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1 Earth-Science Reviews Timing of glaciation during the last glacial cycle: evaluating the concept of a global ‘Last Glacial Maximum’ (LGM) Philip D. Hughes a,* , Philip L. Gibbard b , Jürgen Ehlers c a Geography, School of Environment and Development, The University of Manchester, Oxford Road, Manchester M13 9PL, United Kingdom b Cambridge Quaternary, Department of Geography, University of Cambridge, Downing Place, Cambridge CB2 3EN, United Kingdom c Geologisches Landesamt, Billstraße 84, D-20539 Hamburg, Germany ABSTRACT It has long been known that mountain glaciers and continental ice sheets around the globe reached their respective maximum extent at different times during the last glacial cycle, often well before the global Last Glacial Maximum (LGM; c. 23- 19 ka), which is formally defined by peaks in global sea- 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21
Transcript of Web viewIt is unfortunate that the word 'Glacial' in LGM has led to misleading correlations for ......
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Earth-Science Reviews
Timing of glaciation during the last glacial cycle: evaluating the concept of a global
‘Last Glacial Maximum’ (LGM)
Philip D. Hughes a,*, Philip L. Gibbard b, Jürgen Ehlers c
a Geography, School of Environment and Development, The University of Manchester,
Oxford Road, Manchester M13 9PL, United Kingdom
b Cambridge Quaternary, Department of Geography, University of Cambridge, Downing
Place, Cambridge CB2 3EN, United Kingdom
c Geologisches Landesamt, Billstraße 84, D-20539 Hamburg, Germany
ABSTRACT
It has long been known that mountain glaciers and continental ice sheets around the
globe reached their respective maximum extent at different times during the last glacial
cycle, often well before the global Last Glacial Maximum (LGM; c. 23-19 ka), which is
formally defined by peaks in global sea-level and marine oxygen isotope records.
However, there is increasing evidence from around the world that it was not only
mountain glaciers which were asynchronous with the global LGM but also some regions
of the large continental glaciers. The Barents-Kara Ice Sheet in northern Eurasia
together with a majority of ice masses throughout Asia and Australasia. reached their
maximum early in the last glacial cycle, a few thousand years before the global LGM
period. The East Antarctic Ice Sheet also reached its maximum extent several millennia
before the global LGM. In numerous mountainous regions at high-, mid- and low-
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latitudes across the world, glaciers reached their maximum extent before Marine
Isotope Stage (MIS) 2, in MIS 5, 4 and 3. This is in contrast to most sectors of the
Laurentide Ice Sheet, the Cordilleran Ice Sheet, the SE sector of the Fennoscandinavian
Ice Sheet and the Alpine Ice Sheet in central Europe, which appear to have reached
their maximum close to the global LGM in MIS 2. The diachronous maximum extents
of both mountain glaciers and continental ice sheets during the last glacial cycle, means
that the term and acronym Last Glacial Maximum (LGM) has limited
chronostratigraphical meaning when correlating glacial deposits and landforms.
Contents
1. Introduction
2. Establishing geochronological frameworks for glacial sediments and landforms
2.1 Radiocarbon dating
2.2 Cosmogenic exposure age dating
2.3 Optically stimulated luminescence dating
2.4 U-series dating
2.5 Other techniques relevant to dating Pleistocene glacial sediments and landforms
3. The maximum extents of glaciers during the last glacial cycle – a global view
3.1 North America
3.1.1 Laurentide Ice Sheet
3.1.2 Cordilleran Ice Sheet
3.1.3 Alaska
3.1.4 Rockies
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3.1.6 Hawaii
3.2. Central America
3.3 South America
3.3.1 Colombia
3.3.2 Venezuela
3.3.3 Ecuador, Peru & Bolivia
3.3.4 Chile/Argentina
3.4 Europe
3.4.1 British-Irish Ice Sheet
3.4.2 Fennoscandinavian Ice Sheet
3.4.3 Alps
3.4.5 Pyrenees
3.4.6 Iberia
3.4.7 Corsica
3.4.8 Italian Apennines
3.4.9 Romania and the Balkans
3.4.10 Turkey
3.5 Asia
3.5.1 Barents-Kara Ice Sheet
3.5.2 NE Russia
3.5.3 Pamirs, Tien Shan, Altai
3.5.4 Himalaya and Tibet
3.5.5 Japan and Taiwan
3.6 Africa
3.6.1 Morocco
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3.6.2 Ethiopia
3.6.3 Equatorial Africa
3.6.4 Southern Africa
3.7 Australasia
3.7.1 Papua New Guinea
3.7.2 Australia
3.7.3 New Zealand
3.8 Antarctica
4. Discussion
5. Conclusions
1. Introduction
The concept of a ‘last glacial maximum’ has been current for decades and is used in a formal
sense to refer to the period during the last glacial cycle (Weichselian, Wisconsinan and
equivalents, i.e. the past ~120 ka or since Termination-II) when ice masses reached their last
maximum global extent (Bowen, 2009). CLIMAP Project Members (1976) first coined the
term ‘last glacial maximum’, in simulations of climate at the surface of the Earth “when the
continental glaciers reached their maximum extent in the last ice age, approximately 18,000
B.P.” In a later paper (CLIMAP Project Members, 1981) the first letters were capitalised to
‘Last Glacial Maximum’ and the acronym ‘LGM’ applied, implying the term's quasi-formal
status – a situation which has continued ever since.
The term ‘Last Glacial Maximum’ (abbreviated to LGM) is widely accepted as referring to
the maximum in global ice volume during the last glacial cycle which corresponds to the
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largest minima in the marine isotope record at c. 18 14C ka BP (Martinson et al., 1987) and
the associated global eustatic sea-level low also dated to c. 18 14C ka BP or to 21 cal. ka BP
(Clark and Mix, 2002; Yokoyama et al., 2000). Mix et al. (2001) considered the event should
be centred on the radiocarbon calibrated date of 21 cal. ka BP, and should span the period
23-19 or 24-18 cal. ka BP depending on the dating applied (e.g. MARGO project members,
2009). However, other research on global sea-level minima places the global ice maximum
slightly earlier at between 26-21 ka (Peltier and Fairbanks, 2006). Thus, the actual definition
of the term 'LGM' is open to debate depending on what criteria are used to define it and today
it has no formal stratigraphical status despite attempts to assign it as a chronozone (Mix et al.
2001). Nevertheless, there is indisputable evidence that a major global glaciation did occur
during the broader definition of MIS 2 (e.g. Clark et al., 2009), which spans the period 29 to
12 ka (Martinson et al. (1987). Also see various other definitions reviewed in Sanchez Goñi
and Harrison (2010), their Table 2.
The last glacial cycle, is defined as the period between the penultimate and current
interglacial or generally between Termination II and Termination I (Fischer et al., 1999).
More precisely, it encompasses the last cold stage between 110.8 to 11.7 ka; lower boundary
based on Martinson et al., 1987 and upper boundary based on Lowe et al., 2006). With an
ever increasing geographic spread of glacial studies, many authors adopted the term ‘local’
last glacial maximum to explicitly label glacial maxima that did not coincide with the above
global LGM definition. Gillespie & Molnar (1995) observed that, in many mountain areas
around the world, the ‘local’ last maximum glaciation during the last glacial cycle predated
the global LGM. This observation was reiterated with new evidence from several
mountainous areas of the world in Thackray et al. (2008a). Even so, there is widespread
belief that the majority of the large continental ice sheets reached their maximum close in
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time. In fact, Clark et al. (2009) asserted that “Nearly all ice sheets were at their LGM
positions from 26.5 ka to 19 to 20 ka” – a situation which appears to provide justification and
a firm basis for the concept of a global LGM. Clark et al. (2009) used 5704 14C, 10Be, and 3He
ages that span the interval from 10,000 to 50,000 years ago (10 to 50 ka) to constrain the
timing of the Last Glacial Maximum during MIS 2. Unfortunately, the situation during the
earlier part of the last glacial cycle prior to 50 ka was omitted (MIS 5d to early MIS 3) from
their study. In contrast, Ehlers et al. (2011) recognised that in several areas, some glaciers
were more extensive during MIS 4 in the early part of the last glacial cycle, prior to 50 ka.
With the advent of new radiometric dating techniques which push back the previous
radiocarbon limit of ~50 ka, there is increasing evidence to suggest that glacier advances
were more extensive earlier in the last glacial cycle, not only in mid-latitude mountain areas
(e.g. Gillespie and Molnar, 1995), but also in some of the large continental ice sheets (Ehlers
et al., 2011).
This paper builds on a recent edited compilation of global glaciations by the authors (Ehlers
et al. 2011) In the introduction to that volume the authors highlighted major differences in
the extent and timing of glaciations in different parts of the world – particularly with respect
to the global LGM. This paper elaborates on this theme and provides a focused review of the
timing of global glaciations during the last glacial cycle and highlights the numerous
complications that arise in understanding and differentiating the behaviour of mountain
glaciers, alpine ice caps, and continental and polar ice sheets. The aims are to 1) examine
the evidence for the timing of the most extensive phases of glaciation around the globe, 2)
critically evaluate the definition and concept of the global LGM from a glacier perspective,
and 3) discuss the palaeoclimatic significance of asynchronous glaciations during the last
glacial cycle.
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2. The maximum extents of glaciers during the last glacial cycle – a global view
The following sections review the evidence for the timing of glaciations during the last
glacial cycle around the world. This review must be selective given the quantity of data
available. Case studies were selected on the basis of geographical representation. Where
conflicting or contradictory evidence exists, it is included. this is not intended to be an
exhaustive review and it is inevitable with a global survey such as this that some important
literature in specific regions might be overlooked or regions where little is known about the
geochronology of glaciations during the last glacial cycle are not mentioned, such as the
Caucasus where very little dating has been carried out (Gobejishvili et al., 2011). Still, this
should not detract from the broader aim of the paper, which is to evaluate the validity and
applicability of the concept of a global LGM with reference to a wide range of glacial
deposits and landforms from around the world. Local glacial events are sometimes correlated
with stages and substages in the ocean oxygen-isotope record as a global reference of
geological time in the Pleistocene. It is important to realise that this does not necessarily
imply any palaeoenvironmental link, but simply time-equivalence, a situation highlighted in
the later discussion section. Complicating this correlation is the fact that the boundaries of the
marine isotope stages are not formally fixed in time, with minor variability noted in the
literature (cf. Table 2 in Sanchez Goñi and Harrison, 2010; Gibbard and West, 2000; Gibbard
and van Kolfschoten 2005). In the following review the boundaries are taken as follow: MIS
2/3, 29 ka; MIS 3/4, 59 ka, MIS 4/5, 74 ka (based on Martinson et al., 1987 and Sanchez
Goñi and Harrison, 2010) (see Fig. 1).
2. Establishing geochronological frameworks for glacial sediments and landforms
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This paper is concerned with reviewing the timing of glaciations across the globe during the
last glacial cycle. In the last decade there has been an exponential increase in the number of
dates associated with glaciations around the world. This is most striking difference between
the successive volumes reviewing global glaciation by Ehlers and Gibbard (2004) and Ehlers
et al. (2011). Some techniques directly date glacial landforms whilst others provide ages for
associated deposits. Each of the different techniques have strengths and weaknesses and some
techniques are used much more than others (especially radiocarbon and cosmogenic exposure
age dating). This section provides a very brief overview of the main contexts, problems and
prospects for the main dating techniques. The purpose of this is to provide the reader with
some background on the different geochronological approaches that are commonly used to
determine glacial histories around the world.
2.1 Radiocarbon dating
Until the late 20th Century, radiocarbon dating dominated the geochronological frameworks
established for glaciated terrains. This technique is limited to dating Late Pleistocene organic
material that is <50 ka in age. The theory of radiocarbon dating is well known and widely
reviewed elsewhere (e.g. Walker, 2005). This section focuses on the contexts where
radiocarbon ages are commonly applied and recent challenges that have arisen with this
technique in glaciated settings.
In some contexts radiocarbon ages can provide maximum ages for glaciation, when
organic material has become incorporated within moraines during a glacial advance (Szabo et
al, 2011), although this is more commonly applied to date mountain glacier advance and
retreat during the Holocene since trees are more abundant in glacier forelands during
interglacials than glacials (e.g. Hormes et al., 2001).
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More commonly for the Pleistocene glacial sequences, radiocarbon dates provide
minimum ages for deglaciation. Radiocarbon dates from basal organic deposits in glacial
lakes offers the opportunity to date the duration of lacustrine sedimentation in a glaciated
basin. This can provide important information about the deglacial history of an area. In the
case of the Laurentide ice sheet, deglacial histories were based on thousands of radiocarbon
ages (Dyke, 2004). However, in this region and others radiocarbon dating has been replaced
by cosmogenic exposure dating as the dominant technique. Nevertheless, radiocarbon dating
is still widely used to date basal sediments in lakes of glacial origin. Usually, this is by
scientists interested in lacustrine sediment records (e.g. Moreno et al., 2009) than by
geomorphologists interested in the ages of associated moraines, although sometimes lake
sediments are also utilised directly by geomorphologists (e.g. Serrano et al., 2012).
Radiocarbon dates from glacial lake sediments provide only minimum ages for
deglaciation. If the time elapsed between glacier retreat and ecological development in a
catchment is large then lake basins will simply be filled with minerogenic sediments.
Sometimes this will be coarse-grained and many coring investigations (especially when
manually extracted) bottom in sands or gravels whose remaining thickness is unknown. For
example, in the Pindus Mountains of Greece, Willis (1992) retrieved a 12 m core from marsh
sediments that bottomed in sands. The oldest radiocarbon age from the base of this sequence
was 9.89 ± 0.12 14C ka BP (11.3 ± 0.11. cal. ka BP) whereas the moraines within which the
basin is enclosed contain calcite cements that are Middle Pleistocene in age (Hughes et al.,
2006a).
In some areas, apparent conflicts have arisen between the results arising from
different dating techniques, such as radiocarbon and cosmogenic dating techniques ―
especially where ice was cold-based and non-erosional, preserving relic landforms and
sediments (and associated 14C ages predating the LGM). This has been the case in Arctic
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Canada where old radiocarbon ages (at the limit of the technique) have suggested that some
areas remained ice free during the last cold stage. However, cosmogenic exposure ages has
revealed that these areas were ice-covered, only that the ice was cold based resulting in
minimal erosion and preservation of ‘old’ organic material (Davis et al., 2006b) – see section
3.1.1 below. In other areas there have been contradictions between deglacial chronologies
based on radiocarbon ages from glacial lake sediments and exposure ages obtained using
cosmogenic nuclide techniques. This is particularly true for Iberia and the Pyrenees – see
section 3.4.5 below.
All radiocarbon ages require calibration for comparison with calendar ages obtained
using other techniques. In this paper all original radiocarbon ages are quoted and these are
also calibrated into calendar years (at 1 σ error) using the IntCal09 calibration (Reimer et al.,
2009) via the Calib 6.0 calculator (http://calib.qub.ac.uk/calib/).
2.2 Cosmogenic exposure age dating
Cosmogenic exposure dating has revolutionised glacial geomorphology in the last two
decades. Previously, glacial sediments were notoriously difficult to date, often relying on
minimum limiting ages from radiocarbon dating of glacial lake sediment sequences.
Comprehensive reviews of the principles of cosmogenic exposure dating are provided in
Gosse and Phillips (2001) and Cockburn and Summerfield (2004) with a thorough overview
on the application of this technique to glacier chronologies provided by Balco (2011). The
technique is based on the fact that cosmic radiation is continually bombarding the Earth.
Most of this radiation originates from within our galaxy (including from our Sun) with a
small proportion of higher energy radiation originating from outside our galaxy. When it
reaches the upper atmosphere this galactic cosmogenic radiation interacts with the Earth’s
magnetic field producing a shower of secondary radiation that reaches the Earth’s surface.
