Quaternary Science Reviews 20 (2001) 1067}1091
Snowline depression in the tropics during the Last Glaciation
Stephen C. Porter*
Department of Geological Sciences and Quaternary Research Center, University of Washington, Seattle, WA 98195-1360, USA
Abstract
Five primary methods have been used to reconstruct Pleistocene snowlines or equilibrium-line altitudes (ELAs) in the tropics
(23.53N}23.53S) during the last glaciation, but each has inherent errors that limit the accuracy of the results. Additional potential
errors in determining ELA depression involve estimates of modern snowline altitude, dating resolution, topographic reconstruction of
former glaciers, orographic e!ects, the presence of rockfall debris on glaciers, and calculation of regional ELA gradients. Eustatic
sea-level lowering during the last glaciation is an additional factor in#uencing estimates of ELA depression (!ELA). In cases where
modern snowline lies above a mountain summit, only a minimum value for !ELA can be obtained. At 12 tropical sites in Africa, the
Americas (to 103S latitude), and Paci"c islands, estimates of average !ELA range from 440 to 1400 m, but most fall in the range of
800}1000 m (mean $1""900$135 m). In a regional study of ELA depression in the southern tropical Andes (8}223S), an average
!ELA of ca. 920$250 m has been reported. Based on the assumption that glacier mass balance was controlled solely by
ablation-season temperature, and assuming a full-glacial temperature lapse rate of !63C/km, depression of mean annual temperature in glaciated alpine areas was ca. 5.4$0.83C. If adjusted for a sea-level fall of !120 m at the glacial maximum, this value is
reduced to 4.7$0.83C. The "gure is based on the (unlikely) assumption that accumulation on alpine glaciers has been invariant;
nevertheless, it is similar to values of temperature depresson (5}6.43C) for the last glaciation obtained from various terrestrial sites, but
contrasts with tropical sea-surface temperature estimates that are only 1}33C cooler than present. ! 2001 Elsevier Science Ltd. All
rights reserved.
1. Introduction
Recurring questions regarding the magnitude of tropical climate change during the last glacial age have
emerged since publication of the CLIMAP Project Members (1976, 1981) reconstruction of ice-age sea-surface
temperatures (SSTs). The CLIMAP reconstruction,
which focused on the last glacial maximum (LGM) revealed large areas in the tropics to have had SSTs as
warm as, or even slightly warmer than, those of
the present. The CLIMAP project considered the
LGM to date to 18,000 !"C yr BP [21,648 (21,484)
21,313 cal yr BP; equivalent calibrated ages ($1") have
been obtained using CALIB 3.03 (Stuiver and Reimer,
1993) for ages (18,000 yr, and using Stuiver et al. (1998)
for ages '18,000 yr]. Rind and Peteet (1985) subsequently noted con#icts between ice-age paleotemperatures generated by a general circulation model
* Tel.: #1-206-543-1904; fax: #1-206-543-3836.
E-mail address: [email protected] (S.C. Porter).
experiment, which used the CLIMAP SSTs as boundary
conditions, and low-latitude terrestrial paleoclimate
proxy data. Their analysis employed pollen evidence and
estimates of snowline depression from four tropical sites
(Hawaii, the Colombian Andes, equatorial Africa, and
New Guinea), and led them to conclude that the
CLIMAP reconstruction underestimated the amount of
tropical temperature depression, which likely amounted
to 5}63C.
Renewed interest in this topic has been generated
by evidence and modeling that point to colder tropical
temperatures than those implied by the CLIMAP
reconstruction (e.g., Guilderson et al., 1994; Stute et al.,
1995; Thompson et al., 1995; Bush and Philander, 1998;
Farerra et al., 1999). Basic to much of the discussion about colder glacial-age tropics has been the
question of snowline depression (e.g., Broecker, 1995;
Hostetler and Mix, 1999; Lee and Slowey, 1999), yet
most of the snowline data used in the arguments has not
been rigorously evaluated. Since Rind and Peteet (1985)
questioned the CLIMAP conclusions more than a decade ago, additional information has emerged that now
permits a more thorough assessment. This paper focuses
0277-3791/01/$ - see front matter ! 2001 Elsevier Science Ltd. All rights reserved.
PII: S 0 2 7 7 - 3 7 9 1 ( 0 0 ) 0 0 1 7 8 - 5
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S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
on the set of paleoclimatic data bearing on snowline
depression in the tropics during the last glaciation. In
assessing these data, potential sources of error are also
considered in arriving at site-speci"c and globally averaged values.
1.1. Glacier equilibrium-line altitudes
The snowline, de"ned as the lower limit of perennial
snow on the landscape, is equivalent to the "rn limit
on temperate alpine glaciers, which is the lower limit
of snow at the end of the ablation season. On such
glaciers, the "rn limit approximates the equilibrium line,
the locus of points along which the annual mass balance
is zero. In most recent paleosnowline studies, the equilibrium line is regarded as synonymous with the snowline,
and its altitude, following Meier and Post (1962), is
designated the equilibrium-line altitude (ELA). The difference between the modern ELA (ELA ) and that of
!
some earlier time (e.g., the last glaciation, ELA ) is
"
a measure of equilibrium-line (i.e., snowline) depression
(!ELA).
The mass balance of a glacier, and #uctuations of the
glacier's equilibrium line, are controlled by a number
of climate-related processes. For most low-latitude
temperate glaciers, the most important controls are accumulation-season precipitation and ablation-season temperature. Together these parameters encompass a range
of possible conditions controlling the ELA. Therefore,
a unique value for past precipitation or temperature
cannot be derived from the !ELA alone (Porter, 1977;
Seltzer, 1994). In most published paleosnowline studies,
no di!erence in precipitation is assumed (in most cases
probably erroneously) between the present and the
LGM, and a change in temperature is obtained by assuming a "xed atmospheric lapse rate. In cases where
independent evidence for one parameter (i.e., LGM precipitation or temperature) is available from another climate proxy, then !ELA can provide an estimate of the
other parameter.
2. Methods
In studies of glaciated
"ve common methods have
former ELAs. Because the
approach, the results they
comparable.
low-latitude mountains,
been used to reconstruct
methods di!er in their
produce are not strictly
2.1. Cirque-yoor altitude
When a glacier just "lls a cirque, its steady-state ELA
typically lies not far above the average altitude of the
cirque #oor (Fig. 1a). Therefore, cirque-#oor altitude has
sometimes been used as a convenient proxy for former
ELAs (e.g., PeH weH and Reger, 1972; Nogami, 1972, 1976;
Fox and Bloom, 1994). While this approach is reasonable
in situations where Pleistocene glaciers terminated at
cirque thresholds, in such cases the cirque glaciers disappeared when snowlines rose above cirque levels at the
end of the Pleistocene, meaning that site-speci"c ELA
depression cannot be calculated directly. Furthermore, in
many glaciated tropical mountain ranges and on large
volcanoes, glaciers expanded beyond cirques to form
valley glaciers, and in these circumstances ELAs lay
below (often well below) the altitudes of cirque #oors. In
such cases, snowline reconstructions based on cirque#oor altitudes may substantially underestimate actual
snowline depression.
2.2. Upvalley limits of lateral moraines
For a glacier in a balanced (steady-state) condition,
the upvalley limit of its contemporary lateral moraines
lies at the equilibrium line, below which ice-#ow paths
are diverging and ascending. If lateral moraines of a
former glacier are well preserved, then the altitude
of their upvalley limits may closely approximate the
former ELA (Fig. 1b). Whereas this method has been
used with success in some areas (e.g., Andrews, 1975;
Mahaney, 1990), in many alpine regions lateral moraines
are absent or poorly preserved and at best provide only
lower limiting estimates for contemporaneous ELAs.
Meierding (1982) considered the highest lateral moraine
altitude to be the least reliable of several methods for
determining Pleistocene ELAs in the Front Range of
Colorado.
2.3. Glaciation threshold
The glaciation threshold (GT) for a speci"ed area (normally a 7.5# topographic quadrangle or its equivalent:
e.g., ca 60 km# at 453 latitude) is the mean altitude between the lowest mountain with a glacier on it and
the highest without (Fig. 1c). Although this method is
not applicable to isolated peaks, such as volcanoes,
it has proved useful for assessing regional snowline
trends across mountain ranges (e.g., "strem, 1966;
Porter, 1975, 1977; Rodbell, 1992). Studies have shown
that the GT essentially parallels the regional ELA trend,
but commonly lies 100}200 m higher (Meierding, 1982;
S. C. Porter, unpublished data). A limiting problem
when using the GT to determine Pleistocene snowline
depression is the need to identify and map former glacierized and nonglacierized peaks for a speci"c time (e.g.,
the LGM) throughout a rather broad region. Such extensive "eldwork normally is impractical, and so subjective
assessments of the extent and age of past glaciation
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
1069
Fig. 1. Common methods used to derive past equilibrium-line altitudes in the tropics. See text for details. Cirque-yoor method: The ELA of a cirque
glacier is inferred to lie above, but not far above, the cirque #oor (CF). However, if a glacier expands beyond the cirque threshold, the ELA will be lower
than the cirque #oor. Lateral-moraine method: The upglacier limit of a lateral moraine approximates the ELA of the glacier that constructed the
moraine. Glaciation-threshold method: The average altitude between the highest nonglacierized summit (Sn) and lowest glacierized summit (Sg) de"nes
the glaciation threshold (GT) in a restricted area. Altitude-ratio method: In the median-altitude variant of this method, the ELA lies midway in altitude
between the head of the glacier (A ) and the terminus (A ). In the terminus-head altitude ratio (THAR) approach, the THAR equals the ratio of the
#
$
altitude di!erence between the terminus and the ELA divided by the total altitude range of the glacier. The ELA can be estimated by adding the
altitude of the terminus to the product of the total altitude range and an assumed THAR. Accumulation-area ratio method: In using this method, an
accumulation-area ratio (AAR) is used, based on the ratio of the accumulation area (Sc) to the total area of the glacier (where Sa is the ablation area).
