
New Zealand’s glaciers
Figure 29.21. Schematic of longitudinal glacier surface and bed profiles illustrating how Tasman Lake is apt to
continue expanding until its length increases by about 150% from its present size (2013) and the glacier terminus
finally retreats from the lake. This expansion of the lake and stabilization of Tasman Glacier may occur sometime
between 2025 and 2034, according to Fig. 29.20C, at which point the glacier will be about half as long as it is now.
However, the retreat trend is far from simple as a result of undulations in bed topography, the likely future
detachment of Hochstetter Glacier (the most important tributary, not shown, feeding Tasman Glacier) and other
tributaries such as Ball Glacier (shown), and the possible formation of a lake when the tributary detaches.
calving. According to Fig. 29.20C, this could occur
as early as 2023 (exponential growth projection) or
toward the mid-2020s (third-order polynomial
projection) or mid-2030s (second-order polynomial
projection). However, disarticulation can proceed
very rapidly, and it would not be surprising if
most of the projected retreat occurs in the next
few years as the calving front reaches the deepest
part of the basin. By using approximate present and
future stable accumulation area ratio (AAR)
values, the retreat is predicted to continue until
around 2045.
We calculated lake growth and glacier shortening
from a different perspective. Based on Tasman
Glacier’s AAR value estimates (Dyurgerov et al.
2009) Tasman Lake could more than double its
present lake area before the glacier regains near
equilibrium with current climate. Considering the
insulating properties of heavy debris cover, lake
growth has to be seen as the greatly delayed
response of Tasman Glacier to climate change that
commenced over a century ago. With the removal
of ice support from the valley sides extensive slumping of lateral moraines and some collapse of bedrock has taken place (Kirkbride and Warren 1999).
The thermal influence of the lake and its ability to
reduce basal shear stress and hence induce calving
will decrease or stop dramatically as the lake
approaches a length of 16 km and the terminus
becomes grounded (Fig. 29.21). However, with ice
thicknesses reaching 600 m in depth below the
Hochstetter icefall (Anderton 1975) grounding of
the ice front is unlikely. A more likely scenario is
Hochstetter ice cutting off Tasman ice, with this
arm of the glacier dwindling and developing a
second lake above and confined by Hochstetter
ice. An ice-dammed lake in this location has the
potential to create very dangerous jokulhlaup
floods, and therefore any such development would
have to be monitored carefully. Large rock
avalanches or landslides into the lake—a major
cause of glacier lake outburst floods in the Himalaya—is a further cause for concern at Tasman
Lake.
Apart from these possibilities, the hazards found
in some other regions of the world, where large
glacial lakes are forming, are unlikely to occur at
Tasman Lake. For example, a massive thick wedge
of impounding material immediately downstream
of Tasman Lake (a broad thick alluvial deposit
Remote-sensing case studies
called ‘‘alluvial fan-head impounding’’ by Kirkbride) makes a glacier lake outburst flood of the
type most common in the Himalaya practically impossible for Tasman Glacier. Though the lake is
deep, it is situated within an overdeepened basin
and is well contained by debris and bedrock
impoundments. A large landslide and resulting
tsunami would be the only significant outburst
mechanism. Comparatively few lakes on the West
Coast are situated such that small glacial lake outburst floods (GLOFs) may be possible, but these are
the exception in New Zealand. Other common
glacial lakes in New Zealand reside in stable bedrock basins.
A hint as to the complexity and massive development of the terminal sedimentary wedge is suggested by the adjacent Mueller Glacier (Harvie
2011), where interpretations of ground-penetrating
radar (GPR) data indicate the presence of a very
thick complex stratified sedimentary unit rather
than a simple impounding moraine. Thus, the idea
of a melting ice-cored moraine, or fluvial downcutting through a moraine, is not presently applicable to Tasman Glacier.