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The primary radiation reaching the earth from space is dominated by protons whilst the
secondary radiation reaching the earth’s surface consists predominantly of neutrons
(Cockburn and Summerfield, 2004). These neutrons (along with much smaller amounts of
muons) interact with minerals in rock surface producing in situ terrestrial cosmogenic
nuclides. The longer the exposure of the rock surface, the greater the quantity of in situ
terrestrial cosmogenic nuclides in that rock surface. Whilst cosmogenic exposure age dating
has made an enormous contribution towards understanding glacial chronologies around the
world, there are several issues which must be noted in order to be able to critically review the
datasets that have emerged.
Secondary cosmogenic radiation is strongly attenuated with depth through rock
surfaces. Thus, the top few centimetres of intact rock surfaces are most suitable for
cosmogenic exposure age dating. An obvious limitation to this technique occurs when the
cosmogenic signal is not zeroed by glaciers because of insufficient erosion and thus rock
surfaces yield an inherited cosmogenic nuclide signal. Glaciers erode bedrock and when this
involves the removal of a thickness of more than c. 2 m from the rock surface then any
previous build-up of in situ terrestrial cosmogenic nuclides is removed and the cosmogenic
signal is zeroed. The same principal applies to rock fragments that originate from glacier
erosion and that are transported through the glacier system and subsequently deposited as
glacially-transported boulders (erratics, moraine boulders etc). If glacier erosion has not been
sufficient to zero the cosmogenic nuclide signal then apparent ages will be older than the time
elapsed since the last exposure of that surface.
Conversely, erosion of moraines can result in the exhumation of boulders revealing
their surfaces to the atmosphere after the moraine itself was formed. This results in exposure
ages that are younger than the glacier event which formed the moraine. Similarly, toppling of
boulders can also produce ages which are younger than the age of the moraine-building
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event. This often produces a wide scatter of ages from the same moraine. Putkonen and
Swanson (2003) argued that or identical dating accuracy, six to seven boulders are needed
from old and tall moraines (40 to 100,000 ka, 50 to100 m initial height) but only one to four
boulders from small moraines (20 to 100 ka, 10 to 20 m initial height). Many papers present
less samples than this from moraine surfaces. Whilst this does not necessarily mean that
exposure ages are wrong, some caution must be given to interpreting results of such papers.
In the case of moraine boulders, Heyman et al. (2011) argued that most boulders are
more likely to be influenced by incomplete exposure rather than inheritance, yielding ages
that are younger than the true age of the moraine. Thus, Heyman et al. (2011) argue that
glacial boulder exposure ages should be viewed as minimum limiting deglaciation ages. In all
instances, the geomorphological context can be crucial when interpreting noise and scatter
with exposure age datasets from moraine boulders (e.g. Briner et al., 2005a; Zech et al.,
2005a; Dortch et al., 2009). As for bedrock surfaces, this also depends on the
geomorphological context as well as the sample size. Inheritance can be a particularly
important issue when ice was cold-based and non-erosive. Even when rock has clearly been
eroded by warm-based ice, the problem can still persist and Balco (2011) noted that there is
“no obvious way to determine in advance of sampling whether a striated bedrock surface has
been fully reset by subglacial erosion or not”. It is often only after analysing data from
multiple samples that inheritance can be negated. In their study of glaciated bedrock surfaces
in the SW British Isles Rolfe et al. (2012) found a tight cluster of 10Be and 26Al exposure ages
from bedrock surfaces from ice-moulded granite bedrock and old outliers were able to be
identified. However, for both bedrock and moraine surfaces from mountains in NW England
results have yielded a large scatter of 36Cl exposure ages with inheritance appearing to be a
significant problem in a majority of samples (Ballantyne et al., 2009; Wilson et al., 2013). In
several cases where there has been complex exposure history, erratic boulders and bedrock
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yield different ages and paired isotopes (10Be and 26Al) can help identify complex exposure
history (Fabel et al. 2002; Briner et al., 2003; 2005b; Miller et al., 2006). However, even
paired isotopes are not able to identify complex exposure histories on timescales of several
tens of thousands of years (i.e. within the last glacial cycle – see discussion in Rolfe et al.,
2012). To get around the problem of inherited 10Be and 26Al in determining exposure of rock
surfaces previously covered by a cold-based area of the Laurentide ice sheet, Miller et al.
(2006) exploited in situ 14C in rock surfaces since this nuclide has a much shorter half-life
than 10Be or 26Al.
Cosmogenic exposure ages from moraine boulders are usually interpreted as relating
to a glacier event such as an advance resulting in the building of a moraine. However, there is
ambiguity whether such a surface age represents the timing of ‘retreat’ or ‘advance’. For
example, does the exposure age represent the timing of retreat from the moraine? This can be
particularly important if a glacier has been stationary for some considerable time at a moraine
position. Even so, given that the error associated with the dating technique is c. ± 9-10%,
such high-resolution problems are not so significant, especially for older moraines. Based on
modern analogues it is rare that glaciers are situated at the same moraine position for longer
than hundreds of years. However, care must be taken when interpreting cosmogenic exposure
ages because at the resolution of the technique ages are likely to be able to pick out complex
and rapid glacier behaviour. These types of issues, and others, were highlighted by Kirkbride
and Winkler (2012) in their review of the correlation of late Quaternary moraines.
Some note must be made of the differences between in analytical (AMS only) and
absolute (AMS + production rate) age errors. They are c. 2-5% for former and >10% for
latter. When comparing ages between samples from a single locality or glaciated region, then
age errors do not need to include production rate errors because the locality will have a
similar production rate at any given altitude. However, when comparing ages from across the
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Earth’s surface where production rates are so varied the uncertainty in production rates means
that the errors are higher. This is an impediment to global correlation and crucial when
discussing the LGM when 10% error on 20 ka means a >2 ka error. This means that it is not
possible to differentiate between glacier events at millennia-scale resolution. Thus it would
not be realistic to try and separate out events at 21 ka and 18 ka, for example, using exposure
ages from different parts of the world. . Nevertheless, cosmogenic exposure ages are certainly
able to differentiate between global scale events that are >5 ka (> c. 20%) deviant from the
global LGM within MIS 2.
There have been changes in calculations of production rates over the past 10 years
which change some ages (e.g. Ballantyne, 2012). This is especially important for older ages
(>200 ka) but it does not impact significantly on ages <100ka, although is significant for
cases where fine-scale resolution is important (such as for the Younger Dryas Chronozone –
see review by Ballantyne, 2012). As noted above, this level of resolution is in any case
questionable for a global review such as this. Finally, recent changes in the 10Be half-life and
AMS standard calibrations (Fink et al., 2007; Nishiizumi et al., 2007; Chmeleff et al., 2010)
will not alter ages cited in this review, which focuses on the last glacial cycle only.
2.3 Other techniques
Several other techniques have been applied to date deposits associated with glaciers that
formed during the last glacial cycle. Whilst not as prevalent as radiocarbon dating or
cosmogenic exposure age dating, OSL and U-series dating are particularly notable and their
contexts are reviewed very briefly below.
2.3.1 Optically stimulated luminescence (OSL) dating
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OSL has been used in many parts of the world to date deposits relating to glaciations of the
last cold stage. When sediments become buried, electrons accumulate in traps within the
crystal lattice of minerals such as quartz and feldspar as a result of ionising radiation from
within the sediment and cosmic radiation. The longer the burial time the greater the
accumulation of electrons within crystal lattices. The electrons are released when stimulated
by light, in the case of optically stimulated luminescence – the most common luminescence
technique used to date glacial and associated sediments. This technique has the advantage of
being able to date sediments panning the entire Quaternary and is widely used in dating
Quaternary sediments (Duller, 2004).
OSL is often used to date proglacial glaciofluvial sands and, more rarely, glacial
deposits themselves. This is because the technique relies on mineral grains such as quartz
being bleached by daylight prior to deposition and there may be doubts about this where
sediments originate from glacial erosion and are then transported and deposited and buried
whilst under ice. Fuchs and Owen (2008) provide a thorough review of the applications of
luminescence dating in glacial and associated sediments. They identify several problems such
as insufficient bleaching, low sensitivity of quartz, and variable dose rates during the history
of the sediment due to changing water content or nuclide leaching. In addition to inadequate
bleaching, Lukas et al. (2007) also suggest that glaciers can re-work inherited (older) moraine
material. In their case study in NW Scotland they were unable to obtain ages for Younger
Dryas sediments using OSL. Nevertheless, Fuchs and Owen (2008) argue that many of the
problems associated with using OSL to date glacial and related sediments can be overcome
with careful site selection, sampling, testing for bleaching and modelling dose rate variability.
2.4 Uranium (U) -series dating
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U-series dating has been used to date secondary carbonates (calcites) that are found
cementing moraines and tills. The most useful and versatile U-series method in Quaternary
studies is the 230Th/234U method based on the 238U series (van Calsteren and Thomas, 2006).
This method of U-series dating relies on the propensity of uranium, a very soluble element in
water, to co-precipitate with calcium during carbonate formation. Its daughter isotope, 230Th,
does not co-precipitate during carbonate formation due to its total immobility in the near-
surface environment (Langmuir and Herman 1980). As a result, at the time of formation, a
secondary carbonate deposit such as calcite cement, contains 234U and no 230Th, the daughter
isotope. The subsequent ingrowth of 230Th provides and its ratio with parent isotope 234U
provides a chronometer with which to measure elapsed time since carbonate deposition. The
practical range of U-series dating using 230Th/234U is c. 350 ka, although with the development
of sensitive analytical techniques of mass spectrometry, the range of this technique is
potentially <100 to 600,000 years (Ku 2000).
The dating of secondary carbonates via U-series relies on several assumptions and
criteria for application Upon deposition, the carbonate should be entirely free of 230Th. Most
non-authigenic 230Th is introduced bound to silicate or organic material. The presence of
detrital 230Th can be deduced from the presence of the isotope 232Th. This long-lived isotope
occurs in water as a trace impurity adsorbed onto particles and, where present, indicates that
contamination has occurred. In these instances, sample dates need correcting to account for
the detrital component. Methods of correcting for non-authigenic 230Th incorporated from
detrital contamination are described in Hellstrom (2006) and van Calsteren and Thomas
(2006). Another assumption is that the system should remain closed to the migration of
uranium and thorium after deposition. This will not be the case if a sample is characterised by
recrystallisation, solution, secondary precipitation or high porosity (Smart 1991). It is
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therefore important that the physical characteristics of the sample are assessed before
considering it for dating.
When used to date secondary carbonates within moraines, the technique assumes that
the calcite cements formed after the moraines were deposited. It is useful for providing
minimum ages for glacial events and since the cements often form during interglacials the
technique can help bracket moraines within certain glacial cycles (e.g. Kotarba et al., 2001;
Dehnert et al., 2010; Hughes et al., 2006a; 2010; 2011). However, it lacks precision (with
respect to the glacier event – not the U-series dating itself) because the calcites can form
sometime after the moraines were laid down. Consequently it has seen limited use for dating
moraines of the last glacial cycle, although minimum ages for deglaciation have been
obtained from parts of the Balkans (Hughes et al., 2010). U-series dating can also be applied
to date glaciofluvial sediments, where groundwater calcretes can be closely related to the
timing of deposition. In this situation, the U-series ages may be closer in time to the
depositional event and help provide a more precise approach to dating glaciofluvial sediments
(e.g. Woodward et al., 2008).
3.1 North America
3.1.1 Laurentide Ice Sheet
The Laurentide Ice Sheet represented one of the largest ice masses on Earth during the last
glacial cycle (Wisconsinan Stage), on a scale comparable to Antarctica (Fig. 2), and covered
almost all of Canada and reaching well into the northern parts of the contiguous United States
(Dyke et al., 2002; Dyke, 2004). The ice volume stored in the Laurentide Ice Sheet alone was
c. 37 x 106 km3 (Sugden, 1977), more than four times the size of the Fennoscandinavian Ice
Sheet which had a maximum volume of c. 8 x 106 km3 during the Late Pleistocene (Siegert,
2001). Understanding the timing of the maximum extent of this ice sheet is therefore crucial
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for understanding its contribution to the signal of global ice volume and global sea levels.
Various papers have attempted to resolve the extent and timing of glaciation during the last
cold stage and a thorough review is given in Dyke et al. (2002) and Dyke (2004), although at
that time questions still remained as to the limits and timing of the maximum phase of
glaciation in many areas. Furthermore, geochronologies were largely based on radiocarbon
dating. Recent advances in cosmogenic nuclides have led to a rapid growth in
geochronological data from areas covered by the Laurentide Ice Sheet (Briner et al., 2006).
Whilst a majority of the geochronological evidence supports the most extensive phase of the
Laurentide Ice Sheet during MIS 2, ice sheet modelling suggests that ice volume during MIS
4 (at c. 65 ka) was only c. 20% smaller (Stokes et al., 2012) (Fig. 3). Some of the evidence
for the timing of the most extensive phases of glaciation is outlined below from key areas of
the former Laurentide Ice Sheet.
Many parts of the southern and eastern sectors of the Laurentide Ice Sheet reached
their maximum extents in MIS 2 and deglaciation occurred after the global LGM. In
Newfoundland, Gosse et al. (1996) dated erratic boulders from summits that were previously
considered to have been nunataks during the last cold stage (e.g. Grant, 1977). Gosse et al.
(1996) found that the boulders yielded exposure ages of 13.1-20.9 ka suggesting deglaciation
following the global LGM. In the wider area of Québec–Labrador, the Laurentide Ice Sheet
retreated onto land after c. 13.5 ka and persisted until 6 ka (Oschietti et al., 2011).
In southeast Connecticut, at c. 41ºN latitude, cosmogenic exposure age and varve
chronologies indicate that deglaciation occurred at 18.5-19 ka (Balco and Schäfer, 2006). At
Martha’s Vineyard in Massachusetts, moraine boulders yield a range of ages that cluster
between 22 and 25 ka. Older and younger outliers in this dataset are considered to be either
caused by an inherited nuclide signal or boulder disturbance, respectively (Balco et al., 2002).
Radiocarbon ages up to 22 cal. ka BP on postglacial sediments north of moraine limits and
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>24 cal. ka BP beyond the outermost moraines on pre-glacial sediments offshore of
northeastern USA show that the most extensive glaciation in this area occurred at 22-25 cal.
ka BP (Balco et al. 2002 and references therein). It is clear therefore that the southeastern
sector of the Laurentide Ice Sheet was in synchrony with global ice volume recorded in the
ocean-isotope and global sea-level records.
In Ohio, the maximum extent of Laurentide ice during the Wisconsinan Stage is
constrained by numerous radiocarbon ages (Lowell et al., 1999; Szabo et al, 2011 and
references therein). Average radiocarbon ages from organic material and trees overrun by ice
in southwestern Ohio range from 23.2 to 19.6 14 ka BP (c. 27.9 to 23.4 cal. ka BP) with ice
reaching its maximum limit just north of Cincinnati, c. 10 km north of 39ºN, at 23.4 cal. ka
BP (Lowell et al., 1999; Szabo et al, 2011).
In Illinois, the ice reached its maximum extent at a similar time. Here, the Lake
Michigan ice lobe blocked drainage of the ancient Mississippi at c. 24.3 cal. ka BP, diverting
the river to its modern course, before reaching its maximum position at 23 ka at c. 23 cal. ka
BP (Curry et al., 2011).
Further north, Laurentide ice had crossed Lake Superior into Wisconsin, USA, by c.
32 ka (Syverson and Colgan, 2011) and reached its maximum extent between 20-18 ka based
on Optically Stimulated Luminescence dating (OSL) (Attig et al., 2011) and cosmogenic
exposure ages (Colgan et al., 2002). OSL samples from two small ice-marginal lakes in the
Baraboo Hills of Wisconsin yielded mean ages of 18.2 ka (Attig et al., 2011). This data
provides evidence that the Green Bay Lobe of the Laurentide Ice Sheet “stood at or very near
its maximum extent until about 18.5 ka” (Attig et al., 2011, p. 384). Colgan et al. (2002)
measured 10Be and 26Al in 22 quartz-rich samples from striated bedrock surfaces in south-
central Wisconsin. At two sites the cosmogenic ages matched the radiocarbon chronology
with exposure ages of between 13 and 21 ka. This is consistent with an ice maximum in MIS
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2. However, at three other sites the cosmogenic ages were much older and Colgan et al.