Empirical studies suggest that a steady-state (SS, when the mass balance"0) AAR of 0.65$0.5 is appropriate for most temperate, relatively
debris-free glaciers. The surface topography of the former glacier is reconstructed based on glacial-geologic data. From the glacier's area}altitude
distribution (here depicted as a cumulative curve) and an assumed AAR, an ELA value is obtained.
are usually based on analysis of topographic maps or
aerial/satellite imagery. Despite the inherent uncertainties and subjectivity involved (Meierding, 1982), this
method has proved useful in assessing snowline depression in some areas (e.g., the Cascade Range: Porter et al.,
1983, Fig. 4-15).
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S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
2.4. Altitude ratios
Use of the median altitude of a former glacier as
a proxy for past snowline altitude is based on the
empirical observation that the "rn limit on temperate
glaciers at the end of the ablation season often lies about
halfway between the head of a glacier and its terminus
(Fig. 1d). HoK fer (1879) used a variation of this approach
by computing the arithmetic mean of the altitude of
a glacier's terminus and the average altitude of the
mountain crest at the glacier's head. The median altitude
method, in theory, is easy to apply if adequate altitude
data are available (i.e., topographic maps with a resolution of ca 30 m or less, "eld data based on altimetry
measurements, or digital-elevation data), and if a former
alpine glacier had a normally distributed area vs. altitude
curve. Nevertheless, whereas determination of the lower
limit of a glacier based on end moraines or outwash
heads may be relatively straightforward, assigning an
upper altitudinal limit to the former glacier in the cirque
region is generally subjective and arbitrary. High, steep
cirque headwalls can lead to a potential range of estimates di!ering by tens to hundreds of meters.
The median altitude method assumes that the ratio of
a glacier's range in altitude above the equilibrium line to
the total altitudinal range of the glacier is 0.5. A variation
of this method has also been used in which the ratio
[termed the toe-to-head (i.e., terminus-to-head) altitude
ratio, or THAR] is some lower value (Fig. 1c). For
example, Meierding (1982) reported that THARs of
0.35}0.40 generated the most accurate results in the Colorado Front Range. The resulting ELAs were ca
100}150 m lower than those derived using the median
altitude (THAR"0.5).
2.5. Accumulation-area ratio
The accumulation-area ratio (AAR) of a glacier is the
ratio of the glacier's accumulation area to the sum of its
accumulation and ablation areas (Fig. 1e). Empirical
studies of modern glaciers have shown that under
steady-state conditions the AAR typically falls between
0.5 and 0.8 (i.e., 0.65$0.15) (Meier and Post, 1962),
meaning that the accumulation area occupies approximately two-thirds of the glacier's total area. In calculating
past ELAs using the AAR method, a steady-state condition is assumed and the glacier's extent and topography
are determined using glacial-geologic data such as lateral
moraines, erratics, and trimlines (Porter, 1981). An initial
(estimated) ELA is selected using the altitude ratio
method. Contours of the glacier surface are then drawn,
consistent with principles of glacier #ow (contours of
a glacier in a balanced state typically are concave upglacier in the accumulation area and convex in the ablation area, with the degree of concavity or convexity
increasing with increasing distance from the equilibrium
line). The area between each pair of successive contours is
then measured and used to generate a cumulative curve
that graphically displays the glacier's area/altitude distribution. Assuming a steady-state AAR of 0.65, the ELA
can be determined from the graphical plot. Error limits
are derived by assuming a range of AAR values
(e.g.,$0.05 or$0.10).
2.6. Comparison of methods
Meierding (1982) assessed the relative reliability of
various paleo-ELA methods based on data from the
Front Range of Colorado. Using "rst-order trend-surface
analyses, he found that the cirque-#oor, median-altitude,
and lateral-moraine methods had the greatest root mean
square error (RSME"97}148 m), whereas the THAR
("0.40) and AAR ("0.65) methods produced the most
consistent results (RSME"ca 80 m). A similar study in
Norway by Torsnes et al. (1993) also concluded that the
AAR method produced the most reliable results. The
only similar comparative study in low latitudes was made
by Osmaston (1989a) in his study of glaciated equatorial
African mountains. He concluded that a modi"ed version
of the altitude-ratio method gave the best results. Overall, the general lack of reliable topographic information
and detailed "eld mapping for many tropical glaciated
areas, as well as limited radiometric age control, means
that errors inherent in most of these methods may be
magni"ed at low-latitude sites.
2.7. Additional potential sources of error
In addition to di!ering results obtained from the several ELA methods outlined above, as well as the errors
peculiar to each method, several other sources of error
enter into the calculation of LGM snowline depression.
2.7.1. Altitude of the modern snowline
On many tropical mountains, glaciers are absent or the
modern snowline altitude is known only approximately.
Furthermore, in a time of generally warming climate, the
transient nature of the snowline means that values obtained from direct observations a decade or more ago, or
from topographic maps based on them, may underestimate the present snowline altitude. Where direct observational data are unavailable, modern ELAs obtained
from recent glacier maps or aerial photography and
employing the median altitude or AAR methods likely
o!er the best estimates. Nevertheless, errors of tens of
meters or more may result.
2.7.2. Age of the LGM glacial limit
In few cases has the limit of the last glaciation been
radiometrically dated in tropical mountains, and in no
instance has it been closely bracketed by dates. Therefore, synchrony of moraines likely built during the LGM
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
commonly is inferred, based mainly on relative-age criteria and the pattern of moraine sequences, and on inferred regional or global correlations.
Recent studies in several (nontropical) mountain
ranges report that glaciers advanced repeatedly during
the last glaciation (marine isotope stage 2), and that the
resulting moraines often are nested or closely spaced.
These ice advances typically date between ca 27,000
and 16,000 !"C yr BP [31,300 and 18,972 (18,876)
18,784 cal yr BP] (e.g., Phillips et al., 1990; Gosse et al.,
1995; Lowell et al., 1995; Swanson and Porter, 1999).
Although comparable moraine sequences may exist in
the tropics (e.g., New Guinea: Blake and LoK %er (1971);
Andes: Clapperton (1987) and Thouret et al. (1996)), none
has yet been closely dated. The actual age of an `LGMa
moraine may lie anywhere within this ca 12,500 yr range.
Nevertheless, the juxtaposition of such moraines implies
comparable snowline depressions ()50 m di!erence)
during these successive advances.
2.7.3. Paleoglacier reconstructions
A potential source of error in the AAR method is
associated with the topographic reconstruction of a former glacier, which is necessarily subjective. Below the
equilibrium line, terminal and lateral moraines and trimlines provide altitudinal constraints along a glacier's former margin, whereas above the former ELA little control
usually exists. Errors in circumscribing the accumulation
area tend to be minimal because of steep valley walls
upglacier from the equilibrium line; thus, an erroneous
altitude estimate for the glacier margin in this zone only
minimally a!ects the lateral extent of the accumulation
area.
Errors related to topographic reconstruction are minimized in the case of small glaciers with normally distributed area}altitude curves. Large, complex glaciers, and
those having a trend perpendicular to the regional ELA
gradient, may generate unreliable results.
2.7.4. Orographic ewects
Small glaciers con"ned to deep cirques on leeward
#anks of mountains, or shaded by steep mountain walls,
may persist at altitudes well below those of glaciers on
fully exposed slopes. In general, ELAs based on geometrically simple glaciers in exposed sites are likely to provide the most regionally consistent ELA values. In
addition, orographic e!ects (e.g., unequal exposure to
sun, unequal accumulation) may lead to a range of tens of
meters in the altitude of the "rn limit on a given glacier.
2.7.5. Anomalies resulting from a cover of rockfall debris
An extensive cover of rockfall debris can insulate an
alpine glacier and greatly reduce ablation (Clark et al.,
1994). Such glaciers tend to be relatively insensitive to
a warming climate and typically advance to lower altitudes than do nearby debris-free glaciers. Clark et al.
1071
(1994) suggested that the steady-state AAR on such glaciers might be reduced from ca 0.65 to as little as 0.10.
Estimates of paleo-ELAs for debris-covered glaciers
based on the AAR or THAR methods therefore may
result in anomalous ELA values and produce erroneous
regional ELA gradients.
2.7.6. ELA gradient
An error can be introduced in calculating !ELA if the
regional snowline gradient is not considered. If there is
no present or past ELA gradient and both the modern
and LGM ELAs are determined for glaciers on the opposite #anks of a mountain range, the !ELA on each #ank
will be the same (Fig. 2, !ELA ). However, consider the
!
common case where the modern and LGM ELAs are
determined for opposite #anks of a mountain (e.g., Porter, 1979) or mountain range (e.g., Porter et al., 1983),
across which there is a marked precipitation gradient. If
regional trend surfaces of present and past ELAS are
determined, then !ELA may be less than if no gradient
exists (assuming uniform lowering on both #anks) (Fig. 2,
!ELA ), or the !ELA on one #ank may di!er from that
#
on the opposite #ank if the modern and paleo-ELA
gradients converge or diverge (Fig. 2, !ELA ). Where
$%"
possible, therefore, trend surfaces of present and former
ELAs should be calculated and their di!erence determined in order to obtain the most reliable estimates of
ELA depression.
2.7.7. Adjustment for lowered sea level
In most ELA reconstructions, sea-level lowering at the
LGM is not considered. However, the eustatic fall of sea
level had the e!ect of raising the altitude of mountain
summits by the amount of the sea-level drawdown. Assuming that sea level fell ca 120 m (e.g., Fairbanks, 1989;
Bard et al., 1990; Rohling et al., 1998), this amount should
be subtracted from the calculated !ELA to obtain an
adjusted ELA depression with respect to the changing
world sea-level datum (Broecker, 1997). For example, if
the present snowline (ELA ) on a glacier lies at 3900 m
!
(Fig. 3) and the reconstructed ELA at 3000 m, then the
"
apparent !ELA"900 m. However, during full-glacial
time, the ELA lay at an altitude of 3120 m, rather than
"
3000 m. Therefore, the di!erence between the present
ELA (3900 m) and full-glacial ELA , (3120 m) is 780 m
!