When climate perturbation forces mass loss at a
glacier, slow-responding tongues such as Tasman’s
at first exhibit little to no terminus retreat, but
rather lose mass by surface lowering (downwasting), with little or no length or area change. In
general, these glacier types maintain their LIA areas
even as climate change causes the ELAs to rise,
until a tipping point is reached when glacier ice
levels at the terminus lower to the river outlet level
allowing thermokarst sinkhole lakes to coalesce and
flow into the outlet river. The steep margins of
sinkholes, unprotected by debris cover, are then
attacked by limnological processes. Frontal calving
retreat of the glacier and disarticulation-type
breakup of the tongue takes place if the lake is deep
enough. The glacier retreats until a new equilibrium
is attained or approached. This type of mass loss
behavior, exemplified by Tasman Glacier, is typical
of large debris-covered glaciers around the world;
both steady or episodic lake growth proceeds
rapidly but is not directly a response to climate
change, though the process is normally set in
motion by climate change.
29.4.3.5
Energy constraints and discussion of
Tasman Lake’s growth rate
Tasman Lake’s growth, while getting close to being
defined by an exponential function, is not entirely
703
smooth, and obviously growth must end eventually.
However, predictions of future behavior based on
empirical curve fits of past behavior—whether
linear or nonlinear—are inherently unreliable if
they are not based on fundamental physics. Conversely, predictions based on physical models may
also prove unreliable because systems such as
Tasman Glacier and Tasman Lake are complex
and any simple phenomenology, while it may be
described analytically, is apt to exclude other
contributory phenomena, including complex feedbacks. Nevertheless, it is a worthwhile exercise to
consider the energy needed to melt the glacier’s ice
in order to achieve lake growth. In the early 1970s,
before Tasman Lake had formed, the ice in the
tongue reached an astonishing thickness of 600 m,
according to a seismic survey reported by Anderton
(1975). Where the tongue still exists, it has lost only
a little of its thickness since then. Across a great
length of the former glacier tongue, hundreds of
meters of ice thickness have been removed and
replaced by deep water (but with no mass equivalence implied).
By 1993, Tasman Lake—having just finished its
first decade of rapid growth—was already up to 125
m deep (Anderton 1975). Other young rapidly
growing lakes in the Mt. Cook area then had similar
maximum depths. Between 1995 and 2008, Tasman
Lake increased its volume by a factor of almost 4,
and by 2010 it had attained a volume of 510 10 6
m 3 (Dykes et al. 2010), when its area had grown to
about 6 km 2 , according to our measurements.
Thus, the mean depth of the lake is around 90 m.
The amount of ice lost from lake areas may average
closer to 180 m; this is estimated as the sum of the
90 m mean depth of the lake (submerged glacier),
plus the typical 60 m height of the calving front,
plus 30 m from thinning of the tongue between 1982
and 2007 (Thomas 2010); in the lake area, this 180
m of ice has been lost in just a quarter of a century
between the start of rapid lake growth around 1982
and 2007. By contrast, a profile line across the
glacier upvalley from the recent calving margin
thinned only by about 30 m in the 25 years from
1982 to 2007 (Thomas 2009), giving an average
thinning rate of about 1.2 m yr1 . Thus, in the area
of the present Tasman Lake, the loss rate of ice
volume per unit area (equivalent to thinning rate)
since 1982 has been of the order of 7.2 m yr1 , thus
showing the efficacy of limnological attack processes.
The entire solar radiation energy budget—averaging about 4.3 GJ m 2 yr1 for Tasman Glacier
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New Zealand’s glaciers
(de Vos and Fortuin 2010; see their fig. 2.1.4)—
could melt about 12,900 kg m 2 yr1 of ice, equivalent to about 14 m of ice annually. This simple
calculation does not include any sensible heat transfer, nor does not include heat that is lost by outflow
from Tasman Lake. However, the calculation suggests that Tasman Lake absorbs most incident solar
radiation (direct and indirect) and has converted at
least half of that into melting ice and growing the
lake basin. From this simple perspective—albeit a
highly incomplete one—Tasman Lake, as it grows
in area, should be capable of melting a greater
amount of ice each year; thus, lake growth should
accelerate. Therefore, in terms of its energy budget,
the lake area itself is a key component of the
system.