(2002) suggest that there is a problem with nuclide inheritance.
In Montana, USA, a total of 27 ice-rafted boulders deposited in ice-marginal Lake
Musselshell (c. 48ºN) have yielded 10Be exposure ages of between 5.2 ± 0.3 to 21.7 ± 0.6 ka.
Davis et al. (2006a) used this and other data to suggest that the lake existed between 20 and
11.5 ka when there were at least three advances of the Laurentide Ice Sheet into this area. The
extent of Laurentide ice incursion into North Dakota was similar to that in Montana and is
recorded by erratic boulders perched on bedrock (Davis et al., 2006a), six of which have
yielded 36Cl exposure ages of 23-28 ka (Manz et al., 2005).
In NW Canada, Zazula et al. (2004) used radiocarbon to date pro-glacial lake
sediments in the northern Yukon to determine the minimum timing of deglaciation. They
interpreted this evidence as indicating a maximum extent of the Laurentide Ice Sheet between
35 and 22 ka followed by a readvance at 22-16 ka. Earlier work by Duk-Rodkin et al. (1996)
suggested that the maximum extent of the Laurentide Ice Sheet in this area was reached by 30
ka, based on 36Cl exposure dating of glacial boulders. Thus, the timing of the glacial
maximum in the Yukon appears to have been slightly earlier than in other sectors of the
Laurentide Ice Sheet with a late MIS 3 maximum, as opposed to an MIS 2 maximum
elsewhere.
In addition to investigating the timing of the lateral extent of the various sectors of the
Laurentide Ice Sheet, cosmogenic nuclide analyses have been applied to determine changes
over time of the thickness of the Ice Sheet in Arctic Canada. It has become apparent through
this research that the Laurentide Ice Sheet was cold-based over large areas, preserving
without modification , basal sediments, glacial deposits (such as moraine boulders),
underlying landforms and, importantly, leaving organic sediments older than the last ice
maximum buried and untouched by glacial erosion. For example, Davis et al. (2006b) have
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shown that perched cobbles yielding cosmogenic exposure ages of c. 14 ka rest on a
glaciomarine delta that is older than 54 ka, indicating preservation of ancient landscapes
under the Laurentide Ice Sheet. High-level tors and other relict periglacial landforms, pre-
dating the last period of ice cover, have also been reported from Baffin Island (Bierman et al.,
1999; Briner et al. 2003; 2005b) and Labrador (Marquette et al. 2004). In both of these
localities bedrock samples yield cosmogenic exposure ages ranging from 60-150 ka, whilst
perched erratic boulders yield exposure ages of 11-17 ka. The latter are consistent with the
thickest phases of the Laurentide Ice Sheet close to the global LGM during MIS 2.
3.1.2 Cordilleran Ice Sheet
The glaciers of western Canada, forming the Cordilleran Ice Sheet, reached their maximum
extents at c. 20 ka, (17 14 ka BP) towards the end of the global LGM (Booth et al., 2003;
Clague and Ward, 2011; Porter, 2011; Rutter et al., 2012). Both radiocarbon and 36Cl ages
indicate that glaciers expanded after 29 ka and reached their maximum extent between 24 and
20 ka (Duk-Rodkin et al. 2004). There are, however, local exceptions. For example, on
northern Vancouver Island glaciers reached their maximum after 16 14 ka BP (Ward et al.,
2003; Al-Suwaidi et al. 2006). In Western Washington, USA, the Cordilleran Ice Sheet
reached its maximum extent by c. 17 ka and began retreating shortly afterwards with
stillstands or readvances at 15.8 and 14.7 ka (Porter and Swanson, 1998). In the Puget
Lowland ice reached a position c. 100 km south of Seattle (Porter, 2011). An earlier advance,
associated with the ‘Possession Drift’, has been ascribed to MIS 4 based on luminescence and
amino-acid ages (Easterbrook, 1994). This glacial phase is thought to have been less
extensive than the later MIS 2 advance (Booth et al., 2003), although Borden and Troost
(2001) suggest that ice during the former phase reached Tacoma to the south of Seattle
implying that the ice extent was only slightly smaller during MIS 4 in this area.
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In the east, the big debate has centred on whether the Laurentide and Cordilleran ice sheets
converged during the LGM or an ice-free corridor existed at this time. Jackson et al. (1997;
1999) dated 19 boulder samples using 36Cl in Alberta, a key area for testing the convergence
versus ice-free debate. All but one sample of 19 yielded ages in the range 12-18 ka, strongly
supporting the idea of ice convergence between the Laurentide and Cordilleran Ice Sheets in
MIS 2. However, a key difference between the Laurentide and the Cordilleran (western,
southern and central sectors) Ice Sheets is that the latter reached its maximum several
thousand years later than the various sectors of the Laurentide Ice Sheet, and after the global
LGM.
Unlike the western, southern and central sectors of the Cordilleran Ice Sheet, the
northern sector appears to have reached its maximum extent early in the last glacial cycle.
Ward et al. (2007) applied 10Be analyses to date the exposure age of glacial boulders in the
Yukon, Canada. Four boulders yielded 10Be exposure ages of 50.4 ± 3.0, 51.4 ± 2.6, 52.9 ±
2.2 and 54.3 ± 2.0 ka. The mean age of the four boulders 53.1±15.9 ka (including 2σprecision
with systematic error). Assuming that these samples do not contain inherited nuclides then
this is compelling evidence of an MIS 4 glaciation in the Yukon and highlights potential for
major regional variability in the timing of the most extensive glaciations. However, it must be
noted that these are based on single rather than paired isotopes and the possibility of complex
exposure histories (pre-exposure and burial) cannot be discounted.
3.1.3 Alaska
In Alaska, the timing of the maximum extents of the ice masses over the Alaska Range,
Ahklun Mountains and the Brooks Range occurred early in the last glacial cycle, during the
Early Wisconsinan (MIS 4) (Kaufman et al., 2011) (Fig. 4). In the Donnelly Dome area of the
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Alaska Range, Matmon et al. (2010) dated the Delta moraine complex using 10Be and yielded
exposure ages between 41.5 ± 4.5 ka and 70.8 ± 7.8 ka (n=8). They interpret this as indicating
the stabilisation of the Delta moraine complex near the MIS 4/3 boundary and that later
glacier advances during MIS 2 did not override these deposits, although post-depositional
processes have resulted in younger exposure ages on other parts of the moraine complex. In
the north-central Alaska Range, Dortch et al. (2010a) found that nine boulder samples, dated
using 10Be, yielded exposure ages between 56 and 52 ka (excluding one young outlier).
Similarly, in the western part of the Alaska Range, Briner et al. (2005a) found that moraines
surfaces yielded 10Be ages suggesting moraine stabilisation between 58 and 53 ka. Thus, it is
clear that ice reached a maximum extent between 58 and 52 ka in the Alaska Range during
the Wisconsinan (Kaufman et al., 2011) – some 30 ka before the global LGM. Similar
findings have been reported from the Ahklun Mountains (Briner et al., 2001; Kaufman et al.,
2001; Manley et al., 2001) and the Brooks Range (Hamilton et al., 2001). Whilst all of the
massifs experienced glacier maxima at a similar time, relative extents compared with later
MIS 2 glaciers advances differed markedly. For example, at the Brooks Range, Seward
Peninsula and Ahklun Mountains, and western Alaska Range the MIS 4 glaciation was much
larger than the MIS 2 glaciation (Fig. 4). In contrast, in the northern parts of the Alaska
Range the difference in size between the two glacier phases was much smaller (Kaufman et
al., 2011).
3.1.4 Western United States
South of the Laurentide and Cordilleran Ice Sheets, smaller ice caps and glaciers formed over
the Rocky Mountains during the Pleistocene (Pierce, 2003). The geochronology of successive
glaciations is now known in several areas – especially in the Cascades and the Sierra Nevada.
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In the NE Cascade Range, Washington, Porter and Swanson (2008) applied 36Cl
analyses to provide 76 exposure ages for granodiorite and gneiss boulders on seven
successive moraines. The moraines of the most extensive glaciation yielded a mean exposure
age of 105.4 ± 2.2 ka (1 σ uncertainty, analytical error only). Porter and Swanson (2008)
argued that the glacier advances were associated with July insolation minima at this latitude
(47.5ºN) (Fig. 5). They also point out that the largest glacier advance during MIS 5d is
consistent with major advances recorded elsewhere in North America, such as in Wyoming
(Chadwick et al., 1997; Phillips et al., 1997). In the Olympic Mountains, Thackray (2001)
also found evidence for large glacier advances in the early part of the last glacial cycle. More
than 60 radiocarbon ages from tills and outwash deposits provide limiting ages for the glacier
sequence in this area. The biggest advance (Early Wisconsinan Lyman Rapids advance)
occurred in MIS 4/5 followed by successively smaller advances during MIS 3 and MIS 2
(Thackray, 2001).
In the Sierra Nevada the moraines associated with the largest glaciers of the last
glacial cycle are largely within 2-3 ka of the commonly cited range of the global LGM (26-21
ka). For example, on the eastern flank of the Sierra Nevada, Phillips et al. (2009) found that
moraine exposure ages ranged from 28 ka to 14 ka and can be subdivided into advances at
28-24, 18.5-17.0 and 16.0-14.5 ka. In the South Fork of the Yuba River, Sierra Nevada, 10Be
and 26Al exposure ages indicate that the maximum extent of the glaciers during the last glacial
cycle occurred at ca. 18.6 ± 1.2 ka (James et al., 2002). Glaciers then rapidly retreated from
their maximum extents, taking only 1000 or 2000 years to disappear (Gillespie and Clark,
2011).
Many other areas of the Rockies exhibited similar glacier chronologies to the Sierra
Nevada with maximum ice extent predominantly occurring in MIS 2 (Thackray et al., 2008b).
In the Wallowa, Wind River and the Sawtooth Mountains, maximum ice extents
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occurred ca. 21 ka, 24–17 ka and 20–23 ka, respectively (Licciardi et al., 2004; Licciardi and
Pierce, 2008; Sherard, 2006), whilst in the Yellowstone region and Uinta Mountains ice
reached its maximum extent ca. 22–16 ka, (see review by Thackray, 2008b and later paper
from Uinta Mountains by Laabs et al., 2011). In many of these areas the glacial
geochronologies have been established using cosmogenic nuclide analyses, such as 3He and
10Be in the Yellowstone area of Montana (Licciardi et al., 2001) and 10Be in the Uinta
Mountains of Utah (Laabs et al., 2009).
Whilst the precise ages of the maximum glacial extents will no doubt vary in future as
different calibration schemes are applied to recalculate ages, the general theme is similar. The
glacier maximum in many areas of the Rockies was close in time to the global LGM within
MIS 2, with only minor millennial scale variability in the timing of deglaciation from this
maximum (e.g. Laabs et al., 2011). The major exceptions are in the NE Cascade Range in
Washington and also in the mountains of Wyoming where glaciers appear to have reached a
maximum as early as MIS 5d. Thus, it is clear that the timing of glacier maximum was
regionally variable in the Rocky Mountains south of the Cordilleran and Laurentide Ice
Sheets.
3.1.6 Hawaii
In Hawaii, an ice cap over Mauna Kea reached its maximum extent at c. 20.6 ka during the
last cold stage with a readvance at c.16 ka (Porter, 2011). 3He and 36Cl ages from lava near
the summit of Mauna Kea and 14C ages from sediments in a nearby lake suggest that the ice
cap had retreated by c. 15-14 ka (Dorn et al., 1991; Porter et al., 2011). Pigati et al. (2008)
also used 36Cl to date the two phases of ice cap glaciation on the same mountain. However,
their ages suggest retreat from the most extensive position at 23 ka and from a smaller ice cap
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position at 13 ka. Despite these minor disparities in the timing of the two ice caps phases the
maximum ice cap phase was in close accord with the global LGM.
3.2 Central America
In Mexico, Pleistocene glaciers formed on several volcanoes (Vázquez-Selem and Heine,
2011). The geochronological evidence provides a mixed picture for the timing of glaciations.
On Iztaccíhuatl volcano, c. 40 km SE of Mexico City, the most extensive glaciation of the
last glacial cycle is dated to 21-17.5 ka based on 36Cl exposure ages (Vázquez-Selem and
Heine, 2011) – correlating closely with the global LGM. However, on La Malinche volcano,
only c. 60 km east of Iztaccíhuatl, the most extensive glaciation of the last glacial cycle is
ascribed to 36-32 14C ka BP (41-36.5 cal. ka BP) by Vásquez-Selem and Heine (2011) who
correlated this glaciation with one recorded on Ajusco volcano, just 10 km south of Mexico
City, where moraines yield ages of >27.2 14C ka BP (31.4 cal. ka BP) (White and Valastro,
1984; White, 1987). However, little more is known about the timing and extent of glaciations
prior to MIS 2. Glaciations in other parts of central America are poorly dated. In Costa Rica
and Guatemala tentative correlations have been made with the glaciations on Izatccíhuatl in
Mexico (Lachniet and Roy, 2011).
3.3 South America
3.3.1 Colombia
The maximum glaciation of the last glacial cycle in the Colombian Eastern Cordillera
predated the global LGM (Helmens, 2011). Geochronology is based on 14C ages from peat,
lake sediments and palaeosols from boreholes and exposures in formerly glaciated areas (van
der Hammen et al., 1980; Helmens, 1988, 1990; Helmens and Kuhry, 1995; Helmens et al.,
1997). The largest glacial advances during the last cold stage took place prior to 30 14C ka
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BP. At least two advances are recorded in the geomorphology. These moraines are situated c.
200 m below moraines ascribed to the LGM (Helmens, 2011). Overall, Helmens (2011)
identified a series of successively smaller glaciations at 43-38, 36-31, 23.5-19.5, 18.0-15.5,
13.5-12.5 and 11-10 14C ka BP (46-42.5, 41-35, 28-23, 21-18.5, 16.5-14.5 and 13-11.5 cal ka
BP. The successively smaller glacier during the later part of the last cold stage (after 30 14C
ka BP or 34.7 cal. ka BP) is thought to be the result of drier climatic conditions in the
Colombian Andes. Glaciers retreated from their LGM positions at c. 19.5 14C ka BP (23.4 cal
ka BP), readvancing shortly afterwards then retreating again just before 15.5 14C ka BP (18.7
cal ka BP.
3.3.2 Venezuela
The last cold stage in the Venezuelan Andes is represented by the Mérida Glaciation which
consists of two sets of moraines. During an Early Mérida Stadial ice descended to c. 2600-
2800 m a.s.l. whilst a later advance during a Late Mérida Stadial descended to c. 2900-3500
m a.s.l. The moraines of the Early Mérida Stadial are known to be beyond the range of AMS
radiocarbon dating (~60 cal. ka BP) (Mahaney et al., 2001; Dirszowsky et al., 2005) yet
younger than <90 ka (Kalm and Mahaney, 2011). The most extensive glaciers of the Late
Mérida Stadial occurred at c. 21.5 cal. ka BP based on calibration of a number of radiocarbon
ages (cf. references and review by Kalm and Mahaney, 2011).