"
("!ELA adjusted for sea-level fall, designated !ELA
%
in Fig. 3). The sea-level factor becomes relevant if !ELA
is used in conjunction with an atmospheric lapse rate to
estimate temperature depression during full-glacial time,
for it will reduce the estimate by ca 10}15% (see below).
3. Tropical mountain glaciers (23.53N}23.53S)
Data on full-glacial snowlines are available for 18
mountain areas in tropical Africa, Central and South
1072
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
Fig. 2. Where no modern (ELA ) or full-glacial (ELA ) snowline gradients exist, !ELA is the same on opposite sides of a mountain or mountain range
!
"
(!ELA ). Where ELA gradients exist, the !ELA may be less (!ELA ) or greater (!ELA ) than in the no-gradient case, or may di!er on opposite sides
!
#
$
of a mountain or mountain range (!ELA , !ELA ).
"
&
Fig. 3. When adjusted for fall of sea level from its modern level (SL ) to its full-glacial level (SL ), !ELA is reduced by an amount equivalent to the
!
"
sea-level fall (!ELA ). A fall in sea level of 120 m had the e!ect of raising the altitude of the summit and the reconstructed !ELA by this amount.
%
"
America, and several glaciated Paci"c islands (Fig. 4).
Some permit only minimum estimates of snowline
depression because the highest summit lies below
the modern snowline, but for more than half, the fullglacial snowline depression can be calculated. In addition, regional data on snowline depression have been
generated for the tropical Peruvian, Bolivian, and
Chilean Andes.
3.1. Africa
3.1.1. Ethiopian highlands
The highlands of Ethiopia, which reach altitudes of
more than 4000 m, are too low to intersect the modern
snowline, but the highest summits developed glaciers
during the last glaciation. The Simen Mountains (13314#
N) culminate in Ras Dejen (4543 m), the highest mountain in Ethiopia (Fig. 5a). Hurni (1989) mapped moraines,
cirques, and periglacial features that presumably date to
the last glaciation (no radiometric dates are available). Of
20 former glaciers, those that formed in NW- to NEfacing cirques terminated as low as 3760 m; those occupying S-facing catchments reached only as low as 4400 m.
Former ELAs were estimated on the basis of the median
altitude of the glaciers (using end moraines and the top of
cirque headwalls), as well as the altitude of the upper end
of lateral moraines, giving an average value of ca 4250 m.
Minimum !ELA was therefore 290 m (Table 1).
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
1073
Fig. 4. Map showing location of glaciated areas in the tropics where ELA and !ELA estimates have been obtained.
"
Glacial landforms (cirques, U-shaped cross-valley pro"les, moraines, striations) were used by Potter (1976) to
map the extent of a former ice cap ('140 km#) atop
Mt. Badda (7352# N; published altitudes are inconsistent
and range from 4350 to 4133 m), ca 160 km southeast
of Addis Ababa (Fig. 5b). End moraines in W-trending
valleys were noted as low as 3650 m. Lateral moraines,
inferred to be of last-glacial age, reach as high as 4000 m.
If a summit altitude of 4350 m is adopted, then minimum
ELA depression was 350 m. A comparable minimum
!ELA results if the median altitude method is used.
3.1.2. Kilimanjaro, Tanzania
Mt. Kilimanjaro (3305#S), the highest mountain in Africa, now supports (5 km# of glacier cover. This volcanic
massif includes two summits; the highest, Kibo (5895 m),
lies west of a lower peak, Mawenzi (5147 m) (Fig. 5b).
During a succession of glaciations, glaciers on the volcano expanded to cover ca 153 km# (Osmaston, 1989a).
End moraines of the last (`Maina) glaciation on Kibo
and Mawenzi reach as low as ca 3250 m.
To determine snowline depression on Kilimanjaro,
Osmaston (1989a) used a modi"cation of Kurowski's
(1891) method, which assumed that net accumulation is
a linear function of altitude and that the ELA lies at the
mean altitude of the glacier area. Osmaston included an
arbitrary weighting factor to take possible nonlinearity of
the accumulation trend into account, and concluded that
this approach, which he called the Altitude}Height}Accumulation (Alt}Ht}Acc) method, was likely to give
more reliable results.
Osmaston's (1989a) analysis disclosed an asymmetrical
distribution of glaciers, and an ELA that slopes gently
!
eastward on Kibo (5455}5360 m) and lies at ca 5030 m on
Mawenzi. ELA gradients slope west}northwestward
"
across Kibo (4540}4575 m) and eastward across Mawenzi
(4300}4240 m), leading to unequal values of !ELA on
di!erent sides of the mountain (Fig. 5c). These he attributed to complex meteorological in#uences. Osmaston
concluded that a !ELA of 770$60 m between the Main
glaciation and a Recent ice advance (i.e., middle-to
late-Neoglaciation) explained his results on most of
Mazwenzi and Kibo. The calculated rise in ELA since
!
the Neoglacial maximum (a minimum of 60 m) increases
the !ELA to at least 830$160 m (Table 1).
3.1.3. Ruwenzori, Uganda
The Ruwenzori Mountains (0320-25#N) reach altitudes
of more than 5000 m and contain many small glaciers
that collectively cover ca 4.5 km# (Fig. 5d). During the
latest (Lake Mahoma) of at least three Pleistocene
glaciations, ice covered ca 260 km# and terminated as
low as 2070 m on the eastern slope (Osmaston, 1989b).
A date of 14,750$290 !"C yr BP[17,981 (17,647)
17,315 cal yr BP] from Mahoma Lake, provides a minimum age for moraines that impound the lake (Livingstone, 1962, 1975).
Osmaston (1989b) calculated ELAs of present and
former glaciers using the Area}Height}Accumulation
method described above. He derived a su$cient number
of measurements to de"ne the regional ELA gradient,
which descends to the east}southeast. The ELA along
!
an approximately west}east transect across the range
descends from ca 4720 to 4270 m (Osmaston, 1989b,
Fig. 7b), the estimated average ELA being ca 4600 m.
!
During the Lake Mahoma glaciation the ELA sloped
"
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S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
Fig. 5. Sites in Africa where ELA and !ELA estimates have been obtained. (a) Ras Dejen, Simien Mountain, Ethiopia; (b) Mt. Badda, Ethiopia; (c)
"
Kilimanjaro, Tanzania; (d) Ruwenzori Mountains, Uganda; (e) Mt. Kenya, Kenya; (f) Mt. Elgon, Kenya}Uganda, and Aberdare Mountains, Kenya.
See text for details.
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
eastward from ca 4100 to 3600 m (Osmaston, 1989b,
Table 10). Along this W}E transect, the !ELA increased
from ca 620 to 670 m.
3.1.4. Mt. Kenya, Kenya
The glaciers of Mt. Kenya (5202 m; 0309#S) have been
shrinking in area and now cover (1 km# (Young and
Hastenrath, 1991). Modern ELAs are estimated to lie at
ca 4700}4725 m (Mahaney, 1990, Fig. 11.8) and glacier
termini at 4650$100 m. Moraines of the Liki II glaciation (Osmaston, 1989b; Mahaney, 1990, Table 11)
extend as low as 3200 m, and are older than 15,000
!"C yr BP [18,003 (17,916) 17,830 cal yr BP]. Osmaston
(1989b, Table 11), using the median altitude method,
calculated an ELA of 4200 m for the glaciers originating
"
on the highest summit (Batian), and a !ELA of 600 m
(Osmaston, 1989b, Table 12). Mahaney (1990, Fig. 11.8),
on the other hand, estimated that the full-glacial ELA,
based on lateral moraine altitudes, lay between ca 3680
(SW slope) and 4000 m (NW slope) (Fig. 5e); most likely it
averaged close to ca 3700 m. Based on his data, the
!ELA ranged between ca 725 and 1020 m. The discrepancy between Mahaney's results (Fig. 5e, Table 1) and
Osmaston's may largely re#ect the di!erent methodologies used.
3.1.5. Other glaciated African summits
Two additional low-latitude mountains, each lying below the modern snowline, developed large glaciers during
the last glaciation (Osmaston, 1989b, Table 11). On
Mt. Elgon (4320 m; 1330#N), which had 75 km# of ice at
the glacial maximum [prior to ca 11,000 !"C yr BP;
12,966(12,917)12,865 cal yr BP], moraines extend as low
as 3350 m. Osmaston (1989b) calculated that the ELA
"
lay at 3600}3900 m, which means a minimum !ELA of
420}720 m (Fig. 5f). In the Aberdare Mountains (4001 m;
0315}45#S), which had 23 km# of ice cover at the glacial
maximum, LGM moraines reach as low as 3200 m.
The calculated ELA is 3700 m and !ELA was '300 m
"
(Fig. 5f).
3.2. Mexico and Central America
3.2.1. Mexican volcanoes
The high stratovolcanoes of Mexico's Cordillera
NeovolcaH nica display evidence of repeated Pleistocene
glaciations. Two of the best-documented records are
from IztaccmH huatl (5286 m; 19305}15#N) and Ajusco
(3937 m; 19312.5#N) (White, 1962; Heine, 1976, 1978,
1984; White and Valastro, 1984) (Fig. 10). Initially, chronologies of glaciation were based on relative-dating criteria, limiting radiocarbon ages (primarily of paleosols,
wood fragments, peat, and calcrete), and correlation with
glacial sequences elsewhere (Heine, 1984). Heine (1984)
reported evidence of glacier advances at ca 35,000
1075
and 12,000 yr BP [ca 39,800 and 14,111 (13,992)
13,880 cal yr BP], but none that correlated with the marine isotope stage 2 maximum. White (1981), using relative-age criteria, correlated the Hueyatlaco moraines
(Diamantes Substage, Second Advance) on IztaccmH huatl
and Santo TomaH s Substage moraines on Ajusco with
Pinedale moraines of the western United States that are
generally regarded as correlatives of the isotope stage
2 maximum (e.g., Richmond, 1986). Subsequently, White
and Valastro (1984) inferred that the Santo TomaH s drift is
more than 25,000 !"C yr (ca 29,000 cal yr) old, based on
a single radiocarbon date of a bulk sample from the
B horizon of a buried soil. The soil is developed on Santo
TomaH s till and lies stratigraphically below tephra on
which another buried soil is developed that dates to
15,090$150 !"C yr BP [18,186 (18,009) 17,832 cal yr BP].
