Invited review
Glacier fluctuations during the past 2000 years
Olga N. Solomina a, *
, Raymond S. Bradley b
, Vincent Jomelli c
, Aslaug Geirsdottir d
,
Darrell S. Kaufman e
, Johannes Koch f
, Nicholas P. McKay e
, Mariano Masiokas g
,
Gifford Miller h
, Atle Nesje i, j
, Kurt Nicolussi k
, Lewis A. Owen l
, Aaron E. Putnam m, n
,
Heinz Wanner o
, Gregory Wiles p
, Bao Yang q
a
Institute of Geography RAS, Staromonetny-29, 119017 Staromonetny, Moscow, Russia
b
Department of Geosciences, University of Massachusetts, Amherst, MA 01003, USA
c
Universite Paris 1 Pantheon-Sorbonne, CNRS Laboratoire de Geographie Physique, 92195 Meudon, France
d
Department of Earth Sciences, University of Iceland, Askja, Sturlugata 7, 101 Reykjavík, Iceland
e
School of Earth Sciences and Environmental Sustainability, Northern Arizona University, Flagstaff, AZ 86011, USA
f
Department of Geography, Brandon University, Brandon, MB R7A 6A9, Canada
g
Instituto Argentino de Nivología, Glaciología y Ciencias Ambientales (IANIGLA), CCT CONICET Mendoza, CC 330 Mendoza, Argentina
h
INSTAAR and Geological Sciences, University of Colorado Boulder, USA
i
Department of Earth Science, University of Bergen, Allegaten 41, N-5007 Bergen, Norway
j
Uni Research Climate AS at Bjerknes Centre for Climate Research, Bergen, Norway
k
Institute of Geography, University of Innsbruck, Innrain 52, 6020 Innsbruck, Austria
l
Department of Geology, University of Cincinnati, Cincinnati, OH 45221, USA
m
School of Earth and Climate Sciences and Climate Change Institute, University of Maine, Orono ME 04469, USA
n
Lamont-Doherty Earth Observatory, 61 Rt 9W, Palisades, NY 10964, USA
o
Institute of Geography and Oeschger Centre for Climate Change Research, University of Bern, Switzerland
p
Department of Geology, The College of Wooster, Wooster, OH 44691, USA
q
Cold and Arid Regions Environmental and Engineering Research Institute, Chinese Academy of Sciences, 320 Donggang West Road, 730000 Lanzhou, China
a r t i c l e i n f o
Article history:
Received 31 May 2015
Received in revised form
2 April 2016
Accepted 12 April 2016
Available online 29 July 2016
Keywords:
Late Holocene
Glacier variations
Neoglacial
Modern glacier retreat
Temperature change
Little Ice Age
Solar and volcanic activity
a b s t r a c t
A global compilation of glacier advances and retreats for the past two millennia grouped by 17 regions
(excluding Antarctica) highlights the nature of glacier fluctuations during the late Holocene. The dataset
includes 275 time series of glacier fluctuations based on historical, tree ring, lake sediment, radiocarbon
and terrestrial cosmogenic nuclide data. The most detailed and reliable series for individual glaciers and
regional compilations are compared with summer temperature and, when available, winter precipitation
reconstructions, the most important parameters for glacier mass balance. In many cases major glacier
advances correlate with multi-decadal periods of decreased summer temperature. In a few cases, such as
in Arctic Alaska and western Canada, some glacier advances occurred during relatively warm wet times.
The timing and scale of glacier fluctuations over the past two millennia varies greatly from region to
region. However, the number of glacier advances shows a clear pattern for the high, mid and low latitudes
and, hence, points to common forcing factors acting at the global scale. Globally, during the first
millennium CE glaciers were smaller than between the advances in 13th to early 20th centuries CE. The
precise extent of glacier retreat in the first millennium is not well defined; however, the most conservative
estimates indicate that during the 1st and 2nd centuries in some regions glaciers were smaller
than at the end of 20th/early 21st centuries. Other periods of glacier retreat are identified regionally
during the 5th and 8th centuries in the European Alps, in the 3rde6th and 9th centuries in Norway,
during the 10the13th centuries in southern Alaska, and in the 18th century in Spitsbergen. However, no
single period of common global glacier retreat of centennial duration, except for the past century, has yet
been identified. In contrast, the view that the Little Ice Age was a period of global glacier expansion
beginning in the 13th century (or earlier) and reaching a maximum in 17the19th centuries is supported
by our data. The pattern of glacier variations in the past two millennia corresponds with cooling in
reconstructed temperature records at the continental and hemispheric scales. The number of glacier
advances also broadly matches periods showing high volcanic activity and low solar irradiance over the
past two millennia, although the resolution of most glacier chronologies is not enough for robust
* Corresponding author.
E-mail address: olgasolomina@yandex.ru (O.N. Solomina).
Contents lists available at ScienceDirect
Quaternary Science Reviews
journal homepage: www.elsevier.com/locate/quascirev
http://dx.doi.org/10.1016/j.quascirev.2016.04.008
0277-3791/© 2016 Elsevier Ltd. All rights reserved.