3.3.3 Ecuador, Peru & Bolivia
In Bolivia and Peru, moraines associated with the largest advance of the last glacial cycle
have yielded 10Be ages suggesting that glaciers reached their maximum at c. 34 ka and were
retreating by 21 ka (J. Smith et al., 2005). They later revised this glacier maximum to 31 ka
(J. Smith et al., 2008). Farber et al. (2005) suggested that in the Cordillera Blanca of central
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Peru glaciers reached their maximum position by c. 29 ka at the latest and persisted at this
position until retreat began at c. 20.5 ka. In the Nevado Coropuna, southern Peru, Bromley et
al. (2009) dated moraine surfaces using 3He. They found that the largest glaciers of the last
glacial cycle reached their maximum position between 24.5 and 25.3 ka and argued that
tropical glaciers in South America reached their maximum close to the global LGM – in
common with glaciers at higher latitudes. In Bolivia, Zech et al. (2007; 2008) also found that
the largest glaciers of the last glacial cycle yielded exposure ages consistent with maximum
expansion close to the global LGM. Zech et al. recalculated the ages presented in J. Smith et
al. (2005) using age models and found that the ages of the largest glaciers was between c. 31-
25 using the Lal/Stone scaling system (Lal, 1991; Stone, 2000) and25-20 ka using the Lifton
(2005) scaling system in the calculator of Balco et al. (2008). Thus, the idea of a glacial
maximum preceding the global LGM in Bolivia and Peru (Smith et al., 2005) is at odds with
one where glaciers are thought to have exhibited similar behaviour to the global ice volume
signal with a maximum coeval with the global LGM (cf. La Frenierre et al., 2011).
3.3.4 Chile/Argentina
In the northern parts of Chilean Lake District glaciers reached their maximum extent between
29.4 and 14.5 14C ka BP (34.3 and 17.7 cal. ka BP) (Bentley, 1997; Denton et al., 1999b).
However, in the southern parts glaciers reached their maximum extent before 49.9 14C ka BP.
Denton et al. (1999) suggested that this old Llanquihue glaciation occurred during MIS 4.
However, in most of southern Chile glaciers are thought to have reached their maximum
close in time to the global LGM (Schäfer et al., 2006; Harrison and Glasser, 2011). For
example, in the far south in the Magellan Straits, successively smaller glacial advances
occurred at 24.6 ± 0.9, 18.5 ± 1.8, and 17.6 ± 0.2 ka (10Be ages ) (Bentley et al., 2005).
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Nearby to the east, in the Bahía Inútil, advances are dated at 20.4 ± 1.2 and 17.3 ± 0.8 ka
(Kaplan et al., 2008a).
Some of the best-dated moraines are found on the western side of the southern Andes
in Argentina. For example, at Lago de Buenos Aires Kaplan et al. (2004) used 10Be and 26Al
analyses to demonstrate that the most extensive glaciers of the last glacial cycle formed
moraines from 23.0 ± 1.2 to 15.6 ±1.1 ka, with the most extensive glaciation correlating
closely with the global LGM. No moraines dating to MIS 4 were found and older, more
extensive, moraines date from the Middle Pleistocene (Kaplan et al., 2005). In their review of
glaciation during the last glacial cycle, Kaplan et al. (2008b) highlight the differences
between the western side of the southern Andes in Chile and the eastern side in Argentina. At
the time of their review it was only in the west that MIS 4 glaciers appeared to have been
larger than those that formed during MIS 2. Kaplan et al. (2008b, p. 655) suggested that “the
west side of the Andes precipitation worked in combination with temperature to push ice
limits farther west during MIS 4 than in MIS 2”. However, Kaplan et al. (2008b) also note
that generalising patterns of relative ice extents in this region requires further research.
Recent work by Glasser et al. (2011) who applied 10Be analyses to date moraine
surfaces in the Lago San Martin valley, Argentina, illustrates that the extent of MIS 2 (and
global LGM) glaciation was smaller than previously thought. Here, the biggest glaciations
occurred much earlier in the last glacial cycle (Fig. 6). This observation is important because
it contradicts earlier findings from other areas on the eastern side of southern Andes, such as
at Lago de Buenos Aires, c. 400 km to the north, where no MIS 4 moraines were found
(Kaplan et al., 2004). Furthermore, it accords with Zech et al. (2008) who recognised that
there was evidence from numerous sources for an early 'local' LGM in Chile. The timing of
glaciations in the southern Andes during the last glacial cycle is clearly a complex issue and
the glacial record may exhibit major regional differences.
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3.4 Europe
3.4.1 British-Irish Ice Sheet
Clark et al. 2012 provide a comprehensive review of the geochronological evidence for the
timing of deglaciation in the British Isles during the last cold stage (Devensian). The British
Isles ice sheet reached its maximum at different times in different sectors. In many parts of
Britain, Ireland and the North Sea ice had reached its maximum extent by 27 ka. For
example, the earliest maximum advances recognised by Clarke et al. (2012) occurred by 27
ka in the North Sea and off the coast of NW Scotland. In central England the ice reached a
maximum between 25-21 ka and in the Celtic Sea between 23-20 ka. However, the situation
in the latter area is not so clear since Rolfe et al. (2012) have reported clear glacial erosional
landforms that yield paired 26Al and 10Be ages suggesting exposure since MIS 3. A similar
situation was suggested by Bowen et al. (2002) for large areas of Ireland and Scotland based
on 36Cl ages. However, a large-scale review by Ballantyne (2010) suggests that many of
Bowen et al.’s 36Cl ages are too old and the majority of data points to a glacial maximum
between 26-21 ka. Indeed, several papers present evidence that does indicate thick and
extensive ice coverage at the global LGM (e.g. McCarroll et al., 2010) with downwasting,
indicated by exposure of mountain summits in Wales, occurring shortly afterwards (Glasser
et al., 2012). Still, the evidence from Lundy in the SW sector of the British-Irish Ice Sheet
does indicate that erosive ice here did retreat earlier than the global LGM in this area. As
noted by Clark et al. (2012) different sectors of the ice sheet exhibited different behaviours in
space and time and unravelling such complexity is a major challenge to modelling the
behaviour of the British-Irish Ice Sheet during the last glacial cycle.
3.4.2 Fennoscandinavian Ice Sheet
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The Fennoscandinavian Ice Sheet represented one of the major global ice masses during the
Weichselian Stage. It would therefore be expected to make a major contribution to the signal
of global ice volume and global sea level change. Furthermore, the Fennoscandinavian Ice
Sheet coalesced with the Barents-Kara (Svendson et al., 2004) and the British-Irish Ice Sheets
(Graham et al., 2011) (Fig. 7).
The dynamics of the Norwegian sector of the Fennoscandinavian Ice Sheet are
reviewed in Mangerud et al. (2011). The ice sheet advanced westward onto the continental
shelf on several occasions during the last glacial cycle. The largest glacial advances were in
MIS 4 and MIS 2. According to Mangerud et al. (2011), the conclusion is now that the
Fennoscandinavian Ice Sheet was as extensive west and south of Norway during MIS 4 as it
was in MIS 2. However, further south in Denmark the largest glacial advance is dated to MIS
3 (Larsen et al., 2009; Houmark-Nielsen, 2010).
In Denmark six major ice advances crossed the country during 65-60, 55-50, 34-29,
29-27 and 23-19 ka (cf. review by Houmark-Nielsen, 2011). The largest of these was the
Ristinge advance. Larsen et al. (2009) and Houmark-Nielsen (2010) used OSL ages to argue
that this advance dates from 56-46 ka (55-50 in Larsen et al. 2009), during MIS 3. The ice
maximum during the later 'main advance' in MIS 2 is dated to 23-19 ka, based on radiocarbon
ages from shells. This advance was slightly smaller than the earlier Ristinge advance
(Houmark-Nielsen, 2011, their Fig. 5.1).
In central Poland the southernmost limits of the Fennoscandinavian Ice Sheet are
marked by the Lezno moraine (Vistulian Stage). Marks (2011) calibrated a radiocarbon age
of 26.1 cal. ka BP (22.2 14C ka BP) (Stankowska & Stankowski 1988), which provides a
maximum limiting age for the ice advance to the Leszno Moraine. Further north in Poland,
10Be ages from boulders on the Pomeranian moraine (c. 53-54 ºN) suggest that ice retreated
from this position at 14.8 ± 0.4 ka (Rinterknecht et al., 2005). However, in southern Sweden,
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10Be analyses appear to suggest that ice was situated north of the Vimmerby moraine (c.
57.5ºN) . from c. 14.4 ± 0.9 ka (Johnsen et al., 2008). This leaves only a short difference in
time between the maximum extent and the retreat position in southern Sweden. Furthermore
it is also possible that southern Sweden was ice-free at the global LGM itself, although
questions remain as to interpretation of OSL ages from this area (Alexanderson and Murray,
2007). Based on the majority of other geochronological evidence this seems unlikely and ice
would have reached Denmark and Poland at the global LGM. Moreover, in northeastern
Germany 10Be ages from the Gerswalder moraine, a recessional remnant of the Pomeranian
moraine, show that ice retreated from this area over the period from 16.6 ± 1.0 to 12.3 ± 0.6
ka – clearly illustrating persistence of Scandinavian ice in this area after the global LGM
(Rinterknecht et al., 2012).
In the southeast, in Poland, the largest ice advance dates from 24-19 ka (Marks, 2011)
during MIS 2 and close in time to the global LGM. Overall, three advances are recorded in
the southeastern sector of the Fennoscandinavian Ice Sheet between 25 and 12 ka. In Belarus
and Lithuania, Rinterknecht et al. (2006; 2007; 2008) used 115 10Be ages and 70 radiocarbon
ages to constrain the timing of these three substantial ice-margin fluctuations. The ice sheet
reached its maximum position after 20.98 ± 0.27 (21.0+/-0.3) cal. ka BP and retreated from
its maximum position after 19.0 ± 1.6 ka (Rinterknecht et al., 2006). In Latvia, the onset of
Late Weichselian glaciation has not been reliably dated, although OSL ages suggest ice
advanced into the country after 25-24 ka (Zelčs et al., 2011).
Thus, the southeastern sector of the Fennoscandinavian Ice Sheet appears to have
been synchronous with the global record of ice volume changes. This is in contrast to the
southwestern sector, which reached its maximum extent in MIS 4 and 3 (Houmark-Nielsen,
2011) – although a major ice advance did occur in this sector at 21-19 ka (Mangerud et al.,
2011) Thus, it is clear than the various sectors of vast ice sheets, like the Fennoscandinavian
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Ice Sheet, exhibit complex responses to the global climate signal, with the global LGM only
one of a series of major advances of varying size.
3.4.3 European Alps
During the last glacial cycle (Würmian Stage), the Alpine Ice Sheet was at its most extensive
between 30 and 18 ka (Ivy-Ochs et al. 2008; Preusser et al., 2011). This is based on
cosmogenic, OSL and radiocarbon ages from across the region. The last glacial cycle is
termed the Birrfeld Glaciation in northern Switzerland and comprises three independent
glacial advances at ~ 105 ka, 65 ka, and 25 ka. These correspond to MIS 5d, 4 and 2 but the
largest advance was in latter interval (Ivy-Ochs et al. 2008; Preusser and Schlüchter, 2004;
Preusser et al., 2011). In areas peripheral to the Alps, such as in the western Vosges, Seret et
al. (1990) also recognised a period of extensive glaciation between 57 and 30 14C ka BP (>57
and 34 cal. ka BP). A series of smaller phases of glaciation in the eastern Vosges has been
dated using cosmogenic 10Be and postdates the global LGM (Mercier et al., 1999) during the
Younger Dryas and Early Holocene. Evidence of an extensive early glaciation is not
replicated everywhere in the Alps and immediate neighbouring areas. For example, there is
no evidence for large glaciers in MIS 5d and 4 in the eastern Alps, although a large MIS 4
glaciation reaching the lowlands in the western Alps has been largely accepted (Ivy Ochs,
2008, p. 562).
Despite the uncertainty over the early advances, many areas show a the maximum ice
extent of the last cold stage was in the Late Würmian, in MIS 2 close in time to the global
LGM (Florineth and Schlűchter, 2000; Ivy Ochs et al., 2008; Buoncristiani and Campy, 2011;
van Husen, 2011). For example, the large outlet Rhône glacier in the French Alps had begun
to retreat from its maximum position at 21.1 ± 0.9 ka (Ivy Ochs et al., 2006). In the Southern
Alps in Italy, the Tagliamento glacier may have reached its maximum slightly earlier,
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between 26.5 and 23 cal. ka BP (Monegato et al., 2007). The Italian Alpine glaciers also
underwent major retreat after c. 21 ka (10Be age; Gianotti et al., 2008). This is consistent with
cosmogenic exposure ages from numerous sites across the Alps, which indicate that moraines
associated with the largest ice extent had stabilised by c. 21 ka (e.g. Ivy-Ochs et al., 2004;
2006). A later readvance is recorded across the Alps during the Gschnitz Stadial before the
onset of the Bølling Interstadial (before 14.7 ka). Ivy Ochs et al. (2006) suggest that the
moraines of the Gschnitz Stadial glaciers stabilised no later than 15.4 ± 1.4 ka based on 10Be
exposure ages. The Gschnitz Stadial was followed by a marked period of downwasting and
was followed by a series of Late-glacial and Holocene advances (Ivy Ochs et al., 2006).
3.4.5 Pyrenees
Radiocarbon ages from lake and bog sequences within former glacier limits on both the
northern and southern slopes of the Pyrenees suggest that maximum glacier advances during
the last glacial cycle were asynchronous with the record of global ice volume – preceding the
global LGM by tens of thousands of years in several areas (e.g. García Ruiz et al., 2003;
González-Sampériz et al., 2006). First, cosmogenic 10Be ages appeared to conflict with the
established chronology that argued for an ‘early’ glacier maximum in the Pyrenees and
suggested a local glacial maximum after 25 ka and close in time to the global LGM at
c. 21 ka. Pallàs et al. (2007) presented 25 10Be exposure ages for glacial erosion surfaces and
boulders in the south-central Pyrenees where the oldest exposure age was 21 ± 4.4 ka. Similar
findings were reported by Delmas et al. (2008) from the Carlit massif in the eastern Pyrenees.
However, more recent information presented by these researchers has revealed evidence for
an earlier glacial maximum during the last glacial cycle during MIS 5 (Pallàs et al., 2010;
Delmas et al., 2011). The initial contradictions between older radiocarbon and younger
cosmogenic ages appear to have been the result of insufficient evidence from comparable
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sites (Hughes, 2013). In the Cinca and Gállego River valleys (south-central Pyrenees and
Ebro Basin, Spain), OSL ages obtained from glacial deposits reveal glacial phases at
85 ± 5 ka, 64 ± 11 ka, 36 ± 3 ka with the largest extent of glaciers during the last glacial at
64 ± 11 ka (Lewis et al., 2009). Even so, a glacial event corresponding to the global LGM in
MIS 2 is clearly expressed in the mountains (Delmas et al. 2008, 2009) although the relative
positions of glacier fronts across the Pyrenees are still controversial. An excellent review of
the Pyrenees glacial history is provided by Calvet et al. (2011).
3.4.6 Iberia
As in the Pyrenees, the glacial history of the Iberian mountains has been the subject of two
competing hypotheses: one where glaciers reached their maxima close to the global LGM and
another where the glaciers reached their maximal positions much earlier in the last glacial
cycle (Jiménez-Sánchez et al., 2012). Unlike in the Pyrenees, consensus has not been reached
and the two scenarios are still strongly debated with apparently contradictory evidence from
different areas.
In the Sierra de Gredos and Serra de Guadarrama, cosmogenic 36Cl exposure ages
reveal clear evidence of an ice advance in MIS 2 reaching a maximum at c. 26 ka (Palacios et
al. 2011; 2012), In the Sierra de Gredos, the maximum advance is dated at 26-24 ka (36Cl)
and stabilised at this position for c. 3 ka followed by retreat after c. 21 ka (Palacios et al.,
2011). A similar situation has been found in the Sierra de Guadarrama, again with a
maximum advance at 26 ka, followed by stability until 20-19 ka. Domínguez-Villar et al.
(2013) combined 23 10Be exposure ages from moraine boulders with U-series ages from two
stalagmites. Two advances can be recognised in the geomorohological records and the
exposure age dataset, the largest occurring at 26.1 ± 1.3 ka and a smaller later advance
occurring at 21. 3 ± 0.7 ka. The most extensive advance at 26.1 ± 1.3 ka occurred within a
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cold and wet period indicated in speleothem records between 29 and 25 ka. This is similar to
findings from other parts of the Mediterranean such as in Italy (Giraudi, 2012) and Greece
(Hughes et al., 2006b).