Recent $'Cl surface-exposure ages for the outer Hueyatlaco moraines on ItaccmH huatl indicate that the main Late
Pleistocene advance probably culminated 19,000}18,000
$'Cl yr ago, and that the inner moraines are 15,000}
14,000 $'Cl yr old (VaH squez-Selem, 1998).
White (1981) calculated modern and past ELAs for
glaciers on the western slope of IztaccmH huatl and the
northern and eastern slopes of Ajusco. Based on mean
altitudes and an AAR of 0.65, he determined an average
ELA of ca 4880 m for glaciers on IztaccmH huatl and
!
3970 m for the Diamantes Second Advance, indicating an
LGM snowline depression of 910 m (Fig. 6a). For Ajusco,
which lies 65 km west of IztaccmH huatl, White calculated
that the ELA for the Santo TomaH s advance was 3270 m.
"
This is ca 155}170 m below ELAs calculated for two
Neoglacial ice advances and ca 665 m lower than the
present ice-free summit. Absence of glacier ice on Ajusco
(White and Valastro, 1984, Fig. 2) implies an ELA of
!
'3937 m, and a full-glacial !ELA of '665 m (Fig. 6b).
3.2.2. Altos de Cuchumatanes, Guatemala
Hastenrath (1974) reported evidence of glaciation in
the Altos de Cuchumatanes (15330#N), a carbonate karst
upland that reaches altitudes of nearly 3800 m (Fig. 6c).
An end-moraine complex, with up to 20 m of relief,
descends to 3470}3600 m. Hastenrath's published reconnaissance maps do not permit detailed topographic reconstruction of the former glaciers. He estimated that the
associated ELA lay at ca 3650 m, which is close to the
median altitude of the glaciated terrain. The corresponding minimum !ELA is 150 m. Although no dates are
available, Hastenrath considers the moraine characteristics comparable to those of presumed LGM age in the
mountains of Venezuela, Costa Rica, and Mexico.
3.2.3. Sierra de Talamanca, Costa Rica
In the Cordillera de Talamanca, culminating in Cerro
ChirripoH (9329#N; 3819 m) and now below the regional
snowline, two groups of moraines delimit former small
Taiwan Shan, Taiwan
Mauna Kea, Hawaii
Ajusco, Mexico
IztaccmH huatl, Mexico
Altos de Cuchumatanes,
Guatemala
Ras Dejen, Simen
Mountains, Ethiopia
Cerro ChirripoH , Costa Rica
Pico BolmH var, Venezuela
Mt. Badda, Ethiopia
Mt. Kinabalu, Borneo
Nevado del RumH z, Colombia
Nevado de Santa Isabel,
Colombia
Mt. Elgon, Kenya}Uganda
Ruwenzori, Uganda
Mt. Kenya, Kenya
Antisana, Ecuador
Aberdare Mountains,
Kenya
Chimborazo, Ecuador
Kilimanjaro (Kibo),
Tanzania
23330#N
19350#N
19312.5#N
19310#N
15330#N
13314#N
8}93N
7352#N
6305#N
4352#N
4348#N
1320#N
0320}25#N
0310#S
0330#S
0340#S
3305#S
1325#S
9329#N
Locality
Latitude
Table 1
Tropical snowline data
5895
6310
4001
5790
5202
4320
5109
4950
5200
4101
4350
&5000
3819
4543
3798
3937
5286
3997
4206
Altitude (m)
'3800
4880
4715
ELA &(m)
!
Median
altitude
Alt}Ht}Acc
Median
altitude
Alt}Ht}Acc
Alt}Ht}Acc
5360}5455
4540}4575
3880}4090
3700
'4001
4800}4900
3480}3860
3680}4200
3600}3900
3600}4100
3500}4000
3550}4000
3665
4000
3800
3500}3550
4250
3650
3270
3970
3350}3450
3780
ELA (m)
"
4600}4830
4700}4725
Median
'3800
altitude
Median
'4543
altitude
Median
'3819
altitude
Median
&4700
altitude
Lateral
'4350
moraines,
median altitude
Median
4570$150
altitude
Median
4800}4900
altitude
Median
4700}4800
altitude
Alt}Ht}Acc
'4320
Alt}Ht}Acc
4270}4720
AAR (0.65)
AAR (0.65)
Cirque #oor
AAR (0.60)
Method
830$60)
810}920
'300
970}1120
Yes
Yes
Yes
Yes
Yes
No
Yes
'420}720
620}670
725}1020
Yes
Yes
No
1000
1075
905$150 (?)
No
'350
No
'305
No
No
'290
&900
No
No
No
No
Yes
5a
6c
6b
6a
9a
9b
Figure
Hastenrath (1973) and
6d
Orvis and Horn (2000)
Schubert (1974, 1984) and 7a
Clapperton (1993)
Potter (1976)
5b
Hurni (1989)
Hastenrath (1974)
Ono (1988)
Porter (1979); Dorn et al.
(1991)
White and Valastro (1984)
White (1981)
Reference(s)
Clapperton (1987, 1993)
'12,000, (30,000
Osmaston (1989a)
Osmaston (1989b)
Osmaston (1989b)
Osmaston (1989b);
Livingstone (1975)
Osmaston, (1989b) and
Mahaney (1990)
Clapperton (1987, 1993)
'12,200
'12,000, (30,000
'15,000, (25,000
'11,000
'14,750$290
5c
6e
5f
6e
5e
5f
5d
Koopmans and Stau!er
9c
(1967)
'16,220$80
Herd (1974, 1982) Thouret 7b
('19,500,(23,000) et al. (1996)
Herd (1974, 1982)
7c
(12, 650$130
(19,080$820
'10,140$120
ca 21,000 ;
'15,090$150
20,300$2300;
18,900$800
ELA gradient Age control(
assessed
'150
'665
910
'400
'425 (935)
!ELA' (m)
1076
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
Mt. Albert Edward, Papua
New Guinea
Mt. Victoria, Papua New
Guinea
Cordillera Blanca, Peru
8325#S
8358#S
7340#S
Peruvian Andes
Peruvian Andes
Peruvian Andes
Peruvian}Bolivian Andes
Chilean}Bolivian Andes
Chilean}Bolivian Andes
Chilean}Bolivian Andes
103S
123S
143S
163S
183S
203S
223S
Median
altitude, cirque
#oor
Median
altitude, cirque
#oor
Median
altitude, cirque
#oor
Median
altitude, cirque
#oor
Median
altitude, cirque
#oor
THAR (0.2,
0.4), lateral
moraines
THAR (0.2,
0.4), lateral
moraines
THAR (0.2,
0.4), lateral
moraines
THAR (0.2,
0.4), lateral
moraines
Cirque #oors,
THAR (0.45)
Cirque #oors
Cirque #oors
Cirque #oors
THAR (0.45)
THAR (0.45)
THAR (0.45)
THAR (0.45)
Alt}Ht}Acc
3500}3550
3650}3700
3600}3650
3650}3700?
&4600
&4600
&4600
&4600
4700}5000
4700}5000
4500}5100
4500}5200
5100}5300
5400}5600
5400}5800
3400}4200
3400}4600
3600}4200
3200}4400
3200}4200
3800}4400
3800}4800
900}1200
700}1100
700}1000
550}1200
800}1100
950}1100
900}1100
600}1400
&4700}5300 3200}4900
1100}1350
900}1150
750}950
3150}3300
3540}3640
440}970
'335}385
'340}390
'470}520
'820}870
'850}1010
830$60)
3850}3900
4620
4620
4620
4200}4400
3500}3600
&4600
4985$120
4240}4300
5030
Yes
Yes
Yes
Yes
Yes
Yes
Yes
Yes
Yes
Yes
Yes
Yes
No
No
No
No
No
Yes
(12,100$190
'12,100$190
'12,100$190
'13,280$190
&Estimated values in italics
'Minimum values in italics
(Ages in !"C yr BP, except for $'Cl ages (in italics)
)Based on Osmaston's (1989a) estimated average value of 770$60 m, and corrected for a minimum 60 m rise of ELA since the Neoglacial maximum.
Peruvian Andes
Co. Oriental (local divide, &4500
W), Peru
Co. Oriental (main divide &4500
E), Peru
5}173S
7308}58#S
7308}58#S
'6000
4036
3990
4121
4368
4509
5147
Co. Oriental (main divide &4500
W), Peru
Sarawaged Range, Papua
New Guinea
6}6320#S
7308}58#S
Mt. Giluwe, Papua New
Guinea
Kilimanjaro (Mawenzi),
Tanzania
Mt. Wilhelm, Papua New
Guinea
6302#S
5345#S
3305#S
Klein
Klein
Klein
Klein
Klein
Klein
Klein
et
et
et
et
et
et
et
al.
al.
al.
al.
al.
al.
al.
(1999)
(1999)
(1999)
(1999)
(1999)
(1999)
(1999)
Fox and Bloom (1994)
Rodbell (1992)
Rodbell (1992)
Rodbell (1992)
Rodbell (1992)
LoK %er (1972)
LoK %er (1972)
LoK %er (1972)
LoK %er (1972)
LoK %er (1972)
Osmaston (1989a)
8
8
8
8
8
8
8
6f
6f
6f
6f
9d
9d
9d
9d
9d
5c
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
1077
1078
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
Fig. 6. Sites in Mexico and Central America where ELA and !ELA estimates have been obtained. (a, b) IztaccmH huatl and Ajusco, Mexico; (c) Altos de
"
Cuchumatanes, Guatemala; (d) Cerro ChirripoH , Costa Rica. See text for details.
valley glaciers (Weyl, 1956; Hastenrath, 1973) (Fig. 6d).