Quaternary Science Reviews 149 (2016) 61e90
statistical correlations. Glacier retreat in the past 100e150 years corresponds to the anthropogenic global
temperature increase. Many questions concerning the relative strength of forcing factors that drove
glacier variations in the past 2 ka still remain.
© 2016 Elsevier Ltd. All rights reserved.
1. Introduction
Glaciers are sensitive climate proxies and variations in their
length, area and volume provide insights into past climate variability,
placing contemporary changes into a long-term context
(Oerlemans,1994, 2001; Hoelzle et al., 2003). A strong argument for
the sensitivity and reliability of the glacier record is made by their
relatively uniform retreat to contemporary warming (Vaughan
et al., 2013).
A detailed analysis of Holocene glacier fluctuations from 17 regions
and their relationship to potentially important forcing factors
was presented in Solomina et al. (2015). Solomina et al. (2015)
demonstrated a general trend of increasing glacier size from the
early-mid Holocene to the late Holocene in the high and mid latitudes
of the Northern Hemisphere and possible forcing by
orbitally-controlled insolation. Glaciers advanced in the second half
of the Holocene between 4.4 and 4.2 ka (ka ¼ thousand years),
3.8e3.4 ka, 3.3e2.8 ka, at 2.6 ka, between 2.3 and 2.1 ka, 1.5e1.4 ka,
1.2e1.0 ka, and 0.7e0.5 ka. These ice expansions generally correspond
with episodes of cooling in the North Atlantic and provide a
record of cool summers with an approximate century-scale
resolution.
Even though the quality and replication of data on Holocene
glacier variations has dramatically improved in recent decades, the
accuracy of dating advances and retreats is still limited. This dating
uncertainty makes it difficult to draw firm conclusions about the
synchronicity of glacial advances from different regions and
therefore direct comparison with high-resolution proxy reconstructions
is a challenge (Winkler and Matthews, 2010). The
chronology of glacier fluctuations for the past two millennia is a
more reliable climate proxy than earlier in the Holocene, as the ages
of moraines are more precisely known, and the spatial coverage of
the record is more dense. The possibility of dating advances to the
calendar year with tree rings and historical data increases the
resolution of the more recent portion of the records. Additionally,
comparisons with high-resolution reconstructions of summer
temperature and winter precipitation from tree rings, ice cores, and
lake sediments helps identify potential climatic factors that may
have driven glacier fluctuations. Glacier variations themselves can
be used for low-frequency records of temperature and precipitation
especially when they are combined with other proxies (e.g. Dahl
and Nesje, 1996; Luckman and Villalba, 2001; Nesje et al., 2001;
Nesje and Dahl, 2003) or associated with modelling studies
(Leclercq and Oerlemans, 2011; Oerlemans, 2012; Leclercq et al.,
2012; Marzeion et al., 2014).
Glacier variations of the past two millennia are usually considered
in the context of Holocene glacier fluctuations as the latest
part of the “Neoglaciation” (after ~4.5 ka) (Porter and Denton,1967;
special issues of Quaternary Science Reviews v. 7, issue 2, 1988 and
v. 28, 2009 “Holocene and Latest Pleistocene Alpine Glacier Fluctuations:
A Global Perspective”; Wanner et al., 2011). Two other
terms are often used in the discussion of the environmental
changes of the past two millennia: the “Medieval Climatic Anomaly”
(MCA) or “Medieval Warm Period” and the “Little Ice Age”
(LIA). Although there is little agreement in the literature on the
definition of these terms, in this paper we will use the following
approximate boundaries: ~950 to 1250 CE for the MWP/MCA and
~1250 to ~1850 CE for the LIA (IPCC AR5, 2013).
In this paper we will focus on the following questions:
- What were the periods of major glacier advances and retreats in
key mountain regions over the past two millennia, and how
synchronous were they regionally and globally?
- What was the magnitude of glacier fluctuations in the past two
millennia and how does this compare with contemporary
glacier retreat?
- How does the timing and magnitude of glacier variations relate
to orbital, solar, volcanic and greenhouse gas forcings?
2. Approach, data, methods, accuracy of the records
The approach used here is generally the same as in Solomina
et al. (2015) although it is applied to the past two millennia and
focused on higher-frequency glacier variations. We provide a
compilation of glacier fluctuations taking into account both
continuous and discontinuous time series for both glacier advances
and retreats of glaciers. We compile and analyze information about
both individual glaciers and regional summaries (Table 1, Supplementary
Materials (SM)). We selected the best-dated and most
complete time series of glacier fluctuations with preference to
those that extend the full two millennia. Periods of advances and
retreats are identified by historical data, tree remains, ages of
morainesdderived from dendrochronology, terrestrial cosmogenic
nulcides and radiocarbon datingdas well as information from lake
and marine sediments. In Solomina et al. (2015, SM) the reader can
find a detailed description of the methods used for the reconstruction
of glacier fluctuations. Below we briefly summarize the
challenges related specifically to the past two millennia.
Historical data (maps, pictures, written documents) used for the
reconstruction of glacier size is valuable and precise, but limited in
time and to specific regions (Ladurie, 1971; Grove, 2004; Zumbühl
et al., 2008; Masiokas et al., 2009a,b; Nussbaumer et al., 2011a,b;
Nussbaumer and Zumbühl, 2012). In several mountain regions,
for example Scandinavia, the earliest historical data are from the
late 17th century. Information based on tree rings can also be of
high quality and chronologically accurate. Unlike historical sources
tree-ring data can cover longer periods, up to several millennia (e.g.