Despite parts of the glacier tongue being 600 m
thick, it is likely that lake/glacier dynamics may
effectively continue to degrade Tasman Glacier in
the next few decades, in much the same way as it has
already done in recent decades. Recession must
continue until a new equilibrium AAR is achieved
where the ablation zone area has diminished to a
size at which its mass loss comes into balance with
the reduced ice discharge supplied from the smaller
accumulation area above the higher snowline. Only
then can the glacier regain control of the lake size.
This re-equilibration might not be attained until the
glacier recedes so far that it detaches from the lake.
In the meantime, if the calving rate is correlated
with lake area as suggested above, then exponential
lake growth is possible, though other growth curves
may prevail if lake depth at the calving front is the
more important factor.
What actually will happen is likely to be more
complicated than a simple exponential model can
simulate. For one thing, as glacier tributaries
detach, the total amount of ice feeding glacier flow
will diminish; this factor could make the growth
rate of the lake irregular and perhaps accelerate it
even faster than would an exponential growth
curve. On the other hand, calving flux is apt to
diminish well before actual detachment occurs, such
that it is possible that lake growth will end not in
detachment but in a quasistable calving state
involving a much-shortened glacier. Furthermore,
additional lakes may form where glaciers detach,
and this may have knock-on effects by melting more
ice and thinning Tasman Glacier. Flotation and
disarticulation are apt to speed up the disintegration process, but surging could also occur
and actually lengthen the glacier and shorten the
lake for a time. There is too much complexity to
find the exponential growth model compelling,
and a linear projection of past behavior is as
good as more complicated models. Fig. 29.21 suggests that lake growth could follow a slower path,
with lake length doubling by 2045 from 2010’s
length. This model contrasts with the much faster
growth predicted in Fig. 29.20. In any case, dramatic changes are expected to continue for some
time to come, and reality is apt to confound predictions.
29.5
SPECIAL TOPICS
29.5.1 Debris production and debris
cover of New Zealand glaciers
A major characteristic of many of the larger New
Zealand valley glaciers is the mantle of bouldery
debris covering their lower tongues (Kirkbride
1989), which is readily seen in some of the satellite
imagery (e.g., Fig. 29.16). The percentage of debriscovered ice area to total glacier area in the Mt.
Cook area (the glaciers of Fig. 29.2) was 19.3%
for the western and northern glaciers at the start
of the ASTER/ETMþ imaging era, and 29.7% for
the eastern and southern glaciers (Chinn 2001). This
percentage difference is important dynamically.
Debris cover increases surface thermal inertia and
reduces the transmission of thermal energy to the
ice, thereby reducing the melt rates of the underlying ice by up to 90%, even at the relatively low
altitudes and high temperatures of the valley floors.
Where debris exceeds a thickness of a few centimeters, debris blankets become effective insulators;
debris cover is commonly up to 2 m thick towards
the glacier termini of large valley glaciers around
Mt. Cook. Referring back to the parameterized
definition of response time given by Jóhannesson
et al. (1989)—where response time equals the
thickness of ice in the ablation zone divided by
the ablation rate—then additional coverage by or
increased thickness of thin surficial debris cover
must increase the response time of glaciers. Hence,
debris cover reduces the response sensitivity
(increases the response times) of glaciers having
negative mass balances. However, debris cover still
allows a relatively rapid response to positive
balance changes if that response is accomplished
by a kinematic wave of ice transferred by glacier
flow (Kirkbride 1998). Effectively, this means that
debris-covered glaciers respond asymmetrically to
positive and negative mass balance inputs.
Special topics
In Chapter 13 of this book on the Chugach
Range, Kargel et al. examine some of the influences
of supraglacial debris on Alaskan glaciers, which
are applicable to New Zealand as well. Let’s suppose that debris-covered areas melt at 10% the rate
of relatively clean ice areas (a typical value), and the
ablation zone represents 60% of the total glacier
area. Assuming now that all other factors are equal,
the 29.7% debris abundance of south and east-side
glaciers would reduce ablation by about 45%; the
19.3% debris cover on west and north-side glaciers
would reduce ablation by 29%. The difference
between the two is 16%, which indicates that
spatially and temporally varying debris cover is
an important aspect of New Zealand’s glacier
dynamics. Debris cover is transported onto glacier
surfaces in a variety of ways. Landslides and other
mass-wasting processes, perhaps in New Zealand
triggered by moderate to high-magnitude earthquakes, sometimes deliver debris directly and suddenly onto glacier surfaces, and thereby more
abruptly affect glacier ablation and dynamics.
