Equilibrium-line altitudes of late Quaternary glaciers in the Southern Alps, New Zealand

QUATERNARY

RESEARCH

5,

Equilibrium-Line

27-47

(1975)

Altitudes

in the Southern

of Late Quaternary Alps, New

Glaciers

Zealand

STEPHEN C. PORTER Department of Geological Sciences and Quaternary Research Center, University of Washington, Seattle, Washington 98195 Received August

28, 1974

Equilibrium-line altitudes (ELA’s) of former glaciers in the Tasman River-Lake Pukaki drainage basin of the Southern Alps were reconstructed from glacial-geologic data on former ice limits by using an assumed accumulation-area ratio of 0.6 x 0.05. Late Holocene (Neoglacial) ELA’s were depressed 140 m below present levels, whereas those of four late Pleistocene ice advances were depressed 500 m (Birch Hill), 750 m (Tekapo), 875 m (Mt. John), and 1050 m (Balmoral). Reconstructed ELA gradients are approximately parallel to one another and range from 19 to 23 m km’. Although vertical movement on active faults and isostatic tilting due to deglaciation have both contributed to modification of reconstructed ELA gradients from their original values, the maximum resulting effect probably amounts to less than 2.0 m km’ and is undetectable from present data.

INTRODUCTION Expansion and recession of glaciers during the * Quaternary glacial ages occurred in response to vertical fluctuations of the snowline that were generated by worldwide changes of climate. Various estimates have been made of the magnitude of late Quaternary snowline fluctuations from glacial-geologic data in an effort to evaluate the difference in climate between the last glacial age and the present interglaciation, with most results having been derived from mountain systems in the Northern Hemisphere (Charlesworth, 1957, p. 652; Flint, 1971, Table 4A). Cited values are not always directly comparable because (a) different methods have been employed to calculate snowline depression and (b) basic data are of variable quality. Information on the magnitude of snowline fluctuations offers a potentially important input to models of Pleistocene climate; therefore, studies made through a wide range of latitude and longitude are desirable so that global patterns can be evaluated. This paper reports results of an investigation on South Island, New Zealand, in

which snowline positions have been derived for intervals of late Quaternary glacier advance in the highest part of the Southern Alps. In this study, calculations were made of the al&&de--of-the equilibrium line of former glaciers along a transect perpendicular to the main divide of the Southern Alps, and the results compared with present-day values obtained from field studies and aerial photographic interpretation. The equilibrium line is a climatesensitive parameter marking the locus of points on a glacier where net mass balance equals zero (Paterson, 1969, p. 31). On temperate mountain glaciers it corresponds closely to the annual firn limit, or snowline, at the end of the ablation season (Meier and Post, 1962). Fluctuations of equilibrium-line altitude (ELA) reflect changes in the mass balance of a glacier which, in turn, are a function of climate. Temperature and precipitation are widely regarded as the primary parameters controlling the altitude of the equilibrium line although obviously topography and other climatic factors also exert an influence. Because of the com27

Copyright o 1975 University Printed in the United States.

of Washington. ..

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in any form reserved.

28

STEPHEN

plex interaction of ablation- and accumulation-related processes at the glacier surface and of various potential positive and negative feedback effects, specific climatic causes for ELA fluctuations generally cannot be isolated (Meier, 1965). Without such knowledge, detailed climatic reconstructions based on ELA data are not feasible. Nevertheless, comparison of the calculated steady-state ELA of an existing glacier with that of its larger Holocene or Pleistocene predecessors can provide a measure of the gross difference in climate between the present and earlier periods of glacier advance. PREVIOUS ESTIMATES Willett (1950) made an estimate of Pleistocene snowline depression in the Southern Alps using the cirque-floor method and cited values ranging from 976 to 1433 m. However, his figures for present snowline, rather than being based on observed firn limits or mass-balance data, were deduced from climatic diagrams of Taylor (1926) and Zotov (1938). Moreover, the cirque-floor method is applicable only to areas of former cirque glaciation, and such areas commonly are too low to provide a direct basis for establishing present snowline altitude. In the heavily glacier-covered Southern Alps, firn limits of former valley glaciers generally lay well below the level of cirque floors at the source of glaciers. Wardle (1970) pointed out potential errors in Willett’s calculations and suggested that his resulting values were probably too large. PRESENT AND FORMER GLACIERS OF THE SOUTHERN ALPS Anderton (1973) identified 527 glaciers in the Southern Alps between Milford Sound and Arthur’s Pass. Their total area was measured as 810 + 4 km2. Nearly a quarter of the glacier-covered area of the mountain range and almost 40% of the total ice mass lies within the