Nicolussi and Schlüchter, 2012), and can be applied in areas where
the historical descriptions of glaciers are absent. Trees growing on
moraines provide minimum ages of these surfaces (McCarthy and
Luckman, 1993; Wiles et al., 1995; Koch et al., 2009), whereas
trees damaged or tilted by advancing glaciers can be used for more
precise identification of the ages of glacial expansions (Koch et al.,
2007b; Nicolussi and Patzelt, 2001; Nicolussi et al., 2006; Masiokas
et al., 2009a,b; Bushueva and Solomina, 2012; Hochreuther et al.,
2015; Solomina et al., 2015). Interesting results have also been
obtained from wood buried in glacial or glacio-fluvial deposits,
although the link between glacier activity and the death of a tree is
not always clear (e.g. Nazarov et al., 2012).
Dendrochronologically-based calendar-dated glacier chronologies
of high quality and accuracy are available now in several regions,
such as Alaska (Wiles et al., 2011; Barclay et al., 2009a,b,
2013), the Alps (Nicolussi and Patzelt, 2001; Holzhauser et al.,
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9062
2005; Le Roy et al., 2015), British Columbia (Reyes and Clague,
2004; Koch et al., 2007a,b) and the southern Andes (Masiokas
et al., 2009a,b). The most distinct information arises from in situ
tree remains overridden and embedded during advances or buried
by end moraines. Due to a possible loss of some external rings
through decay and abrasion, the ages of advances identified from
this kind of material are maximum-limiting ages, but they are
generally rather close to the actual ages of advances. Le Roy et al.
(2015) showed that the trees that are not rooted but reworked
can also be used for accurate dating of glacier advances. The
youngest age from stratigraphically defined wood layers provides
the age of a glacier advance, whereas the age of the oldest ring of
these trees limits the period of their undisturbed growth and icefree
intervals. This kind of information is extremely important to
identify the scale of former glacial recessions as well.
Radiocarbon dating of organic material is the most common
method for developing glacial chronologies, as it is useful both to
define the age of moraines and to construct the age-depth models
for lake sediments and even ice cores. The development of the AMS
techniques has allowed a substantial increase in the accuracy of this
dating by enabling analyses of smaller sample sizes. One of the
challenges of dating with 14
C is the radiocarbon plateaus when 14
C
concentrations remained constant and thus spanning multiple intercepts
of the radiocarbon calibration curve over long period
(Lotter et al., 1992). In the past two millennia precise 14
C dating is
problematic for the period since ~CE 1600, a period that coincides
with several prominent advances.
The emergence of landscapes from beneath ice after centuries to
millennia has led to new information on past ice extent and to
discussions of their intepretations (Miller et al., 2012; Lowell et al.,
2013; Koch et al., 2014). In situ rooted tundra plants that have brief
life cycles (e.g. mosses) collected at the margins of cold-based nonerosive
ice caps date to the time when they were buried by permanent
snow (“the vegetation-kill-ages”) (Miller et al., 2012, 2013).
Miller et al. (2013) suggested that these are direct records of the
time when ice advanced and then did not retreat until the recent
time as the vegetation is exposed. Lowell et al. (2013) argue with
this concept and defend a different interpretation pointing to a
disagreement between the reconstructions based on the “kill ages”
and the ones based on lake sediment records in Greenland.
We report all ages here, including those based on 14
C analysis, as
CE calendar ages. We use a time scale at the beginning of the
Common Era (labeled as CE); events before the year 1 CE are labeled
as BCE. Centuries and millennia considered here are not labeled as
they are all of the Common Era.
In this paper we treated the 14
C ages of glacier advances and
retreats as the calibrated radiocarbon ages reported in the original
publications. We recalibrated inividual 14
C ages that directly
constrain the timing of glacier fluctuations using the uniform
approach of considering one-sigma intervals, with Calib7.1 (Stuiver
et al., 2005). The calibration curves INTCAL04 and INTCAL13
(Reimer et al., 2013) are identical for the past 2000 years, so the
ages published since 2004 do not require recalibration. For consistency
we report these recalibrated ages as CE; the corresponding
original 14
C ages with the laboratory numbers (when available in
the original publications) are reported in SM Table 1 along with the
reservoir correction details.
Most of the ages on advances and retreats are based on the
assessments by the regional experts inferred from collections of 14
C
ages as well as other sources (geomorphological, stratigraphical,
historical information, tree rings, etc.). In these cases we do no
recalculations and no new assessments, but rely on the choices
made by the experts concerning the age range of glacier events
including the use of one or two standard deviations, multiple intercepts,
etc. In addition, we did not recalibrate the ages for tephras
and the ages for lake, peat, and marine sediments if they provide
age control for models of the rates of sedimentation.