Glacier flow, ablation, and upwelling flowlines in
the ablation zone also transport subglacially eroded
debris to the surface, thus adding to surface debris
load. Common compressive flow in the ablation
zone can subsequently thicken the debris layer on
the glacier tongue. Hence, the slightly greater debris
abundance of eastern glaciers, if statistically significant, might be explained by either greater landslide
or subglacial erosional activity of east-side glaciers,
or similar debris production rates but greater ice
throughput and faster flow rates of west-side
glaciers. The latter is certainly true. Lithological
and mechanical property differences may be
another important factor.
Uplift rates in the Southern Alps are among the
fastest in the world. It is evident from the geomorphology of the diversity of glacierized areas in
today’s New Zealand that glaciation has locally
dominated erosional processes and has kept pace
with high levels of tectonic activity, as exemplified
by widespread surface landforms including cirques
and U-shaped valleys, lakes, moraines, fiords, and
massive outwash surfaces. As Brocklehurst and
Whipple (2004) showed, the geomorphic imprint
of glaciation on the landscape is varied, as represented by dramatically differing hypsometric curves
typical of New Zealand’s topography which are
generated by varied land ice processes such as
cirque glaciation (as on Mt. Ruapehu), valley
glaciers on Mt. Cook, and deep (Pleistocene) ice
cap glaciation in Fiordland.
705
Lateral strike-slip and oblique convergent fault
movements have averaged around 58 mm yr1
along the Alpine Fault, but there are variations
both along the fault and over time. On the basis
of 40 Ar/ 39 Ar closure dates, recent uplift rates may
peak at 6–9 mm yr1 on the Mt. Cook Massif (Little
et al. 2005). Uplift may reach 10 mm yr1 according
to Campbell and Hutching (2007) (with 20–30 mm
of horizontal movement). These rates are broadly
consistent with several prior estimates obtained
from different methods. Uplift near Mt. Cook is
countered by heavy precipitation–induced erosion
and removal of overburden, which can mean up to
35 km of removed rock in some places in the past
few million years along the Alpine Fault (Campbell
et al. 2012). As known in mountain geomorphology, uplift is generally not in true equilibrium or
steady state with erosion. Both of these are
dynamic, but over the long term they are commonly
closely linked. To a rough approximation, uplift
along the Alpine Fault does appear to approach a
balance with exhumation (Little et al. 2005), which
here has been driven mainly by glaciation at least
since the start of the Pleistocene. In turn, glaciation
is accelerated by extremely high precipitation rates
and high ice mass throughput, which is in part
elevation dependent and therefore linked to
regional tectonics.
Hicks et al. (1990) found the denudation rate on
Ivory Glacier was 5.6 mm yr1 , which is a bit less
than but comparable with uplift rates. Although
they found that annual precipitation more reliably
correlates with mountain denudation rates than the
extent of glacierization, the precipitation regime
(i.e., mean annual snow precipitation) is closely
correlated with ice mass throughput, which (as
known from studies elsewhere in the world) is the
underlying physical agency responsible for most
mountain erosion wherever glacierization is considerable (e.g., Chapter 13 of this book by Kargel et al.
on the Chugach Range). Thus, glaciation is likely
the primary controlling mechanism that limits the
heights of mountains in New Zealand. Where
glacier mass throughputs are highest, debris production rates are high; despite high conveyance
rates due to large mass throughput, many of Mt.
Cook’s glaciers remain heavily debris laden, a testimony to the extraordinary debris production
rates. By contrast, some glaciers with high debris
production rates can have relatively light debris
loads; this perhaps is due to rapid conveyance
and purging of the supraglacial and subglacially
eroded debris by rapid glacier flow.
706
New Zealand’s glaciers
The connection between hillslope processes and
glacier dynamics, and especially the evolution from
low-relief to high-relief dominated glacial landscape
evolution has been described for the Tibetan
Plateau and the Himalaya (Scherler et al. 2011); a
similar fast-paced evolution of landscape processes
also appears likely to be taking place in the Mt.