C. PORTER

Tasman River-Lake Pukaki drainage basin. The mass is largely confined to a small number of large glaciers (Tasman, Hooker, Murchison, and Mueller). Tasman Glacier, the largest in the Southern Alps, is 28 km long, covers an area of 55 thickness of km2, reaches a maximum about 630 m, and has a mean thickness of about 270 m. However, most New Zealand glaciers are far smaller (median area 0.7 km2 ). The glaciers are concentrated principally along or close to the main divide (Fig. 1); both their number and size decrease southeastward away from the range crest. Isoglacihypses depicting the configuration of the glaciation limit (Qstrem, 1966) are approximately parallel to the main divide but show several southeastward lobations most likely related to distribution of precipitation (Porter, in press). The gradient of the glaciation limit ranges from 13 to 25 m km-l, being steepest in areas of highest relief. During Pleistocene glacial ages, the Southern Alps supported a vast glacier complex that stretched for 700 km through 4 degrees of latitude and averaged about 100 km in width (New Zealand Geological Survey, 1973). Major drainage basins heading along the main divide generated long, thick (>lOOO m) glaciers fed by numerous confluent tributary ice streams. Where topographic control was minimal, as along the coastal lowland west of the mountains, glaciers spread laterally to form a coalescent Piedmont ice system with individual lobes up to 50 km wide. Secondary ranges to the lee of the main divide supported valley and cirque glaciers which were both smaller and higher with increasing distance from the crest of the Southern Alps. Because few radiometric dates are available to assess the timing of glacier advances, there is no general agreement on the number of New Zealand glaciations and on their possible correlation with

QUATERNARY

SNOWLINE

POSITIONS

ON

SOUTH

ISLAND

FIG. 1. Map of central South Island showing distribution of existing glaciers (black), extent of ice during the last (late Otiran) glaciation (ticked line), and isoglacihypses (in meters) depicting the glaciation limit.

better-dated glacial sequences from the Northern Hemisphere. The last major advances of glaciers (late Otiran) west of the main divide are believed to have culminated between about 22,000 and 14C y. a. and, therefore, are 14,000 broadly correlative with advances of late Wisconsin age in North America (Sug-

gate, 1965; Flint, 1971). Subsequent fluctuations were smaller in magnitude and their resulting deposits have been referred to the Aranuian Stage (Suggate, 1961). At least two, and possibly three, major expansions of glaciers preceded the late Otiran advances, but the timing and correlation of these events are largely

30

STEPHEN

speculative (Suggate, Soons, 1973).

significant nonglacial interval preceded the Mt. John advance (Table 1). Wolds drift is considerably more weathered and TASMAN VALLEY GLACIAL modified than the younger formations SUCCESSION and is exposed mainly on interfluves or The Tasman River-Lake Pukaki drainbeneath younger drift beyond Mt. John age basin was selected for this study be- and Balmoral terminal moraines. Consecause it provides a transect from the quently, a very long nonglacial interval is main divide of the Southern Alps to the inferred to have separated its deposition southeast limit of glaciation in the Macfrom that of the Balmoral Formation. kenzie Basin, and because much of the Balmoral drift is considered more than complex moraine succession has been 36,400 + 3150 14C yr old on the basis of studied and mapped in detail (McGregor, a questionable date on organic matter 1963; Mathews, 1967; Gair, 1967; Manassociated with Balmoral deposits at the sergh, 1973; Burrows, 1973; G. D. Mannorth end of the Mary Range near Lake sergh and R. P. Goldthwait, unpub. data). Pukaki. A peat bed resting on Balmoral Furthermore, Tasman Glacier, at the (?) drift along Landslip Creek and overhead of the valley, is one of only two lain by Mt. John drift has a 14C age of glaciers in the Southern Alps for which 34,100 + 2750 yr. The pollen assemblage mass-balance data are available. in the peat indicates a cool-climate flora Pleistocene drift in the Mackenzie Ba- dominated by grasses, but not cold sin and its principal tributary valleys was enough to be regarded as a glacial maxisubdivided by Gair (1967) into the Wolds, mum. Both these dates are considered Balmoral, Mt. John, Tekapo, and Birch minimum limiting dates for the Balmoral Hill Formations and provisionally as- Formation (G.D. Mansergh, pers. comm., signed to the Otiran Glaciation. How1974). Tekapo drift is older than 11,950 ever, on the basis of comparative mor? 200 yr, as indicated by a date on fragphologic and weathering studies made ments of Discaria toumatou wood found during the present investigation, the Mt. in proglacial lake sediments that overlie John, Tekapo, and Birch Hill Formations the Tekapo moraine encircling the lower appear to be close in age, whereas Bal- end of Lake Pukaki. On the basis of this moral moraines are much more subdued, date and the general moraine sequence in eroded, and weathered, suggesting that a the Tasman Valley, the Tekapo FormaLimiting

1965;

C. PORTER

Dates for Quaternary

Time-stratigraphic units (Suggate, 1961; Suggate and Moar, 1970) Hawera Series Aranuian Stage

Gage and

TABLE 1 Rock-Stratigraphic Units of Glacial Origin in the Mackenzie Rock-stratigraphic units (Gair, 1967) Neoglacial Nonglacial

Otiran Stage ------_-_-_pre-Otiran

“Includes

stages

drift

Age (l*C

Basin

yr BP)

800-40”

interval

Birch Hill Formation Tekapo Formation Mt. John Formation Nonglacial interval Balmoral Formation Long nonglacial interval Welds Formation

lichen and tree ring ages, as well as historical

observations.