In the Sierra Nevada, Gómez-Ortiz et al. (2012) also dated moraine surfaces using
36Cl. These exposure ages reveal that deglaciation was in progress at c. 15.4 ka and glaciers
had retreated from the high cirques by 13.2 ka. Moraines equating to the global LGM are
bracketed in time by exposure ages of 31.8 ka to 19.3 ka. However, compared with the Sierra
de Gredos and Serra de Guadarrama, the precise geochronology for the glacier maximum is
less clear, although region-wide correlations in Iberia do seem reasonable. Overall, there is
sufficient evidence to suggest that some glaciers were larger earlier in the last glacial cycle
than during MIS 2 and at the global LGM. For example, in the Sierra Nevada, Gómez-Ortiz
et al. (2012, p. 93) acknowledge that “remains of moraines older than the global Last Glacial
Maximum exist, but their poor preservation makes them an unreliable subject for surface
exposure dating”.
Geochronological evidence of larger glaciations earlier in the last glacial cycle has
been reported from the several sites in the Cantabrian Mountains. This is all based on
radiocarbon ages from sediment cores in basins within the glacier limits. For example, in the
Redes Natural Park of the Cantabrian Mountains, a radiocarbon date of 28.99 0.23 14C
years BP (33.49 0.36 cal. years BP) has been reported from the base of a core retrieved
from ice-dammed lake deposits that formed when drainage was blocked by a lateral moraine
(Jiménez-Sanchez and Farias, 2002). In the same study, radiocarbon dating of proglacial
deposits, interpreted as being synchronous with the local last glacial maximum in the nearby
Comella basin of the Picos de Europa, has yielded an age of 40.88 0.82 14C years BP
(44.64 0.71 cal. ka. BP) (Jiménez-Sanchez and Farias, 2002). More recent work in this area
has confirmed that the glaciers retreated from the Comella Basin by 43-45 cal. ka BP and
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nearby Lago de Enol by 38 cal. ka BP (Moreno et al., 2010). From the former glacial lake
basin of Campo Mayor, Serrano et al. (2012) retrieved a 20 m core. Radiocarbon ages of 31.2
0.44 14C ka BP (35.28 0.44 cal. ka BP) at c. 15.5 m depth provide a minimum age for the
deglaciation of the former lake basin. To the SW of the Pico de Europa, in the Sanabria
National Park, ice from one of the largest ice caps in Iberia (Cowton et al., 2009) had
retreated forming the glacial lake of Lago de Sanabria by 25.6 cal. ka BP (Rodríguez-
Rodríguez et al. 2011). Whilst the timing of the maximum glacial extent is uncertain, since
this is a minimum age for deglaciation, it is consistent with findings from the nearby Picos de
Europa for a local glacier maximum that preceded the global LGM. There seems to be no
doubt that glaciers in this region were at their largest sometime prior to the global LGM and
MIS 2. However, because the geochronologies are based on radiocarbon ages from lake
basins, only minimum ages for deglaciation are known and at some sites several metres of
sediment underlie the lowermost radiocarbon dated layer. Determining the precise timing of
the local glacial maximum in Iberia will depend on obtaining exposure ages from stable
surfaces on these old moraines, with the best prospects on resistant quartz-rich lithologies.
3.4.7 Corsica
Kuhlemann et al. (2008) presented convincing evidence indicating that a major glacial
advance occurred close in time to the global LGM. There is no geochronological evidence of
an earlier glacier advance belonging to the last glacial cycle. Older more extensive glacial
deposits do exist but these are degraded and ascribed to the pre-Würmian (Middle
Pleistocene) cold stages (Conchon, 1986; Kuhlemann et al., 2005), although no
geochronology exists for these deposits.
3.4.8 Italian Apennines
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In the Gran Sasso massif, Central Apennines, Giraudi and Frezzotti (1997) dated the basal
sediments of a core taken through lake sediments dammed at the side of a valley by lateral
moraines associated with the maximum glacier extent. The age of 22.68 0.63 14C ka BP
(27.25 +/- 1.18 cal. ka BP) means that the lateral moraines impounding the lake formed
before this time, a few kiloyears before the global LGM. Lateral moraines cuts alluvial fans
and a radiocarbon age from beneath these fans of 27 14C ka BP (c. 31 cal. ka BP) means that
the glacier advanced after this time. Thus the largest ice advance in this area occurred
between 32-27 cal. ka. BP.
Giraudi (2012) presented glacial and lacustrine evidence from the Campo Felice basin
in the central Apennines which enabled the glacier events of the past 40 ka to be dated more
precisely. He found that the maximum extent of glaciation did not occur during the global
LGM. Instead it occurred earlier, between 33 and 27 cal. ka BP during a period of less
extreme climate with greater moisture availability.
3.4.9 Romania and the Balkans
The mountains of the Balkans were glaciated on several occasions during the Pleistocene and
some small glaciers survive today in areas such as Montenegro and Albania (Hughes, 2007,
2008, 2009). However, whilst Middle Pleistocene glaciations have been dated in Montenegro
and Greece, little is known regarding the timing of glaciations during the last cold stage. This
largely stems from the fact that many of the massifs are formed in limestone and U-series
techniques have been applied to date secondary carbonate cements within moraines. This
approach does not provide precise ages for moraine formation – only the calcite cements
themselves, which form sometime afterwards. Cosmogenic 36Cl has not been applied to date
glacial surfaces in carbonate terrains in this region. Limestone surfaces are prone to elevated
solution weathering and dating old pre-last glacial cycle surfaces using this technique could
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be problematic. However, 36Cl analyses could potentially be useful for dating Late
Pleistocene (limestone)) surfaces as has been shown in the Alps (e.g. Ivy-Ochs et al., 2009).
However, in mountain areas formed in quartz-rich rocks in Romania, Bulgaria and Kosovo,
10Be analyses have provided constraints on the timing of Late Pleistocene glaciations.
In the Retezat Mountains, Romania, Reuther et al. (2007) found no evidence for an advance
correlating with the global LGM. Moraines of the largest Late Würmian (MIS 2) glaciers
were abandoned at c. 16.1 ka (based on seventeen 10Be ages), several thousands of years after
the global LGM. A larger older advance is undated but Reuther et al. (2007) considered this
to be much older based on soil weathering characteristics although they suggest not during
the penultimate glaciation.
In the Šar Mountains, on the Kosovo-Macedonia border, Kuhlemann et al. (2009)
argued that the most extensive moraines in this area correlate with the global LGM. Their
ages from moraines in the Šar Mountains yielded 10Be ages as old as 19.4 ± 3.2 ka when
corrected for erosion, although when uncorrected this oldest age was 16.7 ± 2.3. However,
the small sample size from the lowest moraines (n = 4) and the large range of ages in this
sample (19.4 ± 3.2 ka, 16.1 ± 2.3, 14.7 ± 2.1, 12.4 ± 1.7 corrected for erosion at 10 mm/ka;
16.7 ± 2.3, 14.2 ± 1.7, 13.1 ± 1.6, 11.3 ± 1.4 zero erosion) mean that the precise age of the
moraines is unclear and given the ages a post-global LGM age seems more likely. The
uncorrected ages are similar to the minimum limiting ages for Late Pleistocene moraines in
Montenegro (17.5 ± 0.4, 13.9 ± 0.4, 13.4 ± 0.4, to12.5 ± 0.4 ka) obtained using U-series
(Hughes et al., 2011) and the moraine surfaces dated by Kuhlemann et al. (2009) may
represent recessional moraines.
In the Rila Mountains, Bulgaria, Kuhlemann et al. (2013) again dated moraines using
10Be and yielded 10 ages ranging from 14.4 to 23.5 ka (mean = 17.4 ka; st dev = 2.5 ka).
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They argued that two glacier advances can be recognised at one site Belin Iskar, where the
outermost moraine yielded an exposure age of 23.5 ± 3.6 ka whilst two of the inner ridges
yielded younger ages (14.3 ± 2.0 and 18.1 ± 2.5 ka).
Elsewhere in the Balkans, the timing of the maximum glacial advance in the last
glacial cycle is poorly constrained and in some places is inferred indirectly from ages
obtained from fluvial deposits. In Montenegro, U-series ages constrain the retreat of valley
glaciers from their maximum positions by 17 ka in the Orjen massif (Hughes et al., 2010) and
by 13 ka in the Durmitor Massif (Hughes et al., 2011). However, these are minimum ages and
ice may have retreated long before this. The problem of interpreting minimum ages is helped
in Greece by the dating of fluvial deposits down-valley of glaciated headwaters. Here, a
combination of electron spin resonance and U-series dating indicates several major phases of
aggradation through the last glacial stage (cf. Fig. 7 of Woodward et al., 2008). Hughes et al.
(2003, 2006) concluded that glaciers retreated after c. 25 ka with debris supply exceeding
snow accumulation resulting in the formation of rock glaciers around the time of the global
LGM. Based on glacier-climate modelling and comparisons with the pollen record at nearby
lake Ioannina, Hughes et al. (2006) argued that the period between 30-25 ka was the most
likely interval for the maximum extent of glaciers in the Pindus Mountains. However, no
direct ages have been obtained from glacial landforms in this area to test this hypothesis and
the maximum glacier advances may correlate with older fluvial aggradational units found in
the Voidomatis basin (cf. Hughes and Woodward, 2008; Woodward et al., 2008).
3.4.10 Turkey
In the Kaçkar Mountains in northeastern Turkey, exposure ages from boulders associated
with the most extensive glacier advance in the Verçenik Valley range from 24.0 ± 1.1 to 20.3
± 0.2 ka (26.0 ± 1.2 to 21.9 ± 1.3 ka, when corrected for snow and erosion at 3 ± 0.5mm kyr-
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1). Rapid retreat occurred after 17.5 ± 0.9 (18.8 ± 1.0 ka) and by 14.8 ± 0.6 ka (15.5 ± 0.7 ka)
the main valley became ice-free with glaciers restricted to tributary cirques. (Akçar et al.,
2007). The geochronology from the Verçenik Valley is similar to that found in the
neighbouring Kavron valley (Akçar et al., 2007b).
In western Turkey, in the north-facing valleys of Mount Sandıras (2295 m) 36Cl
exposure ages from moraine boulders show that the last local glacier maximum occurred at
approximately 20.4 ± 1.3 ka (Sarıkaya et al., 2008). Glacier retreat was interrupted by
readvances at c. 19.6 ± 1.6 and 16.2 ± 0.5 ka. In the Degegöl Mountains, in SW Turkey, 10Be
and 26Al exposure ages indicate a glacier maximum in the Muslu Valley at c. 24.3 ± 1.8 ka
(Zahno et al., 2009). This glacial maximum was slightly earlier than elsewhere in Turkey and
highlights regional variability in the timings of glacier maxima. This regional variability is
further emphasised by the glacial record from the Aladaglar Mountains of south-central
Turkey. Here, Zreda et al. (2011) found that 36Cl analyses from a series of seven successive
moraines, including the lowest, gave apparent exposure ages ranging from 10.2 ± 0.2 ka to
8.6 ± 0.3 ka. Zreda et al. (2011) note that these ages may be too young, depending on
production rate estimates, but nevertheless highlight the pace of glacier retreat in this area.
3.5 Asia
3.5.1 Barents-Kara Ice Sheet
In northern Eurasia the Barents-Kara Ice Sheet reached its maximum recorded extent in the
penultimate glacial cycle. During the last glacial cycle, the maximum extent of ice was
reached during the Early Weichselian (Early Valdaian) between 100 and 90 ka (Fig. 7). This
was followed by successively smaller Middle and Late Weichselian (Middle and Late
Valdaian) glaciations at c. 70-65 ka, 55-45 ka and 25-15 ka, respectively (Svendsen et al.,
2004; Larsen et al., 2007; Vorren et al., 2011), the latter glaciation corresponding with the
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global LGM (Fig. 7). Expansions of Eurasian ice sheets observed in the geological record at
c. 90, 60 and 20 ka are consistent with predictions of glaciological model simulations forced
by global sea level and solar changes (Svendsen et al., 2004). The successively smaller
Barents-Kara Ice Sheet contrasts with successively larger phases of the Fennoscandinavian
Ice Sheet, which as noted in the previous section, reached its maximum close to the global
LGM. This situation is linked to increasing aridity over the Barents-Kara region caused by
the build-up of the more westerly Fennoscandinavian Ice Sheet, a situation again confirmed
by model simulations (Svendsen et al., 2004). The dominance of the Fennoscandinavian Ice
Sheet is highlighted by calculated changes in ice sheet volume in northern Eurasia, with the
Late Weichselian (MIS 2 = 10 x 1000 km3) glaciation being 1.67 times the volume of the
Early Weichselian (=MIS 4/3 = 6 x 1000 km3) glaciations (Svendsen et al., 2004, their Fig.
17).
3.5.2 NE Russia
As in the Barents-Kara region, the role of moisture supply on glaciation was the major
influence on glaciations in Siberia. This region experienced limited glaciation during the
global LGM despite very low temperatures because they were countered by an extremely arid
climate (Hubberton et al., 2004; Stauch and Gualtieri, 2008). In the Verkhoyansk Mountains
no evidence of glaciation is recorded correlating with the global LGM. Aeolian dust covering
moraines has been dated using OSL the ages demonstrating that the last Late Pleistocene
glaciations in the central Verkhoyansk Mountains occurred before 50 ka (Stauch et al. 2007;
Stauch and Lehmkuhl, 2011). Similar conclusions have also been reached for the northern
parts of these mountains (Schirrmeister et al., 2002; Hubberten et al., 2004). A further series
of larger glaciations have been dated using Infrared stimulated luminescence dating (IRSL) to
80-90 ka and 100-120 ka (Stauch et al. 2007). This sequence of successively smaller
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glaciations (Fig. 8) bears similarities to the Barents-Kara Ice Sheet. The largest recorded
glaciations in the Verkhoyansk Mountains, occurred during the penultimate glacial cycle
(Saalian).
Unlike in the Verkhoyansk Mountains, glacier advances are evident at the global
LGM in other mountain areas of NE Siberia. Stauch and Gualtieri (2008) suggest spatial
variability in the timing of glaciation in NE Russia is attributable to contrasting sources of
precipitation with Atlantic areas dominating in the Verkhoyansk Mountains whilst Pacific
sources dominate in the Anadyr Mountains, Koryak Mountains and on Kamchatka. In the
Pekulney Mountains (Anadyr Mountains) glaciers were present but were limited in extent
with cosmogenic (36Cl) ages from moraine boulders ranging from 16.2 to 23.6 ka (n = 3,
excluding a young outlier). On stratigraphically older glacial surfaces (erratics and bedrock),
cosmogenic (36Cl) ages ranged from 36.6 to 69.6 ka (n = 7). Brigham-Grette et al. (2003) use
this evidence coupled with limiting radiocarbon ages (> 40 ka) from related river terraces to
argue that the glaciers in these mountains reached their maximum earlier in the last glacial
cycle (Brigham-Grette et al., 2003). Similar findings were reported from the Koryak
Mountains (Gualtieri et al., 2000) and other mountains in this region (Stauch and Gualtieri,
2008). On Wrangel Island, in the Arctic Ocean, 10Be analyses indicate bedrock exposure
since 84.4 ka (Gualtieri et al., 2005), with no evidence of ice cover at the global LGM.
Brigham-Grette (2001) and Brigham-Grette et al. (2001) suggested that glaciers advanced
during or at the end of the Last Interglacial in NE Siberia. This would suggest that the glacial
history of this area is out-of-phase with the record of global ice volume.
In southern Siberia the glacial history is very similar to that found in the mountains of NE
Siberia and the Barents-Kara Ice Sheet. In the central Sayan-Tuva Upland (c. 52ºN, 98ºE),
glacial advances occurred in MIS 4, MIS 3 and MIS 2 (Arzhannikov et al., 2012). As with the
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other areas of Siberia described above, the largest glaciations occurred during MIS 6 and
Arzhannikov et al. (2012) suggest that ice would have survived throughout the last
interglacial in these mountains.