Surface exposure dating using terrestrial cosmogenic nuclides
(TCNs) (10
Be, 14
C, 26
Al, 36
Cl, 3
He, 21
Ne, etc.) is becoming one of the
most widely used approaches to develop glacial chronologies. 10
Be
is the most frequently employed for this purpose and details of the
application of this method are described in several comprehensive
publications (e.g. Gosse and Phillips, 2001; Balco, 2011; Granger
et al., 2013) and in the SM of Solomina et al. (2015). The TCN ages
of moraines are single ages or more commonly the average of
several samples. Recent progress, including the refinement of
regional production rates for TCNs (Balco et al., 2009; Putnam et al.,
2010; Fenton et al., 2011; Kaplan et al., 2011; Goehring et al., 2012;
Young et al., 2013; Blard et al., 2013a,b; Kelly et al., 2013; Stroeven
et al., 2015; Martin et al., 2015), allow moraines to be dated with the
necessary accuracy for the analyses of centennial and sometimes
even multidecadal events for the past two millennia.
In this paper we report the TCN ages of moraines that in most
cases were calculated with the updated local production rates and
the time-dependent Lal/Stone scaling scheme (Stone, 2000).
Recalculation with the most updated local production rate was
made here for the Tropical Andes (Hall et al., 2009; Licciardi et al.,
2009; Jomelli et al., 2014; Stroup et al., 2014, 2015). The ages from
New Zealand by Schaefer et al. (2009) were recalculated and reported
by Putnam et al. (2012) and are reproduced here. The ages
from Greenland (Kelly et al., 2008; Levy et al., 2014) were recalculated
here using the Young et al. (2013) approach. For Central Asia
we report the original ages: Owen et al. (2008) demonstrated that
the difference between different time-dependent scaling models is
about 10% (Lal,1991; Stone, 2000; Desilets et al., 2006; Dunai, 2001;
Lifton et al., 2008, 2014). The TCN ages are reported here in CE with
its requisite standard error. The ka (kiloyears ago) ages are transformed
in CE by subtracting the ka ages from the years of mea-
surement/publication.
Lichenomentry has been widely used to estimate the relative
and numerical age of glacial deposits of the past two millennia
(Beschel, 1950). Recent attempts to improve the lichenometric
technique have been primarily focused on statistical approaches
(Jomelli et al., 2007, 2009; Naveau et al., 2007). On some occasions
the lichenomentric ages generally agree with the reported 10
Be
ages for moraines (Badding et al., 2013; Munroe et al., 2013).
However, several significant problems, including lichen species
identification, growth rate dependence on microenvironmental
conditions and poor chronological control of growth calibration
curves remain unresolved. Osborn et al. (2015) showed that
lichenometric ages may not be reliable and it is a challenge to
distinguish between those that provide realistic and erroneous
ages. We agree with this view, but believe that estimating the age of
deposits using well-defined growth curves can be useful in distinguishing
different generations of moraines (relative dating).
Certainly, lichenometric studies should be undertaken more carefully
and be better documented. Some of the problems mentioned
above may be overcome by applying TCN methods with the
lichenometry. We include lichenometric ages in our paper only if
they are supported by other lines of evidence.
In many cases, information on the former extent of glaciers has
been erased or obscured by subsequent more extensive ice expansions.
Thus other, less direct methods, to infer past glacial activity
have been developed (e.g. Nesje et al., 1991). The use of
lacustrine sedimentary sequences from glacier-fed lakes to reconstruct
Holocene glacier variations in the catchment was pioneered
by Karlen (1976) in Swedish Lapland and widely used in Scandinavia
(Karlen, 1988; Nesje and Kvamme, 1991; Nesje et al., 1991,
1994, 2000a,b, 2001, 2006; Karlen and Matthews, 1992; Dahl
et al., 2003; Lie et al., 2004; Bakke et al., 2005a,b,c, 2010; Vasskog
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e90 63
et al., 2012) as well as in other regions (e.g. Briner et al., 2010;
Larsen et al., 2011; Osborn et al., 2007; Røthe et al., 2015). Physical,
geochemical, and magnetic sediment properties along with
other dating techniques (e.g. 14
C ages calibrated to calendar years
and 210
Pb) can be used to constrain sedimentary based glacier
chronologies. Studies of lacustrine records include loss-on-ignition,
water content, dry weight, dry bulk density, grain-size distribution,
magnetic susceptibility, biogenic silica, X-ray florescence (XRF), and
in situ reflectance spectroscopy. Varves in proglacial lake sediments
are often indicative of glacial activity (Larsen et al., 2011; Zolitschka
et al., 2015), whereas the disappearance of a glacier from a catchment
is usually marked by homogenous organic-rich layers (Karlen
et al., 1999; Zolitschka, 2007).
Peat deposits intercalated with thin mineral layers of glacier
origin exposed along glacier meltwater channels can also yield
information about the timing of glacier activity in the basin (Nesje
and Dahl, 1991a,b; Nesje et al., 1991; Nesje and Rye, 1993; Dahl and
Nesje, 1994, 1996; Matthews et al., 2005). In addition, radiocarbon
dating of palaeosols and peat deposits buried underneath Neoglacial
moraines has been used to reconstruct glacier-size variations
(Mercer and Palacios, 1977; Matthews and Dresser, 1983).