>9,520,

<11,900 > 11,950

> 36,400

QUATERNARY

SNOWLINE

POSITIONS

tion is tentatively correlated with drift of the Kumara 3 advance in Westland which Suggate and Moar (1970) have assigned an age of about 14,500-14,000 yr. A peat sample (I-1306) from a kettle in Birch Hill moraine in the upper Tasman Valley gave an age of 5120 + 140 yr (Gair, 1967), and wood from a probable mudflow diamicton overlying outermost Birch Hill drift in the same area is 5370 + 30 to 5590 + 30 yr old (QL-57, QL59). Radiocarbon dates for organic matter associated with moraines of probable Birch Hill age in the Cameron and upper Rakaia Valleys northeast of the Mackenzie Basin suggest that glaciers advanced between about 11,900 + 200 and 9520 f 95 y. a. (Burrows, in press). Neoglacial moraines fronting Tasman, Hooker, and Mueller Glaciers have been mapped and dated by Lawrence and Lawrence (1965) and by Burrows (1973). Radiocarbon, tree ring, and lichen dates, together with historical records, indicate that the latest major episode of glacier expansion began as early as the mid-12th century A.D. and was punctuated by glacier advances in the mid-13th century, mid-15th century, mid- to late-17th century, mid- and late-18th century, early, mid-, and late-19th century, and about 1930. Older, undated moraine remnants fronting Tasman Glacier and beyond Mueller Glacier at Foliage Hill and the Hermitage may record earlier episodes of glacier advance corresponding to those dated at ca. 2160-1570 and 24730 14C yr BP west of the main divide (Wardle, 1973). CONTEMPORARY ELA’S Few data are available concerning the mass balance of New Zealand glaciers, so firm statements about present-day ELA’s or their regional variations are not possible. Goldthwait and McKellar (1962) reported that the fim limit of Tasman Glacier was close to 1800 m, based on observations between 1957 and 1959. The New Zealand Ministry of Works ob-

ON SOUTH

ISLAND

31

tained vertical aerial photographs close to the main divide in the Mt. Cook region on 30 April 1971 near the end of the 1970-71 balance year. Comparison of firn limits discernable on the photographs with contours of 1:63,360-scale topographic maps (loo-ft contour interval) indicates that 1971 ELA’s of Tasman, Murchison, and Hooker Glaciers were at 1830 f 15 m, 1815 + 15 m, and 1845 f 15 m, respectively. Oblique area1 photographs taken in this same region on 14 April 1972 showed that firn limits at the end of the 1971-72 balance year were virtually identical to those of the previous year. From available photography, isolines of equal ELA could be drawn near the main divide, using data from the three main glaciers as well as from numerous small adjacent glaciers (Fig. 2). Insufficient information exists to derive a decadal range of ELA values, but the three measurements for Tasman Glacier noted here suggest that the ELA probably has fluctuated between about 1800 and 1850 m during the past 15 yr. Because that part of the glacier lying above the deadice zone has had a dominant negative balance during at least part of this interval (Goldthwait and McKellar, 1972), the steady-state ELA of the active portion of the Tasman Glacier probably lies near or below the lesser of these values. Data from the Mt. Cook region indicate that present-day ELA’s lie some 200 m lower than the glaciation limit. Consequently, the configuration of the glaciation limit, as depicted in Fig. 1, probably also closely approximates the pattern of contemporary ELA’s in the Southern Alps, if 200 m is subtracted from the indicated altitudes of the isoglacihypses. DETERMINATION OF FORMER ELA’S In order to calculate the ELA of a former glacier, detailed glacial-geologic mapping is needed to define former ice limits and, from them, to reconstruct glacier topography. Ice limits used in

32

STEPHEN

C. PORTER

PRESENT GLACIERS (1970-1972)

FIG. 2. Present glaciers in the upper Tasman River-Lake Pukaki drainage basin. Areas of extensive stagnant ice shown by stippled pattern. Median altitudes of east- and south-facing glaciers shown by solid contours (m). Dashed lines represent 1970-72 ELA’s (m).

QUATERNARY

SNOWLINE

POSITIONS

this study are based partly on unpublished mapping by R. P. Goldthwait and G. D. Mansergh in the lower Tasman Valley and Mackenzie Basin, and on my own supplementary field mapping and photo interpretation in the Ben Ohau Range, Gammack Range, Burnett Mountains, and upper Tasman Valley. Lateral moraines related to the Balmoral, Mt. John, and Tekapo advances can be traced almost continuously upvalley for nearly 45 km, and afford an excellent means of determining the area1 extent and surface gradient of former glaciers in the lower Tasman Valley. Birch Hill lateral moraines can be traced about 20 km along both sides of the main valley, and moraines of this age built by smaller glaciers generally are prominent and easily recognized landforms in tributaries beyond and adjacent to the main valley. Because Birch Hill and older lateral moraines of Tasman Glacier are largely restricted to valley walls south of Mt. Cook village, glacier limits in the headward parts of the drainage basin were inferred primarily from projected long profiles constructed from data in the lower and middle reaches of the valley system (Fig. 3). These were supplemented by glaciererosional features related to former ice limits, including truncated spurs and the transition from smoothed and abraded valley walls to frost-shattered bedrock interfluve crests. Because valley walls are steep and contours closely spaced, even rather large errors in the vertical positioning of ice limits in former accumulation areas will not change the total glacier-covered area by more than a few percent. Therefore, such errors are not a serious problem in calculating former ELA’s and generally can be ignored. A number of Neoglacial moraine sequences were studied directly in the field, but a majority of those in the drainage basin are relatively inaccessible and were mapped on recent aerial photographs and topographic maps. Those portions of