Cook area as glaciers retreat and thin, but debris
production proceeds rapidly.
The less maritime-influenced glaciers of the
eastern Canterbury mountains are found within
very friable and steeply dipping zones of Torlesse
greywacke. Subsequently, upper glaciers continuously receive rockfall and avalanche deposits,
ensuring copious quantities of supraglacial
debris.
In this highly unstable environment where earthquakes play a normal part of the erosion cycle, one
would expect to see evidence of high-magnitude
movements along the Alpine Fault releasing
massive landslides and initiating mountain collapses. In the schist and Torlesse greywacke zones,
the friability of the rock and tectonic activity of the
Southern Alps produce numerous rock avalanches
(Whitehouse 1983), which suggests that the widely
distributed supraglacial moraines of New Zealand
glaciers are almost entirely comprised of rock
avalanche and rockfall material masking major
individual rock avalanche events. However, unlike
the case of the Chugach Mountains in Alaska,
where earthquake-triggered landslides are an
important process (Kargel et al. 2009, Uhlmann
et al. 2012), to date there is no evidence of a massive
landslide disturbing the mass balance and forcing
the advance of a New Zealand glacier. However,
Shulmeister et al. (2009) suggested that a major
landslide was responsible for formation of one large
moraine at Franz Josef Glacier on the West Coast.
Furthermore, a unique single massive landslide
event did seriously affect the flow of a Pleistocene
glacier in southern Fiordland (Santamaria Tovar et
al. 2008); this was the enigmatic collapse of an
entire small mountain that was the source of the
Green Lakes, all the more surprising because of the
solid crystalline rocks of the region.
To the west of the Main Divide, glaciers cling
to steep slopes of friable schistose rocks. A 5 km
wide zone of mountainous terrain, located between
the Main Divide and the Alpine Fault, continues
to be squeezed upward; it is no surprise that precipitation and debris production rates reach extremely high rates. The fiords in this locality,
sculpted by Pleistocene glaciers, have long been
infilled by outwash gravels. In this steep topography the present glaciers typically debouche into
rocky gorges rather than onto terraced outwash
plains, as is common on the eastern side of the Main
Divide.
The schist terrain thickens towards the south
until the entire terrain from the West Coast through
to Otago is completely made up of schist. This
region is one of intermediate (for New Zealand)
precipitation, and therefore large valley glaciers in
this terrain carry less supraglacial debris than their
northern neighbors.
29.5.2 New Zealand glacier and
climate coupling
Despite the high precipitation gradient and other
climatic heterogeneities across the Southern Alps,
the series of ELA values of index glaciers indicate
that glaciers here tend to respond uniformly as a
single climatic unit. There is a high degree of correlation in the ELA time series of index glaciers (Clare
et al. 2002, Willsman et al. 2008). The zeroth-order
observation is that, viewed on centennial timescales,
New Zealand’s glaciers are in overwhelming retreat.
Although, when evaluated in greater detail, the
retreat trends of fast-response glaciers are clearly
modulated by decadal-scale climate oscillations,
and the retreat behavior in general is complicated
by differing response times resulting from the varied
characteristics of glaciers such as the presence or
absence of large terminal lakes. In this section we
examine centennial and decadal-scale climate trends
and oscillations for New Zealand as a whole, and as
expressed subregionally. Although New Zealand
has many weather stations, we focus on seven
weather stations that have long-term and more
complete records, and then we compare some individual station data with other station data.
The retreat behavior of the largest debris-covered
glaciers (those with the longest response times as
defined by Jóhannesson et al. 1989) might well be
attributed to lingering responses to the end of the
LIA, or at least to the retreat that occurred in the
early 20th century and at the initiation of supraglacial lake formation. However, most of New
Zealand’s glaciers have response times ranging from
a few years to a few decades and so the end of the
LIA plays no role in explaining the last century of
retreat. Thus, we have to identify the sources of
post-LIA climate oscillations and trends to identify
the cause of both long-term recession and intermittent periods of advance.