Tephrochronology was applied to constrain the age of moraines in
regions of active volcanism, such as Iceland (Stotter et al., 1999) and
Kamchatka (Solomina, 1999).
3. Glacier fluctuations as climatic proxies. Terminology
Not all glaciers can be used as climate proxies. Large ice sheets
and ice caps, surging glaciers, ice shelves, outlet glaciers, calving
glaciers and some other types of ice bodies have been often shown
to not align with climate forcings (Paterson, 1994). The outlet glaciers
of the northern and southern Patagonian Icefield are classical
examples of the contrasting behaviour of glaciers reflecting more
their internal dynamic than climatic perturbations (Glasser et al.,
2004). Warren and Aniya (1999) drew attention to the instability,
sensitivity to topography, and often indirect response to climate
variations of calving glaciers. Issues related to supraglacial debris
on glacier dynamics throughout the Himalaya and their relationship
with glacial changes have also been discussed in depth, and we
acknowledge the difficulty in assigning glacier fluctuations to climatic
changes in these regions (Hewitt, 1988; Benn et al., 2005;
Owen and Benn, 2005; Scherler et al., 2012). Ideally classical
land-terminated valley glaciers of a simple configuration and
hypsometry are the most suitable to infer paleoclimatic information
(Oerlemans, 1994, 2001). Most glaciers considered in this paper
are of this kind.
As in Solomina et al. (2015), we do not account for the response
time of glaciers here, since the accuracy of most glacial chronologies
even for the past two millennia is generally less than the
response time of mountain glaciers. The lag in position of the
glacier terminus in relation to a climatic perturbation is generally
estimated as 10e50 years (Oerlemans, 1994) and often only a few
years (Bjørk et al., 2012; Imhof et al., 2012). The limitations of the
use of glacier variations as climate proxies, considering their
response times, the climatic drivers of glacier mass-balance fluctuations,
the reaction of “specific” glaciers such as cold-based,
surging glaciers, etc. were discussed in the SM in Solomina et al.
(2015).
The equilibrium-line altitude (ELA) rise or depression (ΔELA) is a
useful paleoclimatic metric that can be assessed either from the
location of terminal moraines or modelled using lake sediment
properties. However a shift in ELA depends on glacier type and the
climatic controls differ from glacier to glacier and from region to
region (Kuhn, 1989; Oerlemans, 2001; Zemp et al., 2007; Six and
Vincent, 2014).
The terminology used to describe glacier fluctuations differs
among workers. A glacier fluctuation is initiated as a glacier expands,
reaches a stabilized position (commonly marked by an end
moraine) and then retreats. In this review the term “advance” is not
necessarily used literally to connote the onset or progression of a
glacier expansion. Instead, the term “advance” is used loosely; it
most commonly refers to the interval when a glacier is near its
maximum position as evidenced by geomorphic features, although
in some cases where stratigraphic evidence is available, the term is
also used to refer to the early stage of glacier expansion. In the case
of lake sediments it encompasses both. We recognize that many of
the ages used to determine the timing of “advances” actually
constrain the timing of moraine stabilization, and therefore
represent the initiation of retreat rather than an “advance”.
Furthermore, in this review the term “glacier activity” is used to
refer to the relative abundance of evidence for expanded glaciers
rather than the activity of a glacier as measured by the down-valley
throughput (flux) of ice. Taking into account the typical uncertainty
associated with the dating ranging from a half century to a century
we assume that this approach, although very rough, is reasonable at
the current stage of knowledge. In most cases the actual precision
of the ages does not allow to distinguish between “the beginning of
the ice retreat” and “the end of the ice advance”, so we consider
here the ages of moraines as approximate time of glacier advances
as have most of our predecessors (e.g. Grove, 2004 and references
herein).
4. Regional descriptions of glacier variations in the past two
millennia
The regional subdivision in this paper (Fig. 1) is similar to
Solomina et al. (2015), who generally followed the recommendations
of IPCC AR5 (Vaughan et al., 2013). We do not include
Antarctica in this paper, as we could not find enough detailed and
well-dated records for individual glaciers.
4.1. Alaska
In the southern coastal region of Alaska, fluctuations of tidewater
glaciers have been shown to not reflect climate directly
(Wiles et al., 1995; Barclay et al., 2009a) on decadal to century-scale
periods. We thus restrict our discussion in Alaska to those glaciers
that are land-terminating and whose chronologies are considered
to more accurately reflect climate (Barclay et al., 2013).
The chronology of land-terminating glaciers in southeast coastal
Alaska, Eagle, Mendenhall, and Herbert glaciers, as well as outlets
of the Juneau Icefield, show advances between CE ~200 and ~320
according to radiocarbon ages (R€othlisberger, 1986). An advance
occurred in the Wrangell Mountains about CE 300 (Wiles et al.,
2002). Other evidence of ice expansion at this time is rare. Widespread
periods of advance between CE 550 and 720 based on 14
C
and tree-ring dating is recognized in the Alaska Range, St. Elias,
Chugach, Kenai Mountains and adjacent ranges in Canada (Reyes
et al., 2006; Barclay et al., 2009b, 2013; Young et al., 2009;
Zander et al., 2013), with most frequently documented expansion
around CE 600 (Reyes et al., 2006). Trees preserved in glacial sediments
near the margin of many retreating glaciers, primarily along
the southern coast, germinated by the CE 950s following the first
millennium advances (Wiles et al., 2008; Barclay et al., 2009a,b)
implying that ice was less extensive during that time or at about the
same extent as present retreating margins. However, ice retreat
was not uniform here, and several glaciers advanced through this
interval (Wiles et al., 1995; Barclay et al., 2009a,b; Koch and Clague,
2011; Fig. 2).