ON SOUTH

ISLAND

33

former glaciers lying above present-day firn limits are assumed to have had the same topography as the upper reaches of existing glaciers. As an initial step in topographic reconstruction, the median altitude of a former glacier was regarded as a first approximation of a steady-state ELA. This inferred relationship is based on information assembled for 20 temperate, middle-latitude glaciers in North America and Europe indicating that median altitude of glaciers having a normal area/altitude distribution, on average, lies within 50 m of. the ELA (S. C. Porter, unpublished data). Contours were drawn at 100 m intervals using mapped ice-limit data. Contour lines were placed normal to valley walls in the vicinity of the median altitude of the glacier, but were drawn progressively more convex toward the terminus and more concave toward the glacier head. For a large low-gradient glacier like the Tasman during pre-Neoglacial advances, such contouring is not likely to be in error by more than about 50 m, except near the terminus where contours are strongly bowed, providing the ice-limit data are of good quality. The error is likely to be even less in the case of small, steep glaciers, for the control points on opposing lateral moraines are more closely spaced and permit less subjectivity in contouring. Because the most critical part of a glacier for determining the former ELA is its middle reaches near the median altitude, where contours normally are perpendicular to valley walls, contouring inaccuracies will be at a minimum (< 30 m) and should not generate significant errors in calculated ELA’s. The method adopted utilizes an assumed accumulation-area ratio (AAR) to derive former ELA’s. The AAR is the ratio of the area above the equilibrium line (accumulation area) to the area of the entire glacier; as used here, it is regarded as a mean, or steady-state AAR

-

‘\ ‘.

- .

FIG. 3. Long profiles of late Quaternary glaciers in the Tasman River-Lake Pukaki drainage basin. Mapped ice limits vances are based on right-lateral moraines. Ice thickness of Tasman Glacier based on data reported by Anderton (1973).

200

600

IOOO-

2 1400. 5 * ,200.

_

_

_ _

-.

(unbroken

0

x0’

,

/

IO

,

/

/

/

lines) of pre-Neoglacial

/

,

/

20

ad

km

QUATERNARY

SNOWLINE

POSITIONS

related to a steady-state ELA. Meier and Post (1962) reported that AAR’s for glaciers in maritime northwestern North America generally fall between 0.5 and 0.8, and Grosval’d and Kotlyakov (1969) derived similar values (0.5-0.6) for the Tien Shan of central Asia. The mean AAR for 11 glaciers in northwestern North America for which steady-state ELA’s have been estimated is 0.6. Because latitude, climate, and range of glacier size are broadly similar in the Southern Alps and in the coastal mountains of southern Alaska, British Columbia, and Washington, AAR values may also be similar in the two regions. Limited data tend to support this inference, for AAR’s of Tasman, Murchison, and Hooker Glaciers in the 1970-71 balance year, ignoring extensive areas of stagnant ice below the active termini, were 0.65, 0.57, and 0.60. The 1971-72 AAR of Murchison Glacier also was about 0.6. Although AAR’s of former glaciers can no longer be measured, it is assumed in this study that under steady-state conditions late Quaternary glaciers in the Southern Alps had AAR’s of 0.6 f 0.05. Contour maps of former glaciers were used to construct area-altitude curves by planimetry, from which ELA’s could be determined directly or by interpolation, using the inferred AAR. In practice only the topography of the lower half of the glacier is required, together with the total glacier area. The method proved most suitable for steep, narrow glaciers between 5 and 10 km long. The massive trunk glacier that occupied the main valley during the Tekapo and earlier advances had source areas both along the main divide and in tributary ranges as much as 20 km southeast of the divide. Because ELA’s increased southeast from the crest of the Southern Alps, ELA’s of tributary ice streams must have stood progressively higher in that direction. However, the method employed provides only a single

ON SOUTH

ISLAND

35

ELA for the whole glacier and cannot detect possibly different ELA values for various tributary ice streams. This problem would be less serious in the Southern Alps if the principal trunk glaciers had flowed perpendicular to the main divide, but bedrock structure north of Mackenzie Basin has generated a strong north-south alignment of ridges and valleys and caused glaciers to flow at an acute angle to the regional ELA gradient. Glaciers less than about 3 km long also were not ideally suited to this method, for the scale of available topographic base maps (1:63,360) meant increasingly large potential errors in planimetry, the smaller the glacier. Therefore, for glaciers <3 km long, and with an approximately normal distribution of area with altitude, the median altitude of the glacier was regarded as a close approximation to the steady-state ELA. Maps of glaciers in the Tasman ValleyLake Pukaki drainage basin during the Balmoral, Mt. John, Tekapo, Birch Hill, and Neoglacial advances are shown in Figs. 4-8. Isolines showing the regional pattern of ELA’s for each advance were constructed using ELA values determined for independent glaciers. The small size of most glaciers during the maximum Neoglacial advance of the last few centuries necessitated using median altitudes of south- and east-facing glaciers as a proxy for ELA values. Comparison of the maps shows a general parallelism of isolines among the different episodes, but with ELA’s being progressively lower for successively older advances. This relationship is most clearly shown in a cross section constructed perpendicular to the main divide and normal to the ELA isolines (Fig. 9). ELA data is lacking close to the main divide for pre-Neoglacial advances because all glaciers in that zone coalesced as part of the main Tasman glacier system. Consequently, only independent ice streams in the southern part of the Ben Ohau Range, Gammack

BALMORAL

FIG. 4. Reconstructed moral advance. Contour lines (m).

topography of Tasman Glacier and smaller nearby glaciers during Bal interval on glaciers = 100 m. Isolines of equal ELA shown by dashed

MT.