The upper basins of the main glaciated valleys lie at
3450}3550 m and contain several small moraine- and
rock-dammed lakes. A !"C date of basal sediments
from moraine-impounded Lago de Morrenas (3480 m)
indicates that the basin was deglaciated prior
to 10,140$120 !"C yr BP [12,112 (11,808) 11,123
cal yr BP] (Horn, 1993). At a site near El Empalme
(2400 m), ca 60 km northwest, the paH ramo (alpine) pollen
zone dates to 20,750 !"C yr BP (ca 24,200 cal yr BP) and
represents a treeline depression of at least 650 m (Martin,
1964); possibly, this date is close to the time of maximum
snowline depression as well. Assuming that glaciers
headed close to 3700 m (i.e., just below most crest altitudes) and terminated as low as ca 3300}3400 m (Hastenrath, 1973; Orvis and Horn, 2000), the full-glacial ELA
(median-altitude method) lay at ca 3500}3550 m.
This value is comparable to that recently derived by
Orvis and Horn (2000) of 3506}3523 m, and represents
a minimum ELA depression of ca 295}315 m (Fig. 6d).
They reported that the annual 03C isotherm lies at
4900 m, or ca 1400 m above the late Pleistocene ELA,
which therefore is a possible minimum value for ELA
depression.
3.3. Andes of South America
3.3.1. Sierra Nevada de MeH rida, Venezuela
Schubert (1974, 1984) and Schubert and Clapperton
(1990) reported evidence of multiple glacier advances in
the Sierra Nevada de MeH rida (altitudes to 5000 m) between 8315# and 9300#N latitude in the central Venezuelan
Andes. He assigned the mapped drifts to several stades of
the MeH rida Glaciation, of presumed Late Pleistocene age.
The oldest recognized drift (Early Stade), which lacks
dating control and has no clear morainal morphology,
may date to marine isotope stage 4 or 6 (Clapperton,
1993, Table 14.1a). Well-developed moraines of the last
glaciation (Late Stade) form a nested and overlapping
succession between 3000 and 3500 m altitude. A minimum age of 12,650$130 !"C yr BP [15,128 (14,874)
14,621 cal yr BP] was obtained for peat upvalley from the
youngest of these moraines. Peat beneath and above
thick outwash just beyond the outer moraine limit is
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
dated 19,080$820 and 16,500$290 !"C yr BP [22,700
and 19,736 (19,423) 19,153 cal yr BP], respectively, pointing to an isotope stage 2 age for the Late Stade. A date
for basal peat on the #oor of a cirque demonstrates
deglaciation near the range crest by 11,470 !"C yr BP
[13,463 (13,384) 13,317 cal yr BP] (Schubert and Clapperton, 1990).
Schubert and Valastro (1974) mapped the glacial geology in the PaH ramo de la Culata (northern Venezuelan
Andes), where end moraines of the last glaciation are
traceable as low as ca 3150 m and cirque headwalls
upvalley average 4450 m. Based on the median altitude
method, the ELA for Pico BolmH var (&5000 m) would
"
have been at ca 3800 m (Fig. 7a). This is about 900 m
below the modern snowline [ca 4700 m according to
Schubert (1974), but more recently above 4700 m
(Schubert, 1998)]. However, this !ELA value should be
considered only approximate, for the ELA gradient has
not been considered and the modern snowline value is
only an estimate.
3.3.2. Cordillera Central, Colombia
Glaciers and perennial snow"elds on the high Andean
volcanoes Nevado del RumH z (5200 m), Nevado de Santa
Isabel (4950 m), and Nevado del Tolima (5200 m) cover
ca 36 km# between latitudes 4335# and 5310#N (Herd,
1974, 1982). According to Hoyos-Patin$ o (1998), the
snowline on Nevado del RuH iz lies at 4900 m on the
western #ank and 4800 m on the eastern #ank (Fig. 7b).
On Nevado de Santa Isabel, the snowline lies at 4800 m
on the western #ank and 4700 m on the eastern #ank
(Hoyos-Patin$ o, 1998) (Fig. 7c). These values are
60}170 m higher than those measured by Herd (1974,
1982) in 1972}73. Herd's values were based on average
transient summer snowline, which was estimated to lie ca
100 m below the actual snowline.
Herd (1974) mapped the approximate extent of late
Pleistocene glaciers on the massif and obtained a date of
13,760$150 !"C yr BP [16,705(16,498)16,286 cal yr BP]
for peat directly beneath a tephra that overlies the outermost drift of the last glaciation. His mapped glacier limit
approximates the closely nested limits of the early and
late Murillo drifts of Thouret et al. (1996). The late
Murillo advance predates basal peat overlying the moraines having an age of 16,220$80 yr BP [19,247
(19,100) 18,970 cal yr BP], while the early Murillo moraines likely date between 28,000 and 21,000 !"C yr BP
[28,000 and 23,100 cal yr BP] based mainly on regional
tephrochronology. In the Eastern Cordillera of Colombia, the greatest advance of the last glaciation is believed
to have culminated between ca 23,500 and 19,500
!"C yr BP (ca 28,000 and 23,100 cal yr BP) (Helmens
et al., 1996).
Herd (1974) determined the late Pleistocene ELA using
the median-altitude method, assuming that ice divides of
the full-glacial ice caps on the volcanoes represented the
1079
headward limits of individual ice tongues. The reconstructed ELA on Nevado del RumH z, along a section oblique to the modern ELA gradient, rose from ca 3550 m
on the eastern slope to ca 4000 m on the western slope.
On Nevado de Santa Isabel, the Late Pleistocene ELA
along the same transect as the present ELA gradient rose
from ca 3500 to 4000 m (Herd, 1974). In both cases, Herd
calculated the !ELA as ca 950$50 m. However, if the
modern snowline values of Hoyos-Patin$ o (1998) are used,
the mean !ELA increases to 1075 for Nevado del RumH z
and 1000 m for Nevado de Santa Isabel (Figs. 7b, c).
3.3.3. Ecuadorian Andes
The Ecuadorian Andes comprise two parallel ranges,
the Cordillera Occidental (western range; 0322#N}1329#S)
and the Cordillera Oriental (eastern range; 0301#N}
2320#S), separated by intermontane basins and forming
dissected plateaulike surfaces at 3500}4000 m altitude
(Clapperton, 1987) (Fig. 7d). The ranges are surmounted
by 14 glacierized stratovolcanoes more than 4600 m high
(Jordan and Hastenrath, 1998). Clapperton (1987, 1993)
summarized the glacial geology of the mountains and
noted that deposits of `full-glaciala age typically include
3}4 lateral and/or terminal moraines that date broadly to
within the interval 34,000}12,000 !"C yr BP [38,700 and
14,111 (13,992) 13,880 cal yr BP). In discussing his approach for determining ELA depression, he noted that
some earlier studies used cirque-#oor altitudes to calculate !ELAs. He cautioned that during the glacial maximum, most cirques were buried under continuous
ice"elds, implying that cirques likely originated during
earlier intervals of reduced glacier cover. Therefore, such
cirques cannot be used for constructing former snowlines. Exceptions were mountain ridges in southern Ecuador that are too narrow to have supported ice"elds or ice
caps.
Clapperton (1987) estimated the ELA based on "eld
!
observations and aerial photographs, deriving values of
4800}4000 m for Chimborazo (6310 m) in the Cordillera
Occidental and 4600}4830 m for Antisana (5790 m) in the
Cordillera Oriental (Fig. 7d). The values for ELA on
!
Chimborazo volcano compared closely with those based
on the upper limits of lateral moraines. He used the
median-altitude method to calculate ELA depression
along transects across the two cordilleras, noting, however, that the results were `crude approximationsa because the estimated modern ELAs were imprecise and
the reconstructions were based on topographic maps
with a large (40 m) contour interval. In both cordilleras,
eastward-sloping ELA and ELA gradients are present,
!
"
and !ELA is greatest on the eastern #ank of each range
(Fig. 7d). Mean !ELA values increase from west to east
across the Andes, from 810}920 m in the Cordillera Occidental to 970}1120 in the Cordillera Oriental (based on
values in Clapperton's Fig. 7; values in Clapperton's
Table 2 are somewhat greater).
1080
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
Fig. 7. Sites in tropical South America where ELA and !ELA estimates have been obtained. (a) Pico BolmH var, Venezuela; (b, c) Nevado del RumH z and
"
Nevado de Santa Isabel, Colombia; (d) Chimborazo and Antisana, Ecuadorian Andes; (e) Cordillera Blanca and Cordillera Oriental, Peruvian Andes.
See text for details.
3.3.4. Peruvian Andes
Studies of snowline depression in Peru have centered
mainly in the Cordillera Blanca (7340#S) and Cordillera
Oriental (8308#}9358#S) (Fig. 7e). Rodbell (1992) used the
THAR method and derived measurements for ca 20
glaciers in each mountain range using 1 : 100,000-scale
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
topographic maps with 50 m contours, supplemented by
measuring on 1 : 25,000-scale aerial photographs the upper limits of preserved lateral moraines. Results obtained
using THAR values of 0.2 and 0.4 typically di!ered by ca
150}350 m. The accuracy of resulting ELAs was estimated to be $50 m. The age of full-glacial moraines is
unknown, but radiocarbon dates from moraine-dammed
lakes and bogs provide minimum ages for deglaciation
in the Cordillera Oriental of 12,100$190 !"C yr BP
[14,382 (14,114) 13,870 cal yr BP] and in the Cordillera
Blanca of 13,280$190 !"C yr BP [16,140 (15,859)
15,559 cal yr BP].
3.3.4.1. Cordillera Oriental. Rodbell (1992) obtained
ELA values for former glaciers lying east and west of the
"
main divide, as well as west of a local divide lying ca
20 km west of the main divide (Fig. 7e). No glaciers are
present in the study area, but ELA is based on a re!
gional glaciation threshold estimate of 4620 m (Seltzer,
1987). Paleoglaciers on the western side of the local
divide had an average ELA of 3850-3900 m, represent"
ing a !ELA of 750}950 m. Corresponding ELA values
"
west and east of the main divide are 3540}3640 and
3150}3300 m, respectively. The !ELAs are 900}1150 and
1100}1350 m, respectively, with estimated mean values of
ca 850}1000 and 1200 m for the two respective sides of
the range (Rodbell, 1992). These results demonstrate
a W}E snowline gradient, with snowline depression increasing toward the Amazon Basin, which is, and apparently was, the primary source of precipitation.