The majority of the advances recognized in the glacial record
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9064
over the past millennium occurred in the CE 1180se1320s,
1540se1710s and 1810se1880s (Barclay et al., 2009a,b, 2013). Most
forefields along the coast and into the more interior Wrangell
Mountains reached their Holocene maxima during the latter CE
1800s. To the west in the Aklun Mountains, glaciers reached their
maximum Holocene extent as early as CE 1300 with less extensive
moraine building through CE 1800 (Levy et al., 2004). Along the
southern coast, most glaciers reached their Holocene maxima in
the past few centuries. However, some glaciers in the Alaska Range
(Bijkerk, 1984) and one in the Wrangell Mountains (Wiles et al.,
2002) show a more extensive first millennium advance.
In general summer temperature is the primary limiting factor
for land-terminating glaciers along the southern coast of Alaska in
the Kenai, Chugach, St. Elias and the Coast Mountains (Barclay et al.,
1999, 2013; Wiles et al., 2004) (Fig. 2). Advances during the first
millennium CE correspond with cooling recognized in lake sediments
at Farewell Lake in the Alaska Range (Hu et al., 2001) at
about this time. The interval of retracted ice margins about CE 950
corresponds with warming along the southern coast recognized in
a regional dendroclimatic reconstruction (Wiles et al., 2014).
The majority of advances occur within the intervals of cooling
apparent in a February-August temperature reconstruction for
coastal southern Alaska with temperature minima centered on CE
1200, 1450, 1650 and the 1800s (Fig. 2) (Wiles et al., 2014). A JulyAugust
summer temperature reconstruction based on latewood
density in the Brooks Range (Anchukaitis et al., 2013) shows a
general similarity with glacial records of the past nine centuries,
with cooler summers and ice expansion coinciding with the CE
1300 and CE 1800 advances. However, the tree-ring-based temperature
reconstruction (Anchukaitis et al., 2013) and a summer
temperature record derived from sedimentary chlorophyll (Boldt
et al., 2015) show that the CE 1600 advance occurred during a
time of warm summers, and thus this advance may have been
facilitated by increased winter precipitation. In the 20th century,
glacier retreat for land-terminating glaciers has dominated in
agreement with rising temperature (Molnia, 2008).
4.2. Western Canada and US
Evidence for glacier advances during the first millennium CE is
widespread in the western Cordillera of North America outside of
Alaska (Luckman, 2000; Reyes and Clague, 2004; Allen and Smith,
2007; Koch et al., 2007a; Jackson et al., 2008; Menounos et al.,
2009; Clague et al., 2010; Samolczyk et al., 2010; Bowerman and
Clark, 2011; Johnson and Smith, 2012; Maurer et al., 2012;
Coulthard et al., 2012; Munroe et al., 2012; Osborn et al., 2012;
Craig and Smith, 2013; Hoffman and Smith, 2013; Mood and
Smith, 2015), but glacier extent generally appears to have been
smaller than during advances of the past millennium. Some glaciers
were likely smaller before CE 500 than in the late 20th century
(Allen and Smith, 2007; Clague et al., 2010), even though other
glaciers seem to have advanced during that same period (Jackson
et al., 2008; Samolczyk et al., 2010; Maurer et al., 2012; Osborn
et al., 2012; Hoffman and Smith, 2013). Many locations had advances
between CE 400e600, before glaciers retreated to some
degree. However, no sites studied show evidence that glaciers
receded to sizes similar to the late 20th century after the first
millennium advance. Rather, numerous sites provide evidence of
significant advances during the MCA between the advances of the
first millennium and those of the early 2nd millennium CE (Koch
and Clague, 2011; Johnson and Smith, 2012; Munroe et al., 2012;
Osborn et al., 2013).
Most glaciers reached their maximum Holocene extent in the
second millennium CE (Luckman, 2000; Menounos et al., 2009),
prior to CE 1850 (Koch et al., 2007b; Clague et al., 2010; Hoffman
and Smith, 2013), and remained at their maxima until the early
20th century (Luckman, 2000; Koch et al., 2007b; Menounos et al.,
2009). Glaciers fluctuated, but were more extensive than at present,
between about CE 1200e1500, and it appears that several glaciers
advanced synchronously during this period in the Rocky and Coast
Fig. 1. Spatial distribution of time series used in this paper. 1. Alaska; 2. Western Canada and US; 3. Arctic Canada; 4. Greenland; 5. Iceland; 6. Svalbard; 7. Scandinavia; 8. Russian
Arctic; 9. North Asia; 10. Central Europe; 11. Caucasus and Middle East; 12. Central Asia (semi-arid); 13. Central Asia (monsoon); 14. Low Latitudes; 15. Desert-Central Andes of Chile
and Argentina. 16. South of South America (North Patagonian Andes; North Patagonian Icefield, South Patagonian Icefield and adjacent glaciers, Magallanes region e Tierra del
Fuego) 17. New Zealand. For individual time series description see Table 1 in SM.