Special topics
Small and spectacular readvances in all shortresponse glaciers have provided a rare opportunity
to assess glacier response times and dynamics. Subsequent to the end of the LIA, and in response to
warming climate, the mass loss observed within the
population of fast-response glaciers has lagged
behind warming by only a decade or so (Chinn
1999, Fig. 29.6). A few fast-response glaciers had
managed readvances before reversal of the Interdecadal Pacific Oscillation (IPO), and in 1948 and
1967 minor resurgences were recorded only at
glaciers with the highest sensitivity. Changes in
the sign of glaciological trends such as glacier
length, including those in 20th century readvance
events, are a positive indication that these glaciers
kept in phase with decadal-scale climate oscillations. Therefore, responses to climate changes
exhibited by short-response glaciers in New
Zealand clearly indicate close tracking of climate
changes and exclude any possibility of lingering
responses to the end of the LIA being involved.
Hence, these glaciers exhibit direct responses to
global warming related to increased atmospheric
greenhouse gases and oceanic oscillations. The
same cannot be stated, with confidence, for large
low-gradient debris-covered glacier tongues, such
as those of Tasman Glacier and Godley Glacier,
which integrate dynamical responses due to recent
climate change with lingering responses to the end
of the LIA, as well as nonclimatic phenomena
related to lake growth.
Atmospheric circulation patterns exert a strong
influence on glacier mass balances: positive balance
years appear to be associated with the dominant
southwesterly circulation anomaly; and negative
balance years with a strengthened northeasterly circulation anomaly (Fitzharris et al. 1997, Clare et al.
2002). These circulation anomalies and, hence,
glacier balance change correlate with oscillating
hemispheric, atmospheric, and oceanographic
system occurrences, such as IPO and El Niño events
(Fitzharris et al. 1992). Associated with these
hemispheric anomalies are strong teleconnections
between glacier fluctuation and other climatic and
oceanographic events across the Southern Hemisphere (Fitzharris et al. 1992, Tyson et al. 1997)
including of course New Zealand sea surface temperatures. For example, we show in Figs. 29.22 and
29.23 that the climate of New Zealand is closely
linked with South Pacific Basin climatic and
oceanographic indices and very slightly connected
to those of the North Pacific Basin.
The essential link between changing climate and
707
glacier dynamics is best observed in ELA shift
(Dyurgerov et al. 2009), a leading indicator of
changing glacier dynamics. Since terminus
responses lag behind ELA shifts, they tend to integrate and average out ELA shifts over a period of
time (roughly the response time, see Chapter 1 of
this book by Zemp et al.). ELA monitoring is
especially useful because ELA is affected by temperature, precipitation, and other climatic parameters that together are more relevant to glacier
dynamics than any one of these components alone.
If precipitation has remained constant, then temperature change deemed responsible for observed
lowering of ELAs may be estimated from the atmospheric lapse rate. Figs. 29.22 and 29.23 do not
portray multiple climatic and oceanographic
parameters simultaneously, but they do document
a record of oscillating climate systems that must
have an effect on ELA shifts.
Fig. 29.22A indicates that the joint trend of New
Zealand mean temperature and the Antarctic
Oscillation (AAO3) together with their interannual
oscillations correlate well for data since 1945. One
regional indicator related to ENSO, NINO4,4 oscillates inversely with New Zealand temperatures
(note that the NINO4 axis is reversed), but longterm trends are well correlated (once again note the
axis reversal). The long-term trend is an indicator of
global warming, but inversely correlated annual to
decadal-scale oscillations show that tropical sea surface temperatures, and especially ENSO, affect New
Zealand temperatures as well. Fig. 29.22D shows a
century-long trend of warming temperatures (about
0.95 K century1 ) as well as decadal-scale oscillations of several tenths of a kelvin. Thus, the strong
centennial warming curve is combined with higher
frequency decadal variability and interannual variability, where the latter noise-like variation remains
comparable in magnitude with the magnitude of
centennial scale warming. Hence, some winters
can be colder than average, despite overall warming.