FIG. 5. Reconstructed John advance. Contour lines (m).

JOHN

topography of Tasman Glacier and smaller nearby glaciers during Mt. interval on glaciers = 100 m. Isolines of equal ELA shown by dashed

TEKAPO

FIG. 6. Reconstructed topography of Tasman Glacier and smaller nearby glaciers during Tekapc advance. Contour interval on glaciers = 100 m. Iaolines of equal ELA shown by dashed lines (m).

FIG. 7. Reconstructed topography of Tasman Glacier and nearby smaller glaciers during Birch Ii11 advance. Contour interval on glaciers = 100 m. Isolines of equal ELA shown by dashed lines ml

40

STEPHEN

C. PORTER

NEOGLACIATION

---

0

5

lOk!T

J FIG. 8. Reconstructed topography of glaciers in Tasman River-Lake F’ukaki drainage basin during last major Neoglacial advance. Contour interval on glaciers = 100 m. Isolines of equal median altitude of east- and south-facing glaciers shown by solid lines (m).

QUATERNARY

SNOWLINE

POSITIONS

ON SOUTH

ISLAND

41

FIG. 9. Section across the Tasman River-Lake Pukaki drainage basin perpendicular to the main divide showing reconstructed ELA gradients for Pleistocene ice advances, gradients of median altitude of east- and south-facing glaciers at present and during Neoglacial advance, and glaciation limit (dotted line).

Range, and Burnett used to determine of ELA’s. TEMPORAL

Mountains could be the regional pattern

FLUCTUATIONS OF ELA’S The median altitude of present southand east-facing glaciers lies some 80 m below ELA’s of glaciers for the 1971-72 balance year. Assuming a similar separation of median altitude and ELA’s during the Neoglacial maximum, then Neoglacial ELA’s stood about 140 m below those of the present (Fig. 9; Table 2). During the late Otiran advances, ELA’s were depressed 875 m (Mt. John), 750 m

(Tekapo), and 500 m (Birch Hill) below present levels, whereas during the Balmoral advance, the difference was some 1050 m (Fig. 9; Table 2). The glaciation limit through this same region, calculated from altitudinal distribution of existing glaciers, lies about 200 m higher than present ELA’s (Fig. 9; Porter, in press). The derived ELA values can be plotted as a function of time to illustrate longterm trends of equilibrium line position during the late Quaternary (Fig. 10). The resulting picture is necessarily approximate because (1) the dating of the late Otiran advances is imprecise and the age of the Balmoral advance is unknown, and

42

STEPHEN

C. PORTER

contraction (Denton and Karlen, 1973) implying that ELA’s reached levels equal to or above those of today. The area covered by glacier ice at any given time in the Tasman River-Lake Pukaki drainage basin is largely dependent on the position of glacier equilibrium lines and the topography of the drainage basin. In Fig. 11, the area covered by glacier ice during each of the late Quaternary advances is plotted as a function of the difference between present and former ELA’s. From such a curve, estimates can be made of the increase or decrease in glacier cover that would result from long-term changes in the level of steady-state equilibrium lines. The three segments of the curve reflect three different conditions of glacierization. The steep upper part (inferred by dashed line) represents conditions characterized by ELA’s higher than today with consequent reduction in size and eventual disappearance of valley glaciers. The middle section reflects the presence both of large valley glaciers and of smaller cirque and tributary glaciers under conditions when ELA’s lie between those of the present and those of Birch Hill time. The flatter, lower portion of the curve marks conditions of substantially depressed ELA’s when most glaciers in the

TABLE 2 Area of Glacier Cover and Depression of ELA’S During Late Quaternary Ice Advances in the Tasman River-Lake Pukaki Drainage Basin Glaciercovered area (km2)

Period Present glaciers (1970-72) Neoglaciation Birch Hill Tekapo Mt. John Balmoral

187 300 589 1145 1305 1562

? f k + f +

AELA

5 9 18 34 39 47

(+50m)

0 - 140 - 500 -750 -875 -1050

(2) the level to which ELA’s rose between times of glacier advance is unknown. Nevertheless, the data suggest that in the Southern Alps the lowering of ELA’s during the recent Neoglacial ice expansion was about 0.16 that of a full glacial age, whereas during the Birch Hill and Tekapo advances ELA’s were depressed 0.57 and 0.86 as much as during the late Otiran maximum (Mt. John). The position of ELA’s during the postBirch Hill, pre-Neoglacial interval and during subsequent intervals of recession in the late Holocene could not be determined from available glacial-geologic evidence. In the Northern Hemisphere, intervals of higher tree line during the Holocene correlate with times of glacier

-1000

-l200

-

0

n

n

n

n

’ IO

c

*

“C

FIG. 10. Fluctuations

’

’

’ 20

YEARS

’

’

BP

n

’

’ 30

1

B

’ /Tima

rcalr

(x IO’)

of ELA plotted

as a function

of time.

unknown

I

QUATERNARY

SNOWLINE

POSITIONS

ON SOUTH

ISLAND

43

I! 800 -: 1 8 6004

I

I I

400-l

I

200

3 a Ii -200

a

-400 I

-1200 t I 0

200

400

1 600

600

1000

GLACIER-COVERED

FIG. 11. Glacier-covered tion of change in ELA.