3.5.3 Pamirs, Tien Shan, Altai
The Pamir, Tien Shan and Altai are situated to the north of the Himalaya in central Asia. In
the southeastern Pamir moraines associated with the largest glaciers of the last glacial cycle
have been dated to MIS 4 using 10Be analyses. Advances are also recorded in at the MIS
3/MIS 2 boundary , as well as during a time interval in MIS 2 close to the global LGM
(Owen et al., 2012). Similar findings were reported by Zech et al. (2005b) who suggested that
the largest advance of the last glacial cycle occurred during MIS 4 (the 'local last glacial
maximum'). Seven 10Be exposure ages ranging from 52.6 ± 6.8 to 61.2 ± 8.0 ka provide ages
for the timing of ice retreat at the MIS 4/3 boundary. In the Turkestan range of the Pamir-
Alay, Kyrgyz Republic, Narama and Okuno (2006) found that radiocarbon ages for several
soil layers buried in a lateral moraine of the Asan-Usin glacier suggest that glaciers shrank or
stagnated during mid-MIS 3. These authors argued that glaciers were larger during MIS 2
with no evidence of an earlier glacier maximum. This differs from the findings of Zech et al.
(2005b) and Owen et al. (2012) in the Pamirs as well as other mountain areas of central Asia
(see below). Elsewhere in the Kyrgyz Republic, in the Terskey-Aloo Range, Narama et al.
(2007) dated glacial and loess deposits and found that a major glacier expansion (Terskey II)
also occurred here during MIS 2, between 21 and 29 ka. However, in this area a larger glacier
advance (Terskey I) occurred earlier, possibly during MIS 4.
In the Kyrgyz Tien Shan, the timing of glacier advances has been constrained by
Koppes et al. (2008) who applied 10Be analyses to date moraine surfaces. Glaciers in the north
and east of the Kyrgyz Tien Shan last advanced to their maximum positions during MIS 5
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and MIS 4, while in the south and west of the range the maximum advance occurred during
MIS 3 (Fig. 9). There is no evidence of a major glacial advance during MIS 2. Koppes et al.
(2008) dated the maximum extent of glaciers during the last glacial cycle to c. 49 ka in the
northern part of the range and c. 35 ka in the southern part, based on 10Be analyses from
moraine boulders. However, Sanhueza-Pino et al. (2011) argue that the maximum glacier
advance must predate 63 ka based on 10Be ages from landslide surfaces in glaciated valleys.
Either way, it is clear that the largest glaciers of the last glacial cycle occurred well before the
global LGM. This has recently been reinforced by 10Be exposure ages from moraines reported
in Zech (2012) who suggested that this most likely reflects increasingly arid conditions in
Central Asia during the last glacial cycle.
In the Chinese Altai, OSL dates suggest that the glacier maximum occurred during
MIS 3 whilst in the Russian Altai glaciers advanced during MIS 4 and 2 (Lehmkuhl et al.,
2011). In the Darhad Basin of northern Mongolia, glaciers reached their largest extent before
the global LGM. Gillespie et al. (2008) found that glaciers were larger during MIS 3 than
MIS 2. Whilst there is some uncertainty in the dating, Gillespie et al. (2008) provisionally
accept 35–53 ka as the age for the MIS-3 advance(s). Later MIS 2 glaciers were only slightly
smaller with ELAs that were only c. 75 m lower than during the MIS 3 advance.
3.5.4 Himalaya and Tibet
In the Himalaya, Benn & Owen (1998) argued that glacier oscillations reflect periods of
positive mass-balance coincident with times of increased insolation and intensified monsoon
activity. Westerly atmospheric systems are also important (Koppes et al., 2008) and Dortch
et al. (2013) provide a synthesis of the relative effects of monsoons and westerlies on glacier
advances in the Himalaya.
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The most extensive Late Pleistocene glaciation occurred during the early part of the
last glacial cycle (Taylor and Mitchell, 2000; Heidrick et al., 2011; Dortch et al., 2009,
2010b, 2013) and also likely during MIS 3, a time of high insolation (Richards et al., 2000a;
Owen et al., 2002a; Kamp and Owen, 2011; Owen, 2011). In the Karakoram, the glacier
expansions in MIS 3-4 were on a larger scale than those of MIS 2 (Phillips et al., 2000;
Richards et al., 2000a; Owen et al., 2002a, b). In the Everest area of the central Himalaya,
glaciation was also most extensive during the early part of the last glacial and more restricted
during MIS 2 (Finkel et al., 2003). This situation is matched in many other areas of the
Himalaya.
In the Indian Himalaya, at the global LGM many glaciers only advanced to within 10
km of modern glacier fronts (Owen, 2011). Much of this argument is supported by
cosmogenic exposure and luminescence ages from numerous sites in Pakistan and India (cf.
Owen et al., 2008; Kamp and Owen 2011 and Owen 2011, and references therein). For
example, in the Tangtse Valley in Ladakh, roche moutonées yield mean exposure ages of
38.5 ± 2.8, 34.9 ± 2.5, and 34.3 ± 2.6 ka suggesting that the glaciers that retreated during MIS
3 were not overrun by later glaciers (Dortch, et al., 2011). Forty-seven 10Be ages have also
been obtained from moraines the Ladakh and Pangong Ranges in NW India. In the Ladakh
Range, the most extensive glaciation occurred at c. 81 ± 20 ka with smaller glacier advances
following this including one later undated advance followed by a smaller advance at 22±3 ka.
However, there is variability within this region; for example; in the Nubra River valley area
of northernmost Ladakh glacier advances are only recorded at 81 and 45 ka during the last
cold stage (Dortch et al., 2010b). Similarly, in the Pangong Range, Dortch et al. (2013) report
that two glaciations are recorded by moraines with exposure ages of 85±15 ka and 40±3 ka
with no moraines present from MIS 2.
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In Tibet, there is also evidence that glaciers were restricted in size during the global
LGM with larger glaciers earlier in the last glacial cycle (Lehmkuhl and Owen, 2005). For
example, Heyman et al. (2012) showed that glaciation in the Bayan Har range has been
restricted in size for the past 40 ka. The largest glaciers for the last glacial cycle formed
before 60-100 ka with a later smaller phase of glaciation before 40-65 ka. The global LGM is
not represented by moraines in this mountain range. During the global LGM advances of
Tibetan glaciers were much smaller than elsewhere because climate was very arid in this
region accompanied by high sublimation rates (Schäfer et al., 2002). Fig. 10 shows the
composite probability plots from cosmogenic exposure data from throughout the Himalaya
and Tibet. The peaks in the graph illustrate that moraine surfaces older than the LGM are
present in MIS 3, 4 and 5. MIS 2 displays the biggest probability peak simply because of a
greater number of samples and lower error on these samples. The fact that probability peaks
precede this period illustrate the presence of older, more extensive, glacial surfaces within the
last glacial stage. Most recently, Dortch et al. (2013) have also complied a review of over 700
cosmogenic ages from the wider western Himalaya-Tibetan orogeny and found evidence for
six moraine clusters representing advances of diminishing size from 80-20 ka (at 80 ± 5, 72 ±
8, 61 ± 5, 46 ± 4, 30 ± 3 and 20 ± 2 ka). Of these, the advances that are recorded over the
widest areas of the western Himalyan-Tibetan orogen were during MIS 4 and MIS 2, with
the former being the more extensive than the latter. Given the absence of MIS 2 moraines in
some valleys, as noted above, but presence in many others highlights the regional complexity
of the glacier record in the western Himalaya-Tibetan orogen.
3.5.5 Japan and Taiwan
The maximum extent of glaciation in Japan and Taiwan during the last glacial cycle occurred
during MIS 4 and these glaciers persisted between MIS 5d and MIS 3 (Ono et al., 2005;
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Sawagaki and Aoki, 2011). In Japan, this is based on the evidence of marker tephras on the
surface of moraines. Later glaciers during MIS 2 were more limited in extent. Aoki et al.
(2000) dated moraines associated with these smaller glaciers using 10Be analyses and yielded
exposure ages of 17-25 ka.
In Taiwan, in the Hseuh Shan, sediment samples from a lateral moraine have yielded
OSL ages of 51 ± 10 and 56 ± 4 ka (Hebenstreit and Böse, 2003) and a thermoluminescence
(TL) age of 44.25 ± 3.27 ka (Cui et al., 2002) . Younger glacial landforms have been dated
using 10Be and have yielded Weichselian Late-glacial and Holocene ages (Böse and
Hebenstreit, 2011). There is no geochronological evidence of an MIS 2 glaciation.
3.6 Africa
3.6.1 Morocco
In the High Atlas, Morocco, three phases of glacial advance have been recognised: the oldest,
and most extensive moraines yielding a 10Be exposure age of 76.0 ± 9.4 ka; a second
generation of moraines yielded an exposure age of 24.4 ± 3.0 ka, and moraines belonging to a
third advance has yielded three exposure ages of 11.1 ± 1.4, 12.2 ± 1.5 and 12.4 ± 1.6 ka
(Hughes et al. 2011a). A larger dataset indicates that oldest advance is clearly pre-LGM with
10Be exposure ages ranging from of 31.1 ± 3.8 to 76.0 ± 9.4 ka (Hughes et al., 2011b) and
may indicate a glacier maximum in Morocco early in the last glacial cycle (Soltanian Stage).
However, more ages are required to be statistically confident of the true age of these older
moraines. A second generation of moraines recorded in the High Atlas do yield exposure
ages consistent with advance close to the global LGM (24-15 ka) and work is under way to
establish the timing and extent of glaciations in Morocco (Hughes et al. 2011).
3.6.2 Ethiopia
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In the Bale Mountains, Ethiopia, the timing of the last deglaciation has been constrained
using radiocarbon dates from glacial lake sediments. Ice progressively retreated after 17 cal.
ka BP. This illustrates that ice occupied the basin at Lake Garba Guracha (at c. 3920 m a.s.l.)
at the LGM (Tiercelin et al., 2008). However, glaciers extended more than 5 km down-valley
beyond this lake at their maximum with moraines present as low as c. 3400 m a.s.l.
(Osmaston et al. 2005) and the precise timing of the local glacial maximum in this area
remains unknown.
3.6.3 Equatorial Africa
The timing of the largest glaciations of the last glacial cycle in equatorial Africa is subject to
debate. On Mount Kenya, the most extensive glaciation of the last glacial cycle (Liki I
Glaciation) was coeval with MIS 4 according to Mahaney (2011). A radiocarbon age of 33.8
± 2.8 14C ka BP (38.8 ± 2.6 cal. ka BP) from a lake basin provides a minimum age for the
Liki I moraines (Mahaney, 2011). However, this interpretation is contradicted by Shanahan
and Zreda (2000) who argue for a Middle Pleistocene age (MIS 10 or 12) based on three
cosmogenic exposure ages (36Cl). In response, Mahaney (2011) considers that these ages may
be anomalously old inherited ages. A later glacial advance (Liki II) is possibly associated
with the global LGM and radiocarbon ages of 14-15 14C ka BP (18.5-17 cal. ka BP)provide a
constraint on the age of deglaciation (Mahaney 2011). Shanahan and Zreda (2000) report an
average 36Cl age of 28 ± 3 ka for the Liki II moraines in the Gorges Valley. However,
moraines that are apparently considered stratigraphically-equivalent in the Teleki Valley
yield a wide range of 36Cl ages with a mean of 64 ± 40 ka (Shanahan and Zreda, 2000).
Thus, it is still uncertain based on the information available whether or not the most extensive
glacier advance of the last cold stage on Mount Kenya occurred at or close to the global
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LGM. However, 36Cl ages from moraines on Kilimanjaro do cluster close to the LGM with an
average exposure age of 20 ± 1 ka for the moraines of Main Glaciation, on Mawenzi
(Shanahan and Zreda, 2000, their Table 2). The moraines of the larger Oldest Glaciation on
Kilimanjaro have yielded ages much older than the last glacial cycle. A series of more limited
glacial advances on both Mount Kenya and Kilimanjaro date from the Late-glacial and
possibly the Younger Dryas Stadial (Shanahan and Zreda, 2000).
3.6.4 Southern Africa
The Drakensberg Mountains in southern Africa supported small niche glaciers during the
Late Pleistocene (Mills et al. 2009). Radiocarbon ages obtained from basal organic sediment
found within moraines indicate that they are of LGM age. Five calibrated ages range from
20.53 to 13.82 cal. ka BP. However, Mills et al. (2009) note that these can represent
minimum or maximum ages. There is no evidence of older and larger glaciations.
3.7 Australasia
3.7.1 Papua New Guinea
On Mount Giluwe in Papua New Guinea, Barrows et al. (2011) dated moraine surfaces using
10Be. They found that the most extensive moraines yielded exposure ages of 293-306 ka
(Gogon Glaciation), 136-158 ka (Mengane Glaciation), 62 ka (Komia Glaciation) and from
>20.3-11.5 ka (Tongo Glaciation). The Komia and Tongo glaciations were similar in size
although evidence of the former is preserved because glaciers were slightly larger in some
places. On Mount Trikora 13 boulders from three successive moraines were dated using 10Be
and 26Al yielding ages of 14ka to 27ka, spanning the global LGM (Fink et al., 2003) . The
fact that these moraines cross-cut older (but undated), large moraine systems, led Barrows et
al. (2011, p.1032) to speculate that glacier were larger prior to the LGM. However, in the
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Trikora region the ages of these older glaciations are unknown and whether these relate to
advances during the last glacial cycle (Komia glaciation) or earlier (Gogon or Megame
glaciations) is unclear.
3.7.2 Australia
In Australia, glaciation was limited to the southeastern part of the continent, in the Snowy
Mountains and Tasmania. In the Snowy Mountains, there is no clear evidence for glaciations
pre-dating the Late Pleistocene (Colhoun and Barrows, 2011). Late Pleistocene glaciers
formed short valley tongues and the total glaciated area is only ca. 12-15 km2 with the largest
glacier reaching a length of ~1.7 km. The oldest moraines have yielded 10Be exposure ages of
39.9 ± 4.0, 45.3 ± 4.7 and 59.3 ± 5.4 ka. Barrows et al. (2001) argue that this represents an ice
advance (Snowy River advance) during MIS 4. Later glacier advances occurred during MIS 3
(Hedley Tarn advance and at 19.1 ± 1.6 ka (Blue Lake advance) in MIS 2) , and after a brief
still stand advanced again at 16.8 ± 1.4 ka (the Twynam advance) following which the Snowy
Mountains were completely deglaciated (Barrows et al., 2001; Colhoun and Barrows,
2011).). The Blue Lake advance is considered to coincide with the global LGM and 10Be ages
from these moraines recalculated and presented in Barrows et al. (2002) are 17.3 ± 1.6, 18.0
± 1.6 and 19.3 ±1.8. Similar ages have been obtained from moraines at Lake Cootapatamba
(Barrows et al., 2001, 2002). However, unlike at Blue Lake, there is no evidence for earlier
glacier advances.
Glaciations were more extensive in Tasmania which presents numerous cirque, valley
and ice cap glaciations during the Pleistocene. The most extensive glaciations occurred
during the Early and Middle Pleistocene, the earliest dating from c. 1 million years ago
(Fitzsimmons et al., 1990; Augustinius, 1999; Colhoun et al., 2010). During the Late
Pleistocene ice covered an area of 1085 km2 at its maximum extent. There is evidence for
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multiple advances with many moraines correlating closely in time with the global LGM
(Barrows et al., 2002, Fink et al., 2004). However, as in the Snowy Mountains, a number of
erratics on slightly more extensive down-valley moraines appear to give exposure ages which
predate the global LGM by up to 20 ka, although this is based on a limited dataset (Fink et
al., 2004) . Despite the limited number of cosmogenic ages, the largest Late Pleistocene
glacial advance in Tasmania may be correlated to MIS 3, unlike at MIS 4 in the Snowy
Mountains. The best evidence for a MIS 3 advance in Tasmania is reported by Fink et al.