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e90 65
Fig. 2. Regional glacier variations and regional climate proxies. Alaska. Summer temperature (July-August) reconstruction inferred from maximum latewood density of tree rings
for the Firth River, northwestern Alaska (Anchukaitis et al., 2013) (A), February-August temperature reconstruction for southern coastal Alaska based on ring widths (Wiles et al.,
2014) (B), Glacier Expansion Index (GEI, Wiles et al., 2004) derived from glacial histories from the Arctic Brooks Range to the southern coastal regions across Alaska. The record is
based on ages of moraines and intervals of ice expansion dated with radiocarbon, lichenometry and tree rings (C). General intervals of ice advance for the past 2 ka summarized in
the text (D). Vertical stripes e intervals of glacier advances corresponding to coolings. Western Canada. Reconstructed relative glacier extent in western North America (bold black
line) plotted with a reconstruction of 30 year averages of annual mean temperature deviations from a 1904 to 1980 average for temperate North America (30e55N, 75e130W)
based on pollen data (blue line) and on tree-ring data (red line). Light (medium) blue zones indicate 2SE (1SE) uncertainty estimations associated with each 30 year value. Also
plotted is the comparably smoothed instrumental temperature values up to 1980 (fine black line). All data other than glacier extent is modified from Trouet et al. (2013). Arctic
Canada. Cumulative probability density function of 118 calibrated radiocarbon ages on in situ tundra plants collected within 1 m of the margin of retreating ice caps across Baffin
Island, Arctic Canada. Clusters of ages define periods when colder summers lowered snowline, entombing living plants, and remaining across the site until shortly before the year of
their collection (CE 2005 to 2010). The ages record periods of glacier advance. Data from Miller et al. (2012, 2013). Iceland. Temperature anomaly based on composite proxy records
from Haukadaslsvatn and Hvitarvatn. Hvitarvatn varve thickness and glacial advances in Iceland (triangles) (Larsen et al., 2011; Geirsdottir et al., 2013). Scandinavia. Composite
record of Scandinavian glacier variations during the past two millennia based on continuous records from glacier-fed lakes (A). Records from Northern Folgefonna (Bakke et al.,
2005b), Grovabreen (Seierstad et al., 2002), Jostedalsbreen (Nesje et al., 1991, 2001; Vasskog et al., 2012), Spørteggbreen (Nesje et al., 2006), Breheimen (Shakesby et al., 2007),
Jotunheimen (Matthews et al., 2000; Matthews and Dresser, 2008), Austre Okstindbreen (Bakke et al., 2010), Lyngen (Bakke et al., 2005a), Langfjordjøkelen (Wittmeier, 2014), and
Northern Sweden (Rosqvist et al., 2004). For details, see original publications. Summer temperature reconstruction based on the Tornetr€ask pine ring-width chronology (Grudd
et al., 2002) (B). Russian Arctic. Glacier advances in Franz Josef Land (Lubinski et al., 1999) (A). Ice-core records from Windy Dome, Garham Bell (Henderson, 2002): accumulation
(B), melt features e summer temperature proxies (C). Glacier fluctuations in Novaya Zemlya (Forman et al., 1999; Zeeberg and Forman, 2001; Murdmaa et al., 2004) (D), JuneJuly
temperature tree-ring based reconstruction in NW Sibiria (Briffa et al., 2013) (E). Orange vertical stripes e glacier advances corresponding to coolings, gray stripe e advance
corresponding to high temperature and high accumulation. Altay. Glacier fluctuations in Altay Mountains (Nazarov et al., 2012), summer temperature sensitive tree-ring chronology
from Mongun-Tayga Mountains (Myglan et al., 2012a,b). Orange vertical stripes e periods of glacier advances corresponding to summer coolings, gray stripe e glacier retreat and a
warming of the first half of the first millennium CE. Alps. Glacier length changes in the Alps. Great Aletsch Glacier (Holzhauser et al., 2005) (upper), Mer de Glace (Le Roy et al., 2015)
(middle), Tree-ring width chronology in the Alps (sensitive to summer temperature), red curve e 30-year running means, black curve e 50 year averages (Nicolussi et al., 2009)
(lower). Vertical orange stripes outline the glacier advances corresponding to the coolings. Gray stripe e glacier retreat and warming during the MCA. Tibet (temperate monsoon
glaciers). Dendrochronological ages for moraines in Tibet (Br€auning, 2006; Zhu et al., 2013; Xu et al., 2012; Hochreuther et al., 2015; Loibl et al., 2015) in comparison with decadal
summer temperature reconstructions (summer temperature anomalies with respect to long-term average (1000e2005 CE) (Wang et al., 2015) (A). Relative glacier extent based on
14
C dating (Yang et al., 2008) (B), composite temperature records in Tibetan Plateau (standardized deviations with respect to the past two millennia) (Yang et al., 2003) (C),
speleothem d18
O record in Dongge Cave reflecting the south Asian monsoon variability (Wang et al., 2005) (D). Southern South America. Paleoclimatic proxies and number of
glacier advances in South America. Southern Annular Mode (SAM)-like centennial changes in southern South American climate derived from a stratigraphic record of non-arboreal
pollen, Lago Cipreses (51S) (Moreno et al., 2014). The persistently positive (negative) phases of SAM are associated with warm/dry (cold/wet) climate conditions in southern South
America (upper). Number of glacial advances in subregions of South America (lower). Darker cells indicate synchronous advances from different sites for the 100 year periods;
lighter cells indicate isolated or less synchronous events. Each cell indicates the number of glacier advances identified for each 100 year period (R€othlisberger, 1986; Aniya, 1995;
Strelin et al., 2008, 2014; Masiokas et al., 2009a; Aniya and Skvarca, 2012). 1 e Desert central Andes of Chile and Argentina, 2 e North Patagonian Andes, 3 e North Patagonian
Icefield, 4 e South Patagonian Icefield and adjacent glaciers; 5 e Magallanes region e Tiera del Fuego. New Zealand. Summer temperature reconstruction (Cook et al., 2002) (A) and
glacier variations in New Zealand (from Schaefer et al., 2009; Putnam et al., 2012) (B). Vertical stripes indicate most prominent coolings corresponding to glacier advances.