AREA

area of the Tasman River-Lake

drainage basin coalesce into a single large glacier system occupying the Tasman Valley and its principal tributaries. ISOSTATIC AND TECTONIC FACTORS AFFECTING ELA GRADIENTS Isolines depicting ELA trends have nearly uniform spacing on Figs. 4-8 indicating broadly equivalent ELA gradients, ranging from 19 to 23 m km-‘, during successive ice advances. Although exact parallelism cannot be claimed because of potential error inherent in the method used, the apparent close parallelism of reconstructed ELA surfaces suggests that changes of climate leading to glacier advances caused ELA’s to be depressed rather uniformly throughout this region (Fig. 9). However, because the former ice cover was very thick near the main divide of the Southern Alps

1200

I 1400

1600

(km*)

Pukaki drainage

basin plotted

as a func-

and because this is a highly tectonic region, the possible effects of isostatic recovery and faulting on reconstructed ELA gradients should be considered. Mathews (1967) made an estimate of ice volume during the last glaciation along a transect across the Southern Alps from the Tasman Valley to the west coast. Ice thicknesses in the Tasman Valley were determined directly from altitudes of ice-margin features identifiable on topographic maps and presumably were comparable to those indicated for the Mt. John profile in Fig. 3. Ice thicknesses west of the Alpine Fault were computed using the relationship T = pgh sin (Y, where T = basal shear stress, p = mean ice density, g = acceleration due to gravity, h = ice thickness, and 01= surface slope of the glacier, taking 7,” = 0.5 bars. From a computed moving average ice thickness along this transect, the ap-

44

STEPHEN

proximate isostatic deflection was determined. Mathews concluded that the maximum isostatic rebound subsequent to the Pukaki (= Mt. John) advance (e.g., during the past 14,000 yr), minus the potential effect of the present ice load, should be about 100 ft (33 m) immediately southeast of the crest of the main divide. Isostatic tilting away from the axis of uplift might reach a maximum value of about 3 ft mi-’ (0.6 mkm-l). If the reconstructed ELA surfaces have been tilted by even this maximum amount, such an isostatic effect would not be detectable from the ELA data. A potentially more significant effect is that produced by vertical uplift of the Southern Alps along the Alpine Fault. Suggate (1963) assessed the character and magnitude of movement on this major tectonic feature and concluded that, despite the difficulty in accurately measuring vertical displacements on the fault, about 400 ft (120 m) of uplift has occurred in north Westland within the last 10,000 yr. This figure agrees with a separate estimate of at least 1000 ft (305 m) of uplift since the last glacial maximum (Bowen, 1954), judged to be about 25,000 yr old. If a mean uplift rate of 12 m per milennium is assumed for the section of the Alpine Fault immediately northwest of the Tasman Valley, then about 170 m of displacement should have occurred since the Mt. John advance. This would have the effect of decreasing the apparent Mt. John ELA gradient by a maximum of about 1.5 m km-’ in the lower Tasman Valley, some 30-40 km away from the fault trace, assuming that the block of crust between the Alpine Fault and the Mackenzie Basin moved as a single unit. Such a tilt is too small to detect from the ELA data at hand, and there is no indication that reconstructed ELA gradients are gentler with increasing age, as would be expectable if the assumed long-term rate of uplift is valid. Although apparent

C. PORTER

lack of divergence of successively older ELA gradients probably reflects the inadequacy of the data for determining slight tilting of former ELA surfaces, it might also be related to more complex tectonics than postulated. As much as 12 m of vertical displacement postdating the Mt. John advance is evident on the Ostler Fault along the south end of the Ben Ohau Range, and east-dipping outwash terraces of Mt. John age flanking the Ohau River have been back-tilted and dip west away from this same fault (Mansergh, 1973). Faults cutting Neoglacial moraines near Mt. Cook have been displaced vertically 3 m or more, indicating substantial local movements in the recent past. On both the Ostler Fault and the faults near Mt. Cook rocks west of the fault trace have moved up relative to those on the east. Such movement would have the effect of diminishing the maximum amount of tilt of former ELA surfaces to less than 1.5 m km-i. Therefore, although both isostatic movements and Neotectonic faulting have doubtlessly altered reconstructed ELA gradients, the total change probably has been less than about 2 m km-’ in the Tasman Valley area, an amount too small to detect by comparing available ELA data. COMPARATIVE DATA Although values for ELA depression in the Southern Alps compare rather closely with those obtained from some other middle- and lower-latitude glaciated regions, strict comparison of data is hampered by (a) the different definitions of snowline employed by various workers, (b) the variety of methods employed in snowline reconstructions, (c) the lack of consideration of possible isostatic and tectonic effects on reconstructed snowlines, and (d) the all too common practice of evaluating local snowline depression in only two dimensions, thereby ignoring present and former regional