(2004), at Mt Jukes for 4 boulders from a moraine at 38.4 ± 5.9 ka, ~150 m lower in
elevation than a moraine with two exposure ages of 21.3 and 14.9 ka. Further evidence is
found at Mount Field where moraines have yield 36Cl ages on 2 samples of 44.1 ± 2.2 and
41.0 ± 2.0 ka (Mackintosh et al., 2006) and in the Hartz Mountains cirque moraines have
yielded 36Cl ages of 39.3 ± 2.9 and 44.5 ka ± 2.9 ka (Barrows et al., 2002). These ‘MIS 3’
glaciers extended 4 km beyond the later MIS 2 glacier limits whose moraines have yielded
36Cl ages of 18.1 ± 1.0 ka and 18.8 ± 1.3 ka (Mackintosh et al., 2006). At other sites, boulders
from moraine ridges have yielded 10Be and 36Cl ages between 53 and 24 ka (Barrows et al.,
2011) and may correlate with the MIS 3 moraines at Mount Field. In addition, a combination
of geomorphology and radiocarbon dates from the King Valley pointed to an advance close to
the global LGM and an earlier advance >48.7 14C ka BP (represents an infinite limiting age
when calibrated) which extended to Chamouni (Fitzsimmons et al., 1992), although this is
clearly a limiting age and the age of the actual advance may be much older. According to
Colhoun and Barrows (2011), the current evidence suggests glaciations occurred during MIS
2, 3 and possibly 4, and most evidence supports a weak local LGM advance. What is apparent
in both mainland Australia and Tasmania is that glaciers correlating with the global LGM
were preceded by larger glaciers earlier in the last glacial cycle.
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3.7.3 New Zealand
Glaciers reached close to their most extensive positions of the last glacial cycle during MIS 2
(Barrell, 2011). However, the build-up of ice began in MIS 3 before 27 ka (Suggate and
Almond, 2005) and there is now convincing evidence of extensive glaciation in MIS 3 with
significant yet smaller glacier advances during MIS 2. Putnam et al. (2013) presented the
results of 73 10Be ages from moraines at Lake Ohau on South Island, New Zealand. Based on
this data they identified successively smaller glacier advances at 138.6 ± 10.6 ka, 32.52 ±
0.97 ka, 27.4 ± 1.3 ka, 22.51 ± 0.66 ka and 18.22 ± 0.5 ka. The MIS 2 glaciers were therefore
smaller than the MIS 3 advance. Schäfer et al. (2006) argued that the MIS 2 glaciers were in
synchronisation with other northern hemisphere mid-latitude mountain glaciers based on a
mean exposure age of 17.4 ± 1.0 ka for moraines associated with the local last glacier
maximum in the Lake Pukaki area in the Southern Alps of New Zealand,. Shulmeister et al.
(2010a) recalculated ages presented in Schäfer et al. (2006) using a common set of scaling
and production rates resulting in ages older by ~ 0.9 ka. Schulmeister et al. (2010) argue that
exposure ages from the neighbouring Rakaia Valley moraines clearly indicate that ice
recession began at 24 ka and the local glacier maximum occurred at 24-26 ka (Shulmeister et
al., 2010a). One of the most comprehensive deglacial chronologies in New Zealand has been
established for the Cobb Valley, in NW Nelson on South Island. Here, Shulmeister et al.
(2005) presented the result of twenty-one 10Be ages which showed that in this valley, the last
deglaciation commenced no earlier than 18-19 ka, and was followed by numerous short-term
still-stands and/or minor re-advances before complete glacier retreat by 14 ka. Similarly, in
the Tararua Range on New Zealand’s North Island, six 10Be ages ranging from 17.1 ± 1.1 ka
and 23.9 ± 1.7 ka from moraine boulders and bedrock indicate that glaciers retreated from
maximum positions soon after the global LGM (Brook et al., 2008).
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Whilst there is clear evidence of deglaciation in New Zealand after a major ice
advance during MIS 2, there is evidence that in some areas the largest glaciation of the last
glacial cycle occurred earlier, not just in MIS 3 (Putnam et al., 2013) but also during MIS 4
and earlier (Fink et al 2006 and other papers, see below). In the Cascade Plateau area of SW
New Zealand, Sutherland et al. (2007) used 10Be analyses to date the exposure ages of
moraines. They also found that the most extensive glaciation of the last glacial cycle occurred
early, at c. 79 ka, during MIS 5a. Later glacier advances also occurred at c. 58 ka and 22–19
ka. In NW Nelson, South Island New Zealand, 10Be exposure ages suggest that glaciers
expanded to the most extensive positions of the last glacial cycle during MIS 4/3 (Thackray
2009). Similarly, McCarthy et al. (2008) used cosmogenic and luminescence data to argue
that MIS 4 cirque glaciers were similar in size or slightly more extensive than MIS 2 glaciers.
Moreover, Barrell (2011, his Fig. 75.2) identified several areas where MIS 4 glaciers were
more extensive than during MIS 2.
3.8 Antarctica
Antarctic ice extended out to the continental shelf during its maximum extent in the last
glacial cycle (Fig. 2). However, there is intense debate regarding the timing of the maximum
extent of the various ice sheets (see Stolldorf et al., 2012 for review). There is strong
evidence to suggest that the Antarctic Peninsula Ice Sheet expanded during the LGM and
retreated subsequently (Ó Cofaigh et al., 2002; Domack et al., 2005; Evans et al., 2005;
Heroy and Anderson, 2005, 2007; Kilfeather et al., 2011; Davies et al., 2012). The same is
true for the West Antarctic Ice Sheet in the records from the Amundsen Sea (Lowe and
Anderson, 2002; Evans et al., 2006; Graham et al., 2009, 2010; Jakobsson et al., 2011, 2012;
Smith et al., 2011; Bromley et al, 2012; Kirshner et al., 2012) and the Bellingshausen Sea
(Hillenbrand et al., 2010). However, in the eastern Ross Sea, there is evidence that the West
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Antarctic Ice Sheet had retreated from the continental shelf by the time of the LGM (Mosola
and Anderson, 2006; Bart and Cone, 2011). The evidence from the East Antarctic Ice Sheet is
even more complicated, with some papers suggesting ice sheet advance up to the LGM
(Goodwin and Zweck, 2000; White et al., 2011), whilst others suggest no LGM ice expansion
occurred. For example, Gore et al. (2001) found that OSL ages from glaciofluvial and glacial
lake sediments in the Bunger Hills area of East Antarctica appeared to show that deglaciation
occurred as early as 30 ka and that the area was ice-free at the LGM. In fact, based on a 15
erratics on mountain flanks of the Grove Mountain’s situated 800 kilometres inland from the
coast, Lilly et al. (2010) found that LGM ice thickness of the East ant ice sheet was possibly
smaller than it is today. Similarly, Mackintosh et al. (2007) found that exposure ages from
mountain tops in Mac. Robertson Land, East Antarctica, revealed that there had little change
in ice sheet thickness since the LGM, Glaciomarine sediments in the Weddell Sea have
yielded radiocarbon ages that span the past 30.5 cal. ka BP, indicating that the East Antarctic
Ice Sheet in this area retreated from its maximum position before the global LGM (Stolldorf
et al., 2012). Thus, it appears that expansion and retreat of the East Antarctic Ice Sheet by far
the largest global ice mass, was neither in phase with that of the smaller West Antarctic nor
the Antarctic Peninsula Ice Sheets during the last glacial cycle.
4. Discussion
It is clear from this review of glacial records from around the world that glaciers reached their
maximum extent at different times during the last glacial cycle (Fig. 11). Whilst this finding
is not new (cf. Gillespie and Molnar, 1995) nor unexpected, the degree of complexity and
variability both spatially and temporally is striking, with many areas seeing glaciers reaching
their maximum extent early in the last glacial cycle, well before the global LGM in MIS 2.
Furthermore, this is not only a phenomenon that is restricted to mid-latitude mountain
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glaciers but it is also seen in certain sectors of the large continental ice sheets. The Antarctic
polar ice sheets demonstrate their own degree of complexity.
The expected dominance of the Laurentide Ice Sheet as a principal driver of the
changes reflected in ocean δ18O and thus the marine isotope record is supported by the
geochronological evidence from this region. Here, most sectors of the former ice sheet appear
to have reached their maximum extents close to the global LGM during MIS 2. The
southeastern sector of the Fennoscandinavian Ice Sheet and the European Alpine Ice Sheet
also reached their maximum extents close in time to the maximum period of global ice
volume indicated in the marine isotope records. Taken together, this accords with the
interpretations of marine oxygen isotope records as reflecting variations in global ice volume,
supporting the hypothesis proposed by Shackleton (1967). Other areas, where the maximum
extent of glaciation during the last glacial cycle accords with the global LGM (or at least MIS
2), include parts of the British-Irish Ice Sheet, the European Alps, some mountain areas in
southern Europe and Turkey, the West Antarctic Ice Sheet, parts of Argentina, the tropical
Andes, the Sierra Nevada and Hawaii (USA), and the southern sector of the Cordilleran Ice
Sheet (USA). In an increasingly larger number of case studies there is also evidence of earlier
more extensive glacial advances – all ascribed to the last glacial cycle. Even in the areas
listed above as having maximum advances close to the LGM there are often sites within these
which contradict this and display evidence of an earlier glacier maximum.
There are many areas where a majority of records suggest glaciers reached their
maximum extent early in the last glacial cycle such as in MIS 3 or 4. Starting with the major
continental ice sheets, regions where the maximum extents were reached early in the last
glacial cycle include the southwestern sector of the Fennoscandinavian Ice Sheet ( MIS 4/3),
the southeastern sector of the British-Irish Ice Sheet, the north Eurasian (Barents-Kara) Ice
Sheet (MIS 5), the northern sector of the Cordilleran Ice Sheet (MIS 4), the various Alaskan
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ice caps (MIS 4) and the New Zealand Southern Alps (MIS 4) and Tasmania ( MIS 3). In
Asia, glaciers reached their maximum extent of the last glacial cycle well before the global
LGM in many areas (Fig. 9). As with the large northern Eurasian Ice Sheet, the glaciers in
Siberia, central Asia (Pamirs, Tien Shan, Altai), Tibet and the Himalaya all reached their
maximum extent in MIS 4. Similar findings have also emerged from the mountains of Japan
and Taiwan. In fact, for Asia, glacial maxima in MIS 2 are the exception rather than the rule.
In Tasmania, glaciers were more extensive, during MIS 3 than they were in MIS 2. In South
America, the situation is less clear. In the southern Andes there is some evidence that MIS 4
glaciers were larger than during MIS 2. However, in the tropical Andes the reverse appears to
be true, with the largest glaciers now correlated with MIS 2 and the global LGM. This is not
to say that major MIS 2 advances are not recorded in many of these areas (although in Siberia
and Tibet LGM advances are absent), only that earlier glacier advances were more extensive.
All of this new insight rests on interpretations of the rapidly expanding
geochronogical datasets in glaciated areas. There is a possibility that some of the moraines
that pre-date the global LGM may in fact be older than the last glacial cycle. Many of the
techniques used to date the moraines provide minimum limiting ages. Radiocarbon ages from
organic materials resting on till or moraine, or simply in rock basins within glacial limits,
cannot extend to the early part of the last glacial cycle and 14C ages are pegged at 50 ka at
most. Even when ages are within the range of the technique, they still only provide minimum
ages due to the delay in onset of the local ecosystem following deglaciation . Interpretations
of cosmogenic exposure ages are also impeded as surfaces become older, since there is often
an increasing degree of scatter in exposure ages from increasingly older surfaces (Putkonen
and Swanson, 2003). Boulder and bedrock surfaces also erode over time under subaerial
processes limiting the exposure age. Nevertheless, for typical granitic erosion rates of 1-3
mm/ka (e.g. Phillips et al., 2006; Rolfe et al., 2012) the increase in age would be no more
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than 5-10% at 50 ka using 2 mm/ka. The possibility of misleading ‘apparent’ exposure ages
can be refuted in some areas where more extensive pre-last glacial cycle (Middle Pleistocene)
glacial surfaces yield ages significantly older than those younger surfaces dating from the
early part of the last glacial cycle. Despite these problems, it cannot be coincidence that so
many glacial records yield geochronologies suggesting a maximum period of glacier
expansion early in the last glacial cycle. Even in the case of the Laurentide Ice Sheet, total
ice extent during MIS 4 was not much smaller (10-20% less) than in MIS 2 (Stokes et al.,
2012) (Fig. 3).
The asynchrony of even the large continental ice sheets poses questions regarding
their contributions to global ice volume, which based on the marine isotope and global sea
level records, point to a global glacial maximum in MIS 2. The apparent asynchrony in the
maximum extents of the West and East Antarctic Ice Sheets has important implications in this
regard. The East Antarctic Ice Sheet is currently the largest ice mass on Earth, yet this now
appears to have been significantly smaller at the global LGM and there is evidence that
sections of this ice sheet reached its maximum well before the LGM. Only the smaller West
Antarctic and Antarctic Peninsula Ice Sheets appeared to have reached a maximum at the
LGM. This indicates that the marine isotope record of global ice volume was dominated by
other ice masses. The Laurentide Ice Sheet, would have been far more significant Hughes,
2013) – largely because this ice mass was able to expand over a large area of previously
unglaciated continental land. Indeed, Clark and Pollard (1998) assumed the Laurentide Ice
Sheet has dominated the global δ18O signal during the past 2.8 Ma. In contrast, the Antarctic
Ice Sheets are only slightly smaller than during their glacial peak – a fact constrained by the
position of the continental shelf around this continent (Ehlers and Gibbard 2007, compare
their Figs. 7 and 9). A similar argument can be made for the Greenland Ice Sheet (Ehlers and
Gibbard 2007, compare their Figs. 8 and 10). Other Pleistocene ice sheets such as the
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Cordilleran, Fennoscandinavian, British-Irish Ice Sheets and the New Zealand Ice Cap were
also constrained by marine margins on their western flanks. This is important because it is
from the west that these ice masses received precipitation whereas the eastern margins were
simply constrained by lee-side aridity. This is particularly marked in the case of the Barents-
Kara Ice Sheet which grew successively smaller through the last glacial cycle due to its
position to the east and northeast of an increasingly dominant Fennoscandinavian Ice Sheet,
where the ice centre had shifted from over Norway in MIS 4 to over the Gulf of Bothnia in
MIS 2 enabling expansion of its SE sector (Mangerud et al., 2011).
Whilst there is some similarity in the timings of the pre-LGM early glaciations, the
maximum extents of advances are spread from MIS 5d to MIS 3 though no clear unified
glacial maximum is apparent from the early part of the last glacial cycle. This is unlike the
maximum extent of advances during MIS 2, which are closely matched across the world
within the first half of MIS 2 (26-18 ka), albeit with some minor temporal variability. It is
clear that even for the large continental ice sheets, ice build-up occurred early in the last
glacial cycle. The climatic mechanisms explaining asynchronous glaciations are likely to be
complex. The mechanisms are also likely to vary regionally.
In the northern hemisphere, the insolation in MIS 5d was the lowest of the last glacial
cycle (Fig. 5) and coincides with apparent maximum glacier expansions in Siberia and
elsewhere, such as in parts of western United States (Porter and Swanson , 2008) and the
Pyrenees (Pallas et al., 2010). The amplitudes of the insolation fluctuations in MIS 5 were
large and a major insolation peak followed the trough of MIS 5d before a trough again in
MIS 4. In the southern hemisphere, insolation variability was different with a peak in MIS 4.
However, both hemispheres experienced the lowest insolation during MIS 5, after the last
interglacial (Fig. 4). Evidence of ice build-up in the southern hemisphere during MIS 5 is
limited but Sutherland et al. (2007) and Shulmeister et al. (2010b) did present data suggesting
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a glacier maximum in parts of New Zealand in MIS 5a, whilst in Fiordland, Fink et al. (2006)
suggest that MIS 4 was the largest advance – a situated reiterated in Barrell (2011) for much
of New Zealand.