Fig. 29.22E, however, shows a statistically significant difference between the East Coast, where
3
AAO is an atmospheric pressure anomaly involving
the 850 hPa geopotential height centered near 40–50 S
latitude versus that centered over the Antarctic. It thus
relates to southern midlatitude atmospheric transport,
including storm activity.
4
NINO4 is a subregional central/western tropical Pacific
sea surface temperature index, essentially a component of
ENSO.
708
New Zealand’s glaciers
Figure 29.22. Secular trends and oscillations in New Zealand’s mean temperature and climatic indices (temperature record available from NIWA). Figure can also be viewed in higher resolution as Online Supplement 29.14.
warming has taken place, and the interior and West
Coast, which have not significantly warmed in the
past century. Even so, each of the stations plotted
have fairly tight correlations in their interannual
temperature oscillations. Fig. 29.22F affirms that
each of the individual stations correlates with the
seven-station mean temperature anomaly, lending
credence to the climate data record and showing
again that there is a correlation between stations.
However, visual inspection of the scatter of data
points about the correlation lines shows that
climate is more wildly variable at interior and West
Coast stations.
Fig. 29.23 portrays the seven-station temperature
record in relation to time and five different climatic
and oceanographic indices. Each plot shows the
linear least squares best fit to the centennial-scale
trend as well as the 2 statistical uncertainty in the
trend’s slope and aggregate data mean, thus giving
a 95% confidence window for the actual trend
relative to plotted values. As stated above, the significance of the centennial-scale warming trend is
emergent above (or superposed on) oscillating
decadal-scale climatic and oceanographic indices.
If the centennial trend is expressed as a regional
component of global warming, then we may conclude that global warming is more important than
ENSO and all these other indices. However, interannual variability remains high; this fact is indicated by R 2 ¼ 30%; most temperature variability
consists of year-to-year fluctuations, with global
warming controlling just 30% of the variability.
Special topics
709
Figure 29.23. Correlations between New Zealand’s seven-station mean temperature and time and with five
oceanographic and climatic indices. Turquoise zones show 95% confidence intervals in linear trend and mean
(temperature records from NIWA). Figure can also be viewed in higher resolution as Online Supplement 29.15.
The next most important parameter (after time) is
the AAO, which explains 23% of annual temperature variability. The remaining indices, the
Southern Oscillation Index (SOI),5 ENSO, NINO4,
5
The SOI index is calculated using pressure differences
between Tahiti and Darwin. It relates to ENSO and
particularly sea surface temperatures north of Australia
and the strength of southern Pacific trade winds.
6
PDO is a measure of alternating differences between sea
surface temperature of the north Pacific and equatorial
Pacific; the temperature contrast fluctuates on timescales
of 10–30 years and is strongly associated with storm
activity in the north and northeastern Pacific basin. It
too is related to ENSO.
and the Pacific Decadal Oscillation index (PDO),6
explain variability progressively less, respectively,
with PDO being of marginal if any significance.
Thus, we find that New Zealand’s climate is largely
disconnected from that of the North Pacific Basin,
exhibits a moderate connection with the tropical
Pacific, and has a very strong connection with
mid-southern latitudes of the Pacific and the Antarctic. In this sense, New Zealand is very much a
circum-Antarctic region.
Fig. 29.22 specifically shows the close correlation
between the New Zealand temperature record and
the interannual and subdecadal-scale oscillations of
the various climate indices. As PDO is a North
710
New Zealand’s glaciers
Pacific Basin oceanographic/climatic index, New
Zealand has greater affinities with climate and
oceanographic indices defined for high southern
Pacific latitudes than for the North Pacific Basin.
However, decadal-scale oscillations are similar in
all these regions, as they are in the Himalaya, where
ENSO exerts teleconnection and modulation influences on the Indian Monsoon and westerlies. Thus,
any short-response glaciers would be expected to
exhibit similar decadal-scale oscillations within
those respective regions, but they may be out of
phase by a year or two between regions as a result
of asynchronicity of perturbing influences (e.g., different times of arrival of signals from ENSO in
different regions) and differing glacier response
times.