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snowline (ELA) gradients. The latter factor is especially important in comparative analysis, for unless gradients are determined, calculated values of snowline depression can be in error by hundreds of meters. Many published studies of reconstructed Pleistocene snowlines are based on analysis of cirque-floor altitudes (e.g., Porter, 1966; Flint, 1971, Table 18A; Pewe and Reger, 1972), but such an approach results in only minimum values for snowline depression in situations where former glaciers extended beyond cirque basins. In some such cases, cited values are less than half the actual difference between the snowline levels of the present and those of the last glaciation. A further complication in interregional comparisons arises from potential errors in correlation. In few instances are alpine glacial deposits sufficiently well dated that one can be certain of the contemporaniety of two widely distant reconstructed snowline surfaces. Some apparent interregional differences in snowline depression may be at least partly a function of erroneous correlations. ELA studies in four areas of the Northern Hemisphere have been made using essentially the same methods employed in the New Zealand investigation. During the maximum ice advance of the last glaciation in the Cascade Range of Washington, probably approximately contemporaneous with the Mt. John advance in the Southern Alps, ELA’s were some 850 m lower than at present, whereas during a subsequent readvance, tentatively dated as close to 14,000 yr old, the depression was 750 m (S. C. Porter, unpublished data). During a late-glacial readvance that culminated close to 11,000 y. a. ELA’s were 550 m lower than today, and some 450 m lower than levels reached during Neoglacial advances of recent centuries. During two ice advances assigned to the penultimate glaciation, ELA’s were about 1120 and

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980 m below present levels. A complimentary study in the Mt. Jefferson area of the Oregon Cascades showed that maximum depression of ELA’s during the last glaciation amounted to some 900-950 m and that during the past several centuries the lowering was as much as 200 m (Scott, 1974). ELA depression in Swat Kohistan, northern West Pakistan, was about 900 m during the last glacial maximum and about 1000 m during the preceding glaciation (Porter, 1970). In the Colombian Andes, only 5” north of the equator, ELA’s were depressed 950 m at the last glacial maximum and 150 m during the greatest Neoglacial advance (Herd, 1974). A value of 900 m also was obtained by Osmaston (1965) for snowline depression during the last glaciation on Mt. Kilimanjaro in equatorial Africa, using a method similar to that employed here. During the Younger Dryas advance in Norway, between approximately 10,100 and 10,800 y.a., the snowline lay between 450 and 600 m lower than today (Anderson, 1968). In the Alps the snowline was depressed between 600 and 1000 m during several late-glacial advances, according to Heuberger (1968), but the times of the advances are not well dated; the latepostglacial (Neoglacial) snowline lay between 100 and 200 m below present. Reported studies are as yet too few and widely scattered to draw definitive conclusions about the magnitude of snowline fluctuations on a global scale during the late Quaternary. The oft-stated inference that snowline depression was less at high and low latitudes than in middle latitudes and less under continental climatic regimes than in moist maritime regions (Charlesworth, 1957, p. 652) needs reevaluation. This statement was based largely on early studies, like those of Klute (1921; Flint, 1957, Fig. 4-l), which used two-dimensional meridional profiles of present and reconstructed Pleistocene snowlines. Regional snow-

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line gradients, which trend approximately at right angles to these profiles and locally are quite steep, were ignored. Consequently, some cited values of snowline depression could be in considerable error. As a first step, the global pattern of snowline depression might be determined most easily for the last glacial maximum, inasmuch as ice limits for this period generally can be identified readily. Data cited above suggest that in middle and lower latitudes, at least, snowline depression at that time commonly amounted to about 900 + 50 m. If significant deviations from this value were recognized in other areas, such information would constitute an important input for reconstructions of global climatic conditions at the climax of the last glaciation. ACKNOWLEDGMENTS This investigation was carried out during 1973-74 at the University of Canterbury, Christchurch, New Zealand, under a Fulbright-Hays Senior Research Fellowship. Richard Richardson assisted during part of the field studies, and I benefited from many fruitful discussions with Peter W. Anderton, Ian Brookes, Colin J. Burrows, Maxwell Gage, Richard P. Goldthwait, Graham D. Mansergh, and R. P. Suggate. I am grateful to John T. Andrews, Richard P. Goldthwait, and Graham D. Mansergh for critical reviews of the manuscript, and to Minze Stuiver for processing three radiocarbon samples relating to this study. REFERENCES Anderson, B. G. (1968). Glacial geology of western Trams, North Norway. Norges Geologiske Undersokelse 256, 160 p. Anderton, P. W. (1973). The significance of perennial snow and ice to the water resources of the South Island, New Zealand. New Zealand Journal of Hydrology 12, 6-18. Bowen, F. E. (1954). Late Pleistocene and Recent vertical movement at the Alpine Fault. New Zealand Journal of Science and Technology 35B, 390-397.