The continental ice masses and mid-latitude mountain glaciers reached sizes
comparable with MIS 2 during MIS 4, with many of the mid-latitude mountain glaciers in the
northern hemisphere reaching their maximum at this time (Fig. 11). This ice expansion
coincided with minima in solar receipt at 30 and 60° N during MIS 5a and 4 – when
insolation was lower than in MIS 2 (Fig. 4). In the southern hemisphere glaciers also appear
to have been at or close to their maximum extents during MIS 4, but insolation was
unfavourable in comparison to the northern hemisphere (Fig. 11). In the Southern
Hemisphere an insolation peak occurred early in MIS 4 before reducing throughout this
substage, whereas in the Northern Hemisphere the reverse was true with a trough early in
MIS 4 followed by a peak later in the substage (Fig. 11). Based on the current global dataset
(and the error uncertainties associated with the geochronological data) it is not possible to
confidently determine whether these insolation contrasts caused inter-hemispheric variability
in the timings of glaciations within MIS 4. During MIS 3, insolation receipt increased in both
hemispheres and coincided with glacier retreat in most areas, except in areas where monsoons
were an important control on glacier mass balance, such as the Himalayas where some
glaciers reached their maximum extents in this interval (e.g. Owen et al. 2008). MIS 2 was
characterised by another minimum in solar receipt at 30 and 60° N, and coincided with the
largest expansions of the Laurentide Ice Sheet, along with the SE sector of the
Fennoscandanavian Ice Sheet and parts of other ice masses such as the British-Irish Ice Sheet.
The fact that insolation was greater in MIS 2 than in MIS 4 at these latitudes in the northern
hemisphere, yet MIS 2 saw the largest expansions of the Laurentide and other ice sheets is
intriguing. This build-up of large ice sheets in MIS 2 may be explained by positive feedback
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effects caused by the survival of large ice masses during MIS 3 (e.g. Arnold et al., 2002). The
presence of large ice surfaces at high latitudes at a time of decreasing insolation into MIS 2
may have reinforced the effects of this reduction in insolation – a situation which may not
have been matched during the downturn of insolation into MIS 4, when ice masses such as
the Laurentide were building-up from much smaller dimensions (cf. Stokes et al., 2012, their
Fig. 4).
The build-up of the Laurentide Ice Sheet until its maximum in MIS 2 would have
been associated with global cooling as well as an increasing arid global climate (cf. dust
records in Antarctic and Greenland ice cores: Lambert et al., 2008; Wolff et al., 2010; and
Fig. 12 of this paper). Given the fact that global climate was becoming colder and drier in
MIS 2, ice advances seen in many areas of the world at this time are likely to have been
driven by low temperatures rather than increased precipitation (although this remains a
hypothesis requiring further investigation). MIS 4 is also characterised by a sustained period
of elevated concentrations of dust in both the Antarctic and Greenland ice-core records
(Lambert et al., 2008; Wolff et al., 2010) (Fig. 11), suggesting significant global aridity and
reduced hydrological cycle during this interval too (Lambert et al., 2008). The isotope records
from both Antarctica and Greenland show that MIS 4 was a severe and sustained cold
episode comparable to MIS 2 (Petit et al., 1999; Wolff et al., 2010) (Fig. 11). Thus, global
ice build-up during MIS 4 was clearly associated with a severely cold global climate and
culminated with the first recorded Heinrich Event in the North Atlantic, H6 at c. 60 ka
(Hemming et al., 2004).
The behaviour of glaciers around the world during the last glacial cycle requires
simulation within a temporally-dynamic model. Rather than just focusing on the global LGM,
the spatial and temporal variability in the expansions of glaciers needs to be modelled
globally for the whole of the last glacial cycle. This will have implications for understanding
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major atmospheric systems, such as monsoons, depression tracks, and the development of
high latitude continental anticyclones, as well as for understanding the significance of
orbitally-forced solar variability on the global glacier records. However, constraining the
complexity of temperature-precipitation variability for the last glacial cycle, which would
have driven glacier expansion and contraction, is not a simple exercise. It is clear from
vegetation records that millennial-scale environmental changes recorded in the oceans are
clearly imprinted in the terrestrial record (Fletcher et al., 2010). Developing a transfer
function from records such as these to reflect temperature and precipitation changes is the
challenge facing glacier-climate modellers.
Despite the challenges associated with modelling glacier-climate response during the
last glacial cycle, the basic observations demonstrating variability in the timing of the
maximum extents of glaciers around the world challenges the validity of a concept of a
“global LGM”. Given that in so many regions glaciers were not at their most extensive during
the global LGM, the utility of the term and acronym is questionable. The term is misleading
as it clearly does not apply to a large proportion of the global glacial system. This is not to
say that the LGM does not represent the maximum extent of global ice volume. It probably
does, with most of this ice volume locked up in the Laurentide and East Antarctic Ice Sheet
(Hughes, 2013). However, on this basis the LGM is not a spatially representative definition
of a global glacial event. This problem has led to glacial deposits (and indeed all types of
Quaternary deposits) erroneously ascribed to the LGM, which is irrelevant as a stratigraphical
definition in many instances. Of course in many areas, the global LGM coincides with a
major glacial advance, not a maximum advance, just one other substantial advance of several
during the last glacial cycle. Thus, the global LGM is not applicable to glacial stratigraphy as
it is not readily mappable on a global scale since glaciers reached their maximum at different
times. The diachronous glaciations during the last glacial cycle mean that the term has little
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stratigraphical meaning. Mix et al. (2001, p. 635) were well aware that the “local glacial
maxima or temperature minima do not provide a reliable stratigraphic link to the globally
integrated glacial maximum”. Consequently, the LGM was defined in Mix et al. (2001) based
on a combination of dated corals and the isotopic record δ18O in marine sediments, rather
than in glacial sediments themselves, following on from the earlier definitions of CLIMAP
(1976; 1981). Mix et al. (2001, p. 635) argued that the “best estimate of such an integrated
[ice] maximum is based on the equivalent sea level low stand (i.e., after removing local
isostatic effects), followed by benthic foraminiferal δ18O”. Unlike glacial records on land, this
isotope marker is readily mappable on a global scale (e.g. Imbrie et al., 1984; Lisiecki and
Raymo, 2005). However, since the LGM (and isotope boundaries in general) cannot be easily
defined by isochronous boundaries defined in a stratotype sediment or ice-core sequence, it
cannot be used as a chronostratigraphical unit (cf. Salvador, 1994, p. 78; Ehlers et al. 2011).
This is despite attempts to define the LGM as a chronozone by Mix et al. (2001). Some
agreed definition of the LGM is desirable but it does not readily fulfil chronostratigraphical
principles. The global LGM could potentially be better defined within event stratigraphy.
This approach has been adopted for the Greenland ice-core sequence (cf. Björck et al., 1999;
Lowe et al. 2008). Despite this, the purpose of this paper is not to offer a solution to the LGM
problem but to highlight the irrelevance of this term and acronym to the global glacial
records. It is unfortunate that the word 'Glacial' in LGM has led to misleading correlations for
decades in glacial and other Quaternary sedimentary records. Instead, local stratigraphical
frameworks should be constructed using a variety of terrestrial records. This may require the
use of regional terrestrial parasequences that are more continuous than the glacial record (i.e.
lake sediment sequences) as has been suggested by Hughes (2005). Then, from these
comparisons can be made with the marine isotope record which acts as a useful global time-
reference. With geochronological datasets rapidly developing it is becoming increasingly
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apparent that the term LGM is hindering chronological correlation and, in many instances,
providing a 'red-herring' for glacial geologists, geomorphologists and indeed those working
on other terrestrial Quaternary sequences.
5. Conclusions
Glaciers around the world reached their maximum extents at different times during the last
glacial cycle, which spans the period from Termination II to Termination I (from c. 130 ka to
11.7 ka). The Laurentide Ice Sheet maximum corresponds with the global LGM in MIS 2 and
dominates the marine oxygen isotope and sea level records. However, not all the large
continental ice sheets reached a maximum extent at this time. For example, the Barents-Kara
ice sheet in northern Eurasia reached its maximum extent during MIS 4, as did the western
sectors of the Fennoscandinavian Ice Sheet and the northern sector of the Cordilleran Ice
Sheet. In Asia, a vast majority of mountain glaciers reached their maximum positions before
MIS 2, in MIS 5, 4 and 3. The same appears to be valid for New Zealand and SE Australia
and parts of South America. The East Antarctic Ice Sheet also appears to have been out-of-
synchronisation with global ice volume, retreating from its maximum extent before MIS 2.
Whilst many glaciers reached their maximum positions before MIS 2, this is not to say that a
major glacier advance did not occur in the interval. In most areas, with only a few exceptions,
glaciers advanced just before and through MIS 2, reaching the maximum positions of the last
50 ka (Clark et al., 2009). The implications of the widespread evidence for larger glaciers
advances prior to MIS 2, several of which occurred early in the last glacial cycle, are
important because they imply that the global LGM is an ice volume event and not
representative of the maximum spatial extent of continental glaciers. This means that the term
global LGM has limited chronostratigraphical value when used to correlate glacial deposits
and landforms, and is misleading. This situation highlights the problems of using the marine
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isotopic record of global ice volume for correlating terrestrial records on land (cf. Gibbard
and West 2000). This situation is not restricted to glacial sediments and landforms, but also
has relevance for any record that is correlated with the marine isotope sequence based on
assumptions of palaeoenvironmental causation and counting backwards or forwards from
apparent climatic marker events. Instead, the marine isotope record has the greatest value for
terrestrial correlation simply as a broad reference scale for global time-equivalence.
One of the major implications of the asynchrony of maximum glacial extent reported
here for the last glacial cycle is the possibility that this pattern might have been repeated
during earlier glaciations, particularly during the Middle Pleistocene (Ehlers et al. 2011). If
this occurred it could explain why some glaciations are represented in one region of the
world, whilst they are apparently unrepresented elsewhere. In particular, this could possibly
explain why glaciations during periods including MIS 8, 10, 14 or 18 are apparently absent
from regions such as north-western Europe, when others (e.g. MIS 6, 12) are strongly
represented. In the absence of a sufficiently robust and accurate geochronology, it might be
difficult to examine or differentiate between the earlier MIS glacial successions due to the
limitations in accuracy of cosmogenic dating, but it is certain that earlier glaciations were
unlikely to have been any less subject to the variable interactions of the controlling variables
than the last glacial period. Nevertheless, the mutual timing of advances and retreats could
potentially vary to a greater or lesser extent. This possibility could go some way to explaining
the variability in extent, ice volume and chronology of glaciation during the Quaternary and
earlier.
Acknowledgements
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We would like to thank David Fink (Australian Nuclear Science and Technology
Organisation) and Glenn Thackray (Idaho State University) for their very helpful and
thorough reviews of this paper. We thank William Fletcher (University of Manchester) for
important insights into the workings of the global climate system during the last glacial cycle
and for pointing us to useful datasets.
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Figures
Figure 1. Marine and terrestrial chronostratigraphy for the past 150 ka. The marine oxygen
isotope record is from Martinson et al. (1987) whilst the terrestrial climato- and
chronostratigraphical subdivisions are from Lowe and walker (1997). Redrawn from an
original figure in Pillans and Gibbard 2012.
Figure 2. Maximum extent of glaciation around the globe during the last glacial cycle
(Weichselian, Wisconsinan, Valdaian and equivalents). The extents depicted here are
diachronous with ice masses reaching their maximum positions at different times. The point
of this figure is to illustrate the relative sizes of the ice masses and their spatial distributions.
Redrawn and adapted from Ehlers and Gibbard (2007).
Figure 3. Ice volume of the Laurentide ice sheet during the last glacial cycle, redrawn and
adapted based on modelling results presented in Stokes et al. (2012). Ice volume values are
eustatic equivalent metres of sea level (conversion factor of 25.19 m per 1015 m3 of ice).The
shaded parts of the curve represent the margins of error in the modelling (see Stokes et al.,
2012, their Fig. 5a). MIS = Marine Isotope Stage. Here and in all subsequent figures the
boundaries between the stages and substages are based on data in Martinson et al. (1987) and
Lowe and Walker (1997, Chapter 7).
Figure 4. Extent of glaciations in Alaska. Note the difference in size between the larger Early
Wisconsinan and Late Wisconsinan glaciations. This is most pronounced for the Brooks
Range and the Ahklun Mountains.
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Figure 5. Insolation values for the northern and southern hemispheres through the last glacial
cycle (based on data in Berger and Loutre, 1991). MIS = marine isotope stage.
Figure 6. Published cosmogenic exposure ages from glacial landform surfaces in Patagonia,
South America. Data has been grouped from multiple valleys in a specified region. Redrawn
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Figure 7. The evolution of the Barents-Kara Ice Sheet (and neighbouring ice masses) during
the last glacial cycle (Weichselian/Valdaian Stage). Redrawn and modified from Vorren et al.
(2011).
Figure 8. The timings and relative extents of glacier advances in NE Russia during the last
glacial cycle. The relative sizes of glacier advances are schematically depicted by the
amplitudes of the black-filled peaks. Redrawn and adapted from Stauch et al. (2011). MIS =
Marine Isotope Stage.
Figure 9. Summed probability plots of Pleistocene 10Be exposure ages from the Kyrgyz Tien
Shan (A) and the Pamir and Alay Range of western Kyrgystan. These graphs are redrawn
from Koppes et al. (2008) with permission from Elsevier. Probability distributions for
individual ages were normalised to 1 (not shown) such that peaks in the summed probability
plots are influenced by total number of samples. Different production rates and scaling
methods were used for the different data sets; ages are erosion corrected (3 mm/ka). Also
these plots aggregate data from numerous valleys and may hide local variability, which relies
on morphostratigraphical considerations to be taken into account. Even so, the peaks in the
probability plots do reveal meaningful patterns. For example, the probability peaks for the
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Kyrgyz Tien Shan reflect moraines formed by major advances in MIS 5, 4 and 3. The timing
of the actual maximum extents varies between areas (since this is composite dataset) but the
presence of large peaks in MIS 4 and 5 illustrates that these moraines were not overridden by
later glaciers. This point is important in interpreting the probability plot form the Pamir and
Alay ranges. Here the largest peak occurs during MIS 2 (indicating the largest number of
samples with high individual probabilities). However, the more subdued peak in MIS 5/4 is
significant because it means that moraines of this age have not been overridden by glaciers of
MIS 2 age. The smaller amplitude of the probability peak compared with MIS 2 simply
reflects a smaller sample size and the larger 1 sigma spread in individual sample mean ages
for older samples. MIS = Marine Isotope Stage.
Figure 10. Summed probability plots of Pleistocene 10Be exposure ages from the Himalaya
and Tibet. See Fig. 9 caption for information on interpreting the data. Redrawn from Koppes
et al. (2008) based on original data from the West [Hunza, Chitral and Nanga Parbat]
(Phillips et al., 2000; Owen et al., 2002b, c), Central [Everest, Khumbu, NW Garwhal and
Lahul Himalaya (Sharma and Owen, 1996; Richards et al., 2000b; Owen et al., 2001; Finkel
et al., 2003), and East [Tanggula Shan, Nyainqentanggulha Shan and Gonga Shan] Tibet
(Owen et al., 2005). MIS = marine isotope stage.
Figure 11. The timing of the maximum extents of glaciers from selected regions around the
world during the last glacial cycle. The black bars illustrate approximate period of maximum
extent and in some instances reflect diachronous advances within several areas (such as in the
Pamirs, Tien Shan and Altai, for example). Where glaciers reached their maximum at
different times within the same region, or where there is conflicting or contradictory data, this
is indicated by multiple time-bars. The marine oxygen isotope record is based on the stacked
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record in Raymo and Liesiecki (2004) and the Greenland data based on NGRIP (North
Greenland Ice Core Project members, 2004). Marine isotope stages and substage boundaries
follow Martinson et al. (1987).
Figure 12. Dust concentrations in the NGRIP ice core for the last 80 ka (based on data in
Ruth et al., 2007). MIS = Marine Isotope Stage.
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