It is important to recall that temperature alone,
or any climate parameter for that matter, cannot
solely determine ELA shifts, and therefore our analysis offers a partial perspective only. Glaciologists
have described the current state and likely future
trends of glaciers by comparing their current ELAs
with their calculated ELAs. The departure of the
accumulation–area ratio (AAR) from the usual 0.5
to 0.7 range of most balanced glaciers provides a
direct measure of the extent of disequilibrium of the
glacier. This approach works well for New Zealand.
For example, the New Zealand glaciers that have by
and large kept their LIA areas are rightly seen to be
grossly out of equilibrium with current conditions;
some of these would be required to lose nearly half
of their areas to approach equilibrium with the
present climate, even if no further warming occurs.
29.6
CONCLUSIONS
A century of glacial recession in New Zealand has
shown large variability in the individual length,
area, and volume response of glaciers. Overall,
New Zealand’s changing glaciers are witness to
the powerful influence of global warming. However,
many glaciers are out of equilibrium with rapidly
changing climate. Satellite-based, airborne, and
field-based observations have all contributed to
our understanding of the glacier state and dynamics
in New Zealand. From a three-decade record of
annual ELA values, one of the more surprising
phenomena revealed was that, despite the massive
differences in climate between the wet windward
west and drier eastern sides of the Southern Alps,
the entire Alpine chain glaciologically behaves very
similarly; it does not show an east/west dichotomy
in glacier behavior when considering similar types
of glaciers. The record of frontal fluctuations is
enhanced by long records from two of the world’s
most sensitive and responsive glaciers (Fox and
Franz Josef ). In contrast, the Southern Alps also
contain some slow-response glaciers (Tasman
Glacier being the archetypal example) that are
undergoing a century-long process of lake development. Measurements show a high degree of variability in dynamical glacier responses to climate
shifts between individual glaciers, as well as some
climatic heterogeneity, but (as just mentioned) no
east–west dichotomy in glacier behavior. The high
retreat rates of lake-terminating glaciers do not
accurately reflect the influence of local climate
relative to the slower retreats typical of landterminating glaciers. The mechanisms of lake
growth and glacier retreat, though probably
initiated by climate warming, have become
decoupled from climate change and will likely proceed in coming decades regardless of how climate
changes in the same period. Similar physical processes exist at Alaskan glaciers (see Chapter 14 of
this book by Wolfe et al. on glacier lakes in the
Chugach Range, Alaska). Glacier terminus
response of fast-response glaciers measured over
decadal-scale periods is not indicative of global
warming or cooling but rather reflects ocean
dynamics and regional climate in the Pacific Basin
near New Zealand. Interannual variability remains
high but is damped by the 5 to 20-year response
times of fast-response glaciers. Debris cover is also
a major controller of glacier dynamics and may be
coequal with climate change as a cause of glacier
dynamical variability.
The degree of disequilibrium of New Zealand’s
large glaciers, some of which remain at or near
their 1890 lengths (but are now much thinner), is
now so large that they must continue to retreat
simply to reach equilibrium with the present
climate. However, climate is dynamic and thus
will drive sustained retreat. Some fast-response
small steeply sloping glaciers might have periods
of minor growth interrupting long-term retreat
due to short-term climatic fluctuations. Thus,
New Zealand’s overall pattern of glacier retreat
and thinning relates to global warming, but the
details for most glaciers are complex and cannot
be related to global warming in any linear or simple
way. The productive use of satellite image analysis
presented here and in other recent work in the New
Zealand region underscores the need for continued
high-temporal frequency satellite imaging of the
region.
References 711
29.7
ACKNOWLEDGMENTS
We thank Bruce Raup and Andreas Kääb for
constructive reviews, and NASA’s Cryosphere
Program for support making this work possible.
ASTER L1A, L1B, and higher level data were
obtained through the online Data Pool at the
NASA Land Processes Distributed Active Archive
Center (LP DAAC), USGS/Earth Resources
Observation and Science (EROS) Center, Sioux
Falls, South Dakota (https://lpdaac.usgs.gov/get_
data). ASTER data courtesy of NASA/GSFC/
METI/Japan Space Systems, the U.S./Japan
ASTER Science Team, and the GLIMS project.
29.8
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