C. PORTER Burrows, C. J. (1973). Studies on some glacial moraines in New Zealand-2. Ages of moraines of the Mueller, Hooker, and Tasman Glaciers (S79). New Zealand Journal of Geology and Geophysics 16,831-855. Burrows, C. J. (in press). Late glacial and Holocene moraines of the Cameron Valley, Arrowsmith Range, Canterbury, New Zealand, Arctic and Alpine Research, 7. Charlesworth, J. K. (1957). The Quaternary Era, London, Edward Arnold. 2 vols., 1700 PP.> Denton, G. H., and Karlen, Wibjiirn (1973). Holocene climatic variations-their pattern and possible cause. Quaternary Research 3, 155-205. Flint, R. F. (1957). Glacial and Pleistocene Geology, 553 pp., New York, John Wiley and sons. Flint, R. F. (1971). Glacial and Quaternary Geology, 892 pp., New York, John Wiley and Sons. Gair, H. S. (1967). Sheet 20, Mt. Cook (1st Ed.), Geol. Map of New Zealand (1:250,000): New Zealand Geological Survey. Gage, Maxwell, and Soons, J. M. (1973). Early Otiran glacial chronology-a reexamination. IX INQUA Congress Abstracts, Christchurch, New Zealand. 111-112. Goldthwait, R. P., and McKellar, I. C. (1962) New Zealand Glaciology, in Antarctic Research. Geophysical Monograph American Geophysical Union 7, 209-216. Grosval’d, M. G., and Kotlyakov, V. M. (1969). Present-day glaciers in the USSR and some data on their mass balance. Journal of Glaciology 8, 23-50. Herd, D. G. (1974). “Glacial and volcanic geology of the Ruiz-Tolima volcanic complex Cordillera Central, Colombia.” Unpublished Ph.D. dissertation, University of Washington. Heuberger, Helmut (1968). Die Alpengletscher im Spat- und Postglazial: Eiszeitalter und Gegenwart 19, 270-275. Klute, Fritz (1928). Die Bedeutung der Depression der Schneegrenze fur eiszeitliche ProbZeitschrift fiir Gletscherkunde 16, leme. 70-93. Lawrence, D. B., and Lawrence, E. G. (1965). Glacier studies in New Zealand. Mazama 47, 17-27. Mansergh, G. D. (1973). Quaternary of the Mackenzie Basin. In “IX INQUA Congress Guidebook for Excursion 7, Central and Southern Canterbury of New Zealand,” pp. 102-111. Mathews, W. H. (1967). Profiles of late Pleistocene glaciers in New Zealand. New Zealand Journal of Geology and Geophysics 10, 146163.

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McGregor, V. R. (1967). Holocene moraines and rock glaciers in the central Ben Ohau Range, south Canterbury, New Zealand. Journal of Glaciology 6, 737-148. Meier, M. F. (1965). Glaciers and climate. In “The Quaternary of the United States” (H. E. Wright, Jr. and D. G. Frey, Eds.), pp. 795806. Princeton, Princeton University Press. Meier, M. F., and Post, A. S. (1962). Recent variations in mass net budgets of glaciers in western North America. IUGG/IASH Committee on Snow and Ice, General Assembly, Obergurgl, International Association of Science and Hydrology 58,63-77. New Zealand Geological Survey (1973) Quaternary geology-South Island, 1:1,000,000 (1st ed.): New Zealand Geologic Survey Miscellaneous Series Map 6. Osmaston, H. A. (1965). “The Past and Present Climate and Vegetation of Ruwenzori and its Neighborhood,” Oxford University, Worcester College, D. Phil. thesis. ostrem, Gunnar (1966). The height of the glaciation limit in southern British Columbia and Alberta: Geografiska Annaler 48(A), 126-138. Paterson, W. S. B. (1969). “The Physics of Glaciers,” 250 pp., Pergamon Press. Pew&, T. L., and Reger, R. D. (1972), Modern and Wisconsinan snowlines in Alaska. Proceedings of the 24th International Geologic Congress, Section 12,187-197. Porter,‘S. C. (1966). Pleistocene geology of Anaktuvuk Pass, central Brooks Range, Alaska. Arctic Institute of North America Technical Paper 18, 100 p. Porter, S. C. (1970). Quaternary glacial record in Swat Kohistan, West Pakistan. Geological Society of America Bull 81, 1421-1446.

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Porter, S. C. (in press). Glaciation limit in New Zealand’s Southern Alps. Arctic and Alpine Research 7. Scott, W. E. (1974). “Quaternary Glacial and Volcanic Environments, Metolius River Area, Oregon,” Unpublished Ph. D. dissertation, Univ. Washington. Suggate, R. P. (1961). The upper boundary of the Hawera Series. Transactions of the Royal Society of New Zealand l, ll-16. Suggate, R. P. (1963). The Alpine Fault. Transactions of the Royal Society of New Zealand 2,105-129. Suggate, R. P. (1965). Late Pleistocene geology of the northern part of South Island, New Zealand. New Zealand Geological Survey Bull 77, 91 pp. Suggate, R. P., and Moar, N. T. (1970). Revision of the chronology of the late Otira glacial. New Zealand Journal of Geology and Geophysics 13, 142-746. Taylor, Griffith (1926). Glaciation in the southwest Pacific. Proceedings of the Third Pan-Pacific Congress, pp. 1819-1825. Wardle, Peter (1970). Pleistocene snowlines in the Fox Glacier area. New Zealand Journal of Geology and Geophysics 13, 560. Wardle, Peter (1973). Variations of glaciers of Westland National Park and the Hooker Range, New Zealand. New Zealand Journal of Botany 11, 349-388. Willett, R. W. (1950). The New Zealand Pleistocene snowline, climatic conditions, and suggested biological effects. New Zealand Journal of Science and Technology 32B, 18-48. Zotov, V. D. (1938). Some correlations between vegetation and climate in New Zealand. New Zealand Journal of Science and Technology 19, 474-487.