3.3.4.2. Cordillera Blanca Rodbell's (1992) estimates of
ELA in the Cordillera Blanca range from 4985$120 m
!
west of the divide to 5030$110 m east of the divide,
indicating no discernable gradient (Fig. 7e). The calculated ELA
is 4400$100 to 4250$110 m
"
(THAR"0.2 and 0.4, respectively) west of the divide and
4200$170 to 4350$150 m east of the divide, giving an
average of ca 4300 m for the range as a whole. The
calculated !ELA is 440}900 and 530}970 m, respectively.
Rodbell (1992) suggested that, based on comparison with
ELA and glaciation threshold values, the average
!
!ELA for the range is ca 700 m.
3.3.4.3. Regional Andean reconstructions.
Nogami
(1976) compared the modern snowline along the entire
Andean cordillera (103N}553S), based on the altitude of
existing glaciers shown on aerial photographs (pre-1976),
with a Pleistocene snowline reconstructed using the cirque-#oor method. The regional pattern disclosed that
past and recent snowline surfaces rise westward between
the northern equatorial Andes and northern Chile
(ca 303S), at which latitude a shift occurs to northeastward-rising snowlines. The di!erence between Nogami's
two surfaces (!ELA) is less than 1000 m.
1081
Fox and Bloom (1994) made a regional study of Late
Quaternary snowlines in the Peruvian Andes (5}173S).
They based their estimate of modern snowline altitude on
the lower limit of snow depicted on 1 : 100,000-scale aerial photographs (1955}63). The data were transferred to
topographic maps (50-m contour interval), and an error
of $100 m was assumed. They showed that the snowline
rises from ca 4700 m in the northern and eastern Andes
to more than 5300 m in the south and west, and has
a westward-rising gradient, especially in the north. The
overall pattern is similar to that derived by Nogami
(1976). Full-glacial snowline was reconstructed based on
the cirque-#oor method, assuming that glaciers occupied
the lowest cirques contemporaneously. The time of this
occupation, however, has not been dated. The reconstruction indicates that snowline depression reached
a maximum of 1400$200 m on the eastern side of the
Peruvian Andes but decreased westward to a minimum
of 600 m in the western ranges and on most of the
Peruvian Altiplano (Fig. 8). An average value was not
given, but based on their Fig. 7 it appears to be in the
range of 800}1000 m in the north, decreasing to
600}800 m in the south.
Klein et al. (1999) subsequently expanded on the work
of Fox and Bloom (1994) to include not only the Andes of
southern Peru, but also Bolivia and northern Chile
(Fig. 8). They adopted Fox and Bloom's estimate of the
modern snowline altitude in Peru (see above), and used
LANDSAT Thematic Mapper imagery to map the lower
limit of snow cover in the southern part of their study
region. Maps with scales of 1 : 50,000 (20 m contours) and
1 : 250,000 (250 m contours) were used. They noted the
lack of adequate dating control for the last glaciation, but
assumed that moraines mapped as (local) LGM were
constructed contemporaneously. In Peru, cirque-#oor
altitudes were used to determine LGM snowlines; however, the age of the last cirque glaciation is unknown, and
so their assumed age equivalence is unproven. Furthermore, many cirques lay at the heads of valley glaciers. In
the southern region, an adjusted THAR of 0.45 was used
to calculate speci"c snowlines, as well as a regional snowline pattern for the last glaciation. In the area where data
from the two methods overlap, the di!erence in estimated
ELA between the methods is ca 175 m. The con"guration
of the reconstructed glacial snowline is similar to that of
Fox and Bloom's (1994) modern snowline. Their calculated snowline depression over the tropical Andes
averaged 920$250 m. However, consistent with the conclusions of Fox and Bloom (1994), the calculated !ELA
is ca 1200 m in the eastern cordillera of Peru and Bolivia,
and 500}800 m to the west [see also Seltzer (1992, 1994),
who calculated an ELA depression of only 300$100 m
on the western slope of the Cordillera Real]. In the arid
ranges along the border of Bolivia and Chile, snowline
depression reached 1000}1200 m (Klein et al., 1999,
Fig. 7).
1082
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
the LGM, an ice cap covered 70 km# of the upper slopes
above ca 3200 m altitude (Porter, 1979) (Fig. 9b). Two
surface-exposure ages have been obtained for boulders at
the surface of the youngest (Makanaka) drift and one for
glacially abraded rock near the summit (20,300$2300,
18,900$800 yr, and 14,700$500 $'Cl yr, respectively;
Dorn et al., 1991). These imply that the LGM occurred
during marine oxygen-isotope stage 2 and that the summit was deglaciated by ca 15,000 yr ago.
Based on the AAR method, the full-glacial ELA, corrected for ca 35 m of postglacial isostatic subsidence
(2.5 m/10$ yr; Porter, 1979), averaged ca 3780 m and had
an eastward-sloping gradient. The minimum !ELA was
ca 425 m. The July freezing isotherm now lies close to
4715 m, about 500 m above the summit, and likely approximates the ELA . Assuming a comparable relation!
ship between July temperature and ELA during the
"
LGM, and using an AAR of 0.60$0.05, !ELA was
935$190 m.
Fig. 8. Map of the Peruvian, Bolivian, and northern Chilean Andes
showing pattern of regional ELAg surface (after Klein et al., 1999). Bold
lines perpendicular to trend of isolines show arbitrary transects depicted in Fig. 10 and included in Table 1.
3.4. Pacixc Islands
3.4.1. Taiwan Shan, Taiwan
Along the Taiwan Shan, which forms the high crest of
Taiwan, 62 peaks exceed 3000 m altitude; the highest, Yu
Shan, reaches 3997 m. Kano (1934}35) identi"ed 35 cirques in the northern sector of these mountains, lying
mainly on the eastern side of the range. Moraines on
cirque #oors and beyond cirque thresholds de"ne the
limits of former cirque and valley glaciers, as well as
a small ice cap. Although no dates were available, Kano
assigned the landforms to the last glaciation and noted
that cirque altitudes range from ca 3500 to 3730 m. He
suggested that glaciers extended down to 3300 m on the
northern and eastern slopes of the highest peak, but were
ca 300 m higher on the southern and western sides of the
range.
Ono (1988, Fig. 1; Y. Ono, pers. comm. 2000) estimated
the full-glacial snowline on Taiwan to lie at ca 3400 m,
midway between an ELA of 3350 m on Xue Shan and
"
3450 m on Yu Shan (to the south), derived using the
cirque-#oor and glaciation threshold methods (Fig. 9a).
Based on this "gure, and an inferred average crest altitude of 3800 m, the !ELA was '400 m.
3.4.2. Mauna Kea, Hawaii
The summit of Mauna Kea (4206 m; 19350#N) on the
island of Hawaii lacks perennial glacier ice. However, at
3.4.3. Mt. Kinabalu, Borneo
Glacial-erosional features below the summit of Mt.
Kinabalu (4101 m; 6305#N) on the island of Borneo were
reported by Koopmans and Stau!er (1967) and Stau!er
(1968). They estimated a glacial-age snowline of
3735$75 m (12,000}12,500 ft) for the mountain based
on the median-altitude method (Fig. 9c). The downvalley
extent of ice was inferred from possible moraines at ca
2835 and 3230 identi"ed on aerial photographs. The
glacial deposits have not been dated. If the landforms are
correctly identi"ed as moraines and the higher one dates
to the LGM, as inferred here, then based on the medianaltitude method, the average ELA lay at ca 3665 m and
"
!ELA was at least 435 m. Koopmans and Stau!er estimated that the ELA lies at ca 4570$150 m, which
!
would imply a !ELA of 905$150 m during the last
glaciation.
3.4.4. Papua New Guinea
New Guinea is the only equatorial Paci"c island
(5}93N) with numerous highlands (ca 3800}4500 m) that
generated a Late Pleistocene glacier cover. Most glacialgeologic studies have focused on the eastern half of the
island (Papua New Guinea), which is more accessible
than Irian Jaya to the west. Although no glaciers exist in
the eastern highlands, LoK %er (1972) inferred that the
snowline lies at ca 4600 m, the reported altitude of the
snowline in the glaciated areas of Irian Jaya (Verstappen,
1964; Allison, 1976). The LGM snowline was determined
by LoK %er (1972) using the arithmetic mean of the altitudes of terminal moraines and the mean altitude of the
catchment area, as well as the altitudes of the lowest
cirque #oors (Fig. 9d), and is similar to estimates based
on the glaciation threshold (LoK %er, 1971). Bowler et al.
(1976) estimated the age of the LGM to be ca
18,000}16,000 !"C yr BP [21,648 (21,484) 21,313}18,972
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
(18,876) 18,784 cal yr BP] based on pollen-derived estimates and limiting !"C ages from two highland sites.
Hope and Peterson (1976) reported that maximum depression of vegetation zones occurred 18,500}16,000
!"C yr [22,000}18,972 (18,876) 18,784 cal yr] ago, and
that in several areas substantial ice retreat had occurred
by 14,500}14,000 yr [17,453 (17,369) 17,287}16,879
(16,792) 16,704 cal yr] ago.
Mt. Giluwe (4368 m), a dome-shaped stratovolcano at
63S latitude, was mantled by a large ice cap (188 km#)
during the last glaciation. Blake and LoK %er (1971) described end moraines that form concentric belts around
the mountain as low as 2750}3000 m, and alluvium, interpreted as outwash, that overlies peat having an age of
23,600$1100 !"C yr (ca 27,400 cal yr) old. LoK %er (1972)
calculated an ELA of 3500}3550 m (Fig. 9d), which
"
represents a minimum !ELA of ca 820}870 m.
For other areas of less-extensive glaciation lying between ca 53 and 93S latitude, LoK %er (1972) estimated the
ELA to lie between 3500 and 3700 m (Fig. 8d). However,
"
recent studies in the Sarawaged Range, using modern
topographic maps, suggest an ELA as low as 3400 m
!
(M. Prentice, pers. comm., 2000). If a snowline gradient
existed across New Guinea, it was very gentle and cannot
be de"ned on the basis of available data. Adopting Verstappen's (1964) value of 4600 m for the modern snowline,
LoK %er (1972) concluded that snowline depression in
these areas during the last glaciation was ca 900}1100 m.
4. Discussion
As summarized above, data bearing on snowline depression in tropical latitudes are restricted to eastern
Africa, Central and South America, and several Paci"c
islands. Di!erent methods have been used in deriving
ELA , ELA , and !ELA, but the results are not strictly
!
"
comparable. The AAR method, often regarded as the
most reliable and consistent, has been used infrequently
in tropical snowline studies. Altitude ratios (median altitude, THAR, Alt}Ht}Acc) have been employed in most
cases. In the few instances where lateral moraines were
used to de"ne former ELAs, the results were similar to
those obtained using altitude ratios. Cirque-#oor altitudes likely are the least-reliable method, especially without associated "eld studies and adequate dating control.
Although cirques may record some average regional level
of glacial conditions (e.g., Porter, 1989), they do not
necessarily record a synchronous glacial event, such as
the LGM. Di!ering methodologies, therefore, introduce
potential variance to !ELA estimates, the magnitude of
which is di$cult to assess. In some temperate-latitude
studies (e.g., Meierding, 1982), methodological di!erences
amounted to 100 m or more in calculated ELA ; compa"
rable di!erences of 100}200 m probably should be expected in tropical snowline studies.
1083
4.1. Estimates of modern ELAs
Estimates of the modern snowline altitude probably
constitute the least-reliable component of the !ELA calculations. In many cases, summit altitude is used as
a minimum altitude for the modern snowline, recognizing that the ELA must lie some unknown distance
!
above a nonglacierized summit. In other cases, the altitude of the modern snowline is estimated, either based on
limited "eld work (usually by noting the level of the
transient snowline during some part of the ablation season), by assuming a relationship between the snowline
and the summer (July Northern Hemisphere) or annual
mean freezing isotherm based on radiosonde data, or by
using the mapped limits of snow and glacier cover depicted on published topographic maps. In the latter approach, generally it is assumed that snow/ice conditions
shown on an array of maps that cover an area or region
are essentially contemporaneous and represent a steady
state, and that the contour interval is suitable for relatively high-resolution interpolation between contours (e.g.,
)30 m).
The lack of a consistent and rigorous method of determining the modern snowline in the tropics could introduce an error of up to several hundred meters in some of
the reported !ELA results. Furthermore, in a time when
global climate is warming, snowline values measured
several decades apart may di!er by tens of meters. Potential errors may also occur when the ELA is inferred to
!
lie at the level of the summer freezing isotherm, if this
assumed relationship is not always valid.
4.2. Reconstructed Pleistocene ELAs
Data from 26 sites in the tropics permit reconstructions of full-glacial snowline (ELA ) (Table 1; Fig. 10).
"
However, for only 11 of these has the ELA gradient been
"
determined; nevertheless, in only two cases (IztaccmH huatl
and Cordillera de MeH rida) may the lack of a calculated
snowline gradient have introduced a signi"cant error
into the !ELA calculation.
For most localities or regions, the age of the reconstructed ELA is unknown, but it is commonly assumed
"
to equate with the `globala LGM (i.e., ca 21,000}15,000
!"C yr BP; ca 24,400}17,453 (17,369) 17,287 cal yr BP).
Available radiocarbon dates (Table 1) generally are
inadequate to verify whether the reconstruction represents full- or late-glacial (or even pre-LGM) conditions.
In South America, regional reconstructions for the
Andes of Peru, Bolivia, and Chile (Nogami, 1976; Fox
and Bloom, 1994; Klein et al., 1999), mainly using the
cirque-#oor method, add additional data for the tropics
that supplement information from speci"c areas. Despite
the acknowledged uncertainties and assumptions involved in these studies, including a potential error of
1084
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
Fig. 9. Sites in tropical Paci"c islands where ELA and !ELA estimates have been obtained. (a) Taiwan Shan, Taiwan; (b) Mauna Kea, Hawaii; (c) Mt.
"
Kinabalu, Borneo; (d) Papua New Guinea. See text for details.
$100}200 m at any site, they serve to emphasize the
pattern, signi"cance, and overall consistency of regional
ELA gradients along and across mountain systems.
4.3. Snowline depression
For 12 of the tropical sites the modern ELA has been
determined or estimated with su$cient con"dence that
!ELA at the LGM can be calculated. For the others, the
reported values are approximate, or only minimum estimates (Fig. 10; Table 1).
To assess snowline depression, the data are divided
into two groups. The "rst group (mainly north of 103S
latitude) includes speci"c estimates at tropical sites in
Africa, the Americas, and Paci"c islands (Fig. 10). The
two summits of Kilimanjaro are considered a single locality, and average values are used for the Cordillera
Blanca and the Cordillera Oriental reported by Rodbell
(1992). The second group (mainly south of 103S latitude)
includes regional reconstructions spanning segments of
the tropical (central) Andes between 73 and 223S latitude.
The mean !ELA for the "rst set of data (n"12) is
900$135 m. The Ruwenzori forms an outlier from the
otherwise reasonably tight data set; if it is excluded, the
resulting mean is not statistically di!erent at 1"
(925$115 m).
The second group of data (n"8) is represented by one
composite data set (Fox and Bloom, 1994; not plotted in
Fig. 10) and 7 transects parallel to ELA gradients at 23
latitude intervals, with values derived from Klein et al.
(1999, Fig. 7) (Fig. 8; Table 1). These data illustrate a substantial regional range of !ELA values, especially perpendicular to the trend of the Andes. Through this sector of
the cordillera, as far south as 183S latitude, the ELA is
"
lowest, and !ELA values are greatest, on the Amazonian
slope. Although the !ELA varies regionally, Klein et al.
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
1085
Fig. 10. (a) Modern and full-glacial ELAs for areas in the tropics (23.53N}23.53S latitude) that supported Pleistocene mountain glaciers. Where
summits now lie below the snowline, minimum ELAs are shown as the summit altitude. In some cases ELA and (or) ELA are shown with a range of
!
"
values, primarily resulting from ELA gradients across mountains or mountain ranges. See text and Table 1 for details. (b) Full-glacial snowline
depression (!ELA) for tropical mountains and mountain ranges that supported Pleistocene glaciers. Minimum values represent summits that lie below
the modern snowline. In areas where ELA gradients exist, the number shown is the median of a range of values. Areas north of about 103S latitude have
a mean !ELA of 900$135 m, whereas a regional study of Andes south of this latitude produced an estimated mean of 920$250 m (Klein et al., 1999).
(1999) calculated an average value of 920$250 m, the
mean being close to that of the "rst group of data discussed above (Fig. 10b). In the following discussion, the
value for the "rst data set will be used to represent global
tropical snowline depression during the LGM.
A snowline depression of 900$135 m for tropical glaciers at the LGM is similar to estimates obtained for
many temperate-latitude late Pleistocene mountain glaciers in both the Northern and Southern hemispheres
(e.g., Porter, 1975; Porter et al., 1983; Furrer, 1991).
Departures from the general average may be related to
any of a number of factors, including local variations in
temperature depression, distance from precipitation
sources, or nonuniform changes in accumulation (i.e.,
precipitation) and radiation (as in#uenced by cloudiness,
surface albedo, and topographic shading) in di!erent
areas. The close similarity of derived !ELA values for
tropical and temperate latitudes argues for a fundamental temperature control of !ELA and implies a reasonably consistent decline of air temperature during the
glacial maximum in extra-polar alpine regions near maritime sources of precipitation.
As discussed earlier, because sea level was ca 120 m
lower than today at the LGM (Fig. 3), a !ELA of
1086
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
900$135 m is equivalent to a !ELA of 780$135 m as
%
a result of the changing ocean reference level.
5. Paleotemperature inferences from snowline data
Paleotemperature values based on snowline depression are frequently cited, but often uncritically. As Seltzer
(1994, p. 159) has emphasized, `climatic interpretations of
ELA depression will always lack unique solutions because of the complexity of the problema (see also Porter,
1977). The most common approach has been to calculate
lowering of temperature (usually annual, summer, or
July, but not always clearly stated) based on an assumed
LGM atmospheric temperature lapse rate. Inferred lapse
rates vary widely. For example, LoK %er (1970) applied
a lapse-rate range of !5 to !63C/km in the highlands
of New Guinea; Porter (1979) used a lapse rate of
!5.33C/km for Mauna Kea, Hawaii; Clapperton (1987)
applied a lapse rate of !6.53C/km in Ecuador, a value
also used by Rodbell (1992) and Seltzer (1987) for the
Peruvian Andes; Wright (1983) and Osmaston (1989a)
used a lapse rate of !73C/km in Peru and East Africa,
respectively; and Fox and Bloom (1994) used a nonlinear
lapse rate for the tropical Andes (!6.53C/km at
!3.5 km altitude to nearly !103C/km at 6 km). Hostetler and Mix (1999) adopted a `nominal tropical lapse
ratea of !5.5 3C/km. This range in lapse-rate values
(!5.3 to !103C/km) by itself translates into a 4.53C
range of values for temperature lowering, assuming
a snowline depression of 1000 m.
Assuming a mean tropical lapse rate of !6$13C/km,
and no change in precipitation, an average snowline
depression of 900$135 m translates into a full-glacial
mean temperature depression of 5.4$0.83C. Using the
full range (550}1400 m) of reported tropical !ELA
(Table 1), the lapse-rate approach produces temperature
depressions ranging from 3.3 to 8.43C. Adjusted for sealevel lowering of 120 m, average temperature depression
is 4.7$0.83C.
In these simple, straightforward calculations, changes
in the accumulation component of glacier mass balance
have been ignored. Intuitively, it would seem to be an
important factor in some alpine regions. However, Seltzer (1994) assessed its importance and concluded that
relatively large changes in precipitation would be required to a!ect ELA substantially. In tropical areas with
high precipitation, the limiting control on glacier extent
likely is the altitude of the 03C isotherm (Hostetler and
Clark, 2000). Of equal or greater importance may be the
low seasonality of tropical climates, which leads to a relatively constant height of the freezing isotherm (Klein
et al., 1999). At these latitudes, ELAs that lie above the
level of the 03C isotherm are sensitive to accumulation
changes, for above this level all precipitation falls as
snow. At lower altitudes, as temperature rises above 53C,
precipitation falls as rain. In the Andes, for example, the
precipitation gradient is not uniform: rainfall reaches
a maximum at ca 1000 m altitude, above which it decreases. A drop in freezing level, therefore, may convert
a larger percentage of the precipitation to snow and
signi"cantly increase accumulation, without a change
in net precipitation. In this way, a relatively uniform drop
in air temperature might produce regionally variable
glacier mass balances leading to nonuniform ELA
depression.
It is apparent that estimating paleotemperatures using
snowline data involves some substantial uncertainties.
Not only are there pitfalls in the use of di!erent methodologies, as well as signi"cant potential ranges of error,
but the multiple factors that control glacier mass balance
and ELA do not permit an unequivocal and unique
paleotemperature solution. Nevertheless, the regional
and global averages for tropical data suggest that the
simplistic lapse-rate approach may at least provide
a "rst-order approximation of regional and global tropical temperature reduction at the LGM.
6. Other tropical LGM paleoclimate data and model
simulations
The mean paleotemperature values based on snowline
depression are in general accord with other paleotemperature estimates from tropical land areas that suggest
full-glacial temperatures were substantially lower than
tropical warm-season SSTs (Fig. 11).
The CLIMAP Project Members (1976, Fig. 3; 1981)
derived a mean LGM tropical SST cooling of ca 1}33C,
relative to modern (Fig. 11). More recent studies of the
tropical oceans report LGM SSTs that were 1.7$0.7 to
2.8$0.73C lower than present based on alkenone data
(Lyle et al., 1992; Sikes and Keigwin, 1994; Bard et al.,
1997) and Mg/Ca data (Lea et al., 2000), broadly consistent with mean CLIMAP estimates (Fig. 11). In contrast,
oxygen-isotope and Sr/Ca data from corals at Barbados
imply LGM SSTs 5 to 63C colder than now (Guilderson
et al., 1994). Crowley (2000), however, has questioned the
interpretation of the Sr/Ca data, noting that an SST
depression of this amount would leave ice-age corals at
or below their limit of habitability.
A variety of terrestrial climate proxy data suggests that
LGM temperature lowering was greater over land than
over the ocean. Representative estimates (Fig. 11 and
Table 2), which are based on pollen data, noble-gas
values in groundwater, and oxygen-isotope records in
glacier ice, range from !5 to !123C (for additional
data, see Farerra et al., 1999). Plotted with these temperature estimates in Fig. 11 are values based on !ELA and
!ELA . These average snowline-based estimates, at 1
%
standard deviation, fall close to many of the other estimates of terrestrial temperature lowering, and therefore
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
1087
Fig. 11. Estimates of temperature depression during the last glaciation (!Tg) at representative tropical sites, based on various climate proxies,
compared with CLIMAP tropical sea-surface temperature (SST) di!erence between modern August and August 18,000 !"C yr BP (CLIMAP Project
Members, 1976, Fig. 3). Temperature depression based on average tropical !ELA is shown by bold line and 1" range by light shading; value adjusted
for sea-level lowering (!ELA ) is shown by dashed bold line, and 1 " range by dark shading. Error estimates (1"), when reported, are shown by vertical
%
dashed lines. 1 * eastern Atlantic Ocean alkenone data (Sikes and Keigwin, 1994); 2 * Indian Ocean alkenone data (Bard et al., 1997); 3 * central
Paci"c Ocean alkenone data (Lyle et al., 1992); 4 * western and eastern Paci"c Ocean Mg/Ca data (Lea et al., 2000); 5 * Hawaiian foram !!(O data
(Lee and Slowey, 1999); 6 * Barbados !!(O and Sr/Ca data (Guilderson et al., 1994); 7 * noble gases in Oman groundwaters (Weyhenmeyer et al.,
2000); 8 * pollen in Guatemala lakes (Leyden et al., 1993); 9 * noble gases in Nigerian groundwaters (Edmunds et al., 1999); 10, 11, and 12 * pollen
data from Panama, Brazil, and Ecuador, respectively (Colinvaux et al., 1996); 13 * noble gases in Brazil groundwaters (Stute et al., 1995); 14 * !!(O of
Huascaran ice cap, Peru (Thompson et al., 1995); 15 * pollen data from Brazil (Colinvaux et al., 1996); 16 * noble gases in Nambia groundwaters
(Stute and Talma, 1998).
Table 2
Representative tropical SST and terrestrial paleoclimatic data for the LGM
Data No.
Latitude
Location
!SST (3C)
Data
Reference
SST data
1
2
3
4
5
6
03
203N}203 S
0357#N
0319'}2348#N
21.36#N
13315#N
E Atlantic Ocean
Indian Ocean
Central Paci"c Ocean
W & E Paci"c Ocean
Hawaiian Islands
Barbados
!1.8
!1.7$0.7
!1
!2.8$0.7
!2
!5
Alkenone
Alkenone
Alkenone
Mg/Ca
Foram !!(O
Sr/Ca
Sikes and Keigwin (1994)
Bard et al. (1997)
Lyle et al. (1992)
Lea et al. (2000)
Lee and Slowey (1999)
Guilderson et al. (1994)
Terrestrial Data
7
8
9
10
11
12
13
14
15
16
23330#N
16355#N
11330'}13330#N
93N
03
63S
73S
93S
21}223 S
24330#S
Oman
Guatemala
Nigeria
Panama
Brazil
Ecuador
Brazil
Peru
Brazil
Namibia
!6.5$0.6
!6.5 to !8
!6
*!5
!6
!6
!5.4$0.6
!8 to !12
!6 to !9
!5.3$0.5
Noble gas
Pollen
Noble gas
Pollen
Pollen
Pollen
Noble gas
!!(O of ice
Pollen
Noble gas
Weyhenmeyer et al. (2000)
Leyden et al. (1993)
Edmunds et al. (1999)
Colinvaux et al. (1996)
Colinvaux et al. (1996)
Colinvaux et al. (1996)
Stute et al. (1995)
Thompson et al. (1995)
Colinvaux et al. (1996)
Stute and Talma (1998)
support the conclusion that LGM surface land temperatures typically were depressed more than SSTs of adjacent oceans. Similar values have been reported in
a modeling study of LGM climate that used a global
coupled ocean}atmosphere model of intermediate complexity (Ganopolski et al., 1998). The simulation showed
that tropical land areas cooled an average of 4.63C (comparable to ELA , Fig. 11), whereas SSTs between 203N
%
1088
S.C. Porter / Quaternary Science Reviews 20 (2001) 1067}1091
and 203S cooled by 3.33C in the Atlantic, 2.43C in the
Paci"c, and 1.33C in the Indian Ocean.
Farerra et al. (1999) used a variety of climate-proxy
data to calculate average cold-season cooling at the
LGM that ranged from !2.5 to !3.03 at present sea
level to ca !63 at 3000 m altitude, suggesting nonlinear
lapse rates. Such a relationship is apparent in a recent rise
in the tropical freezing level, which is closely linked to an
increase in SST (Diaz and Graham, 1996). Whereas the
observed tropical SST change was ca 0.2}0.33C, temperature change at the level of the freezing isotherm, based
on a standard lapse rate of !63/km, was ca 0.63C, or
two to three times as great (Crowley, 2000). Modi"cation
of tropical lapse rates at the LGM could well be related
to a change in the mean altitude of the tropical inversion,
in turn associated with a shift in the position and strength
of subtropical high-pressure cells (e.g., Hostetler and
Clark, 2000).
Betts and Ridgway (1992) evaluated several factors
that may be related to the discrepancy between tropical
snowline depression and SSTs. They computed that a decrease in average tropical SST by 23 at the LGM and an
increase in tropical sea-surface pressure by 14 mbar (the
result of a 120 m fall in sea level) would account for an
800 m depression of the freezing isotherm. Signi"cantly,
this result is consistent with recent (post-CLIMAP) estimates of SST lowering cited above (1.7}2.83C) and the
mean !ELA (780$135 m) derived in the present study.
%
The mean !ELA value is also close to !ELAs (ca
%
800}870 m) reported by Hostetler and Clark (in press)
based on mass-balance modeling of several tropical glaciers in New Guinea, Hawaii, Africa, and the Andes.
This review has focussed on the derivation and assessment of tropical paleosnowlines. The limited quantity
and limitations of the existing information point to the
need for additional studies to enlarge and improve the
data set. However, even with adequate data, translating
snowline depression into estimates of land-surface temperature depression is likely to persist as a challenging
problem. Among the important questions yet to be answered are: (1) what was the degree to which the mass
balance of tropical glaciers at the LGM was in#uenced
by a change in precipitation? and (2) were LGM lapse
rates di!erent than today, and were they linear or nonlinear? At present, suitable evidence to answer these questions remains elusive.
Acknowledgements
The initial draft of this paper was written while I enjoyed hospitality of Nick Shackleton in the Godwin Laboratory, Cambridge University, and Lloyd Keigwin in the
McLean Laboratory, Woods Hole Oceanographic Institution. I thank Matsuo Tsukada for translation of a Japanese article used in this review, and Geo! Seltzer for
information about his snowline studies in South
America Helpful comments by reviewers Alan Gillespie,
Michael Prentice, and Geo! Seltzer are greatly appreciated.
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