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9066
mountains (Luckman, 1995; Koch et al., 2007b; Coulthard et al.,
2012). Some moraines in western Canada were possibly deposited
during this advance (Koch et al., 2007b; Osborn et al., 2007; Koehler
and Smith, 2011). Major moraine-building phases in the mountains
of western North America date to CE 1690e1730, 1830e1850,
1860e1890, and 1910e1930 (Luckman, 2000; Luckman and
Villalba, 2001; Osborn et al., 2001; Allen and Smith, 2007; Koch
et al., 2007b; Osborn et al., 2007, 2012; Jackson et al., 2008;
Clague et al., 2010; Bowerman and Clark, 2011; Koehler and
Smith, 2011; Maurer et al., 2012; Munroe et al., 2012; Coulthard
et al., 2012; Mood and Smith, 2015).
A tree-ring record from the central Canadian Rockies extends
from CE 950e1994 and provides a reconstruction of May-August
maximum temperature. Cold periods occurred in CE 950e1000,
1200e1320,1450e1880, with the CE 1690s being exceptionally cold
(Luckman and Wilson, 2005). A tree-ring chronology from the
central British Columbia Coast Mountains spans CE 1225e2010 and
allows the reconstruction of regional June-July air temperature.
Intervals of below average temperatures occurred CE 1350e1420,
1475e1550, 1625e1700, and 1830e1940 (Pitman and Smith, 2012).
A reconstruction based on tree rings and pollen data (Viau et al.,
2012) for temperate North America for the past 1500 years shows
below average temperature for CE 480e750 and CE 1100e1900
with distinct cool periods around CE 1300, 1650e1700, and
1800e1900 (Trouet et al., 2013) (Fig. 2).
Annual precipitation (July-June) has been reconstructed for the
southern Canadian Cordillera for up to the past 600 years with treering
chronologies. Wet conditions were reconstructed for most
chronologies for CE 1680e1700, 1750e55, 1780e1790, 1800e1830,
and 1880e1890. Earlier wet periods, with fewer chronologies
dating back this far, are CE 1540e1560, 1600e1620, and 1660e1670
(Watson and Luckman, 2004). Winter precipitation (NovemberFebruary)
in eastern Washington State was reconstructed for the
past 1500 years from lake sediment oxygen isotope data, and above
average wet periods occurred CE 560e730, 770e840, 900e960,
1050e1150, 1300e1460, and 1570e1620 (Steinman et al., 2012).
Spring droughts for southern Vancouver Island on the west coast of
British Columbia have been reconstructed from tree rings. The reconstructions
record severe and multi-decadal droughts in CE
150e200, 540e570, 760e810, 1440e1570, and 1845e1850 (Zhang
and Hebda, 2005).
The climate reconstructions outlined above make it difficult to
simplify the direct forcing of climate on glacier fluctuations due to
the very large area under discussion with complex climatic pattern
and great variety of glacier types. In the Canadian Rocky Mountains
glacier advances coincide with low summer temperature in CE
1200se1300s, late CE 1600s through early CE 1700s, and in the 19th
century, while precipitation changes do not show any significant
correlation with the ages for glacier advances (Luckman and
Villalba, 2001). In general, especially during the second half of the
past millennium, it appears that below-average summer temperatures
and above-average wet periods can account for much of the
reconstructed glacier advances; however, there are exceptions such
as during CE 800e1400, when relatively warm summers and
wetter winters are reconstructed, which likely explain the significant
glacier advances in western North America during this time
(e.g. Llewellyn Glacier in northwest British Columbia, glaciers in
Garibaldi Provincial Park in southwest British Columbia, Robson,
Kiwa, and Peyto glaciers in the Canadian Rockies (Koch and Clague,
2011).
4.3. Canadian Arctic
A sufficient number of 14
C ages of in situ rooted tundra plants
found at glacier margins are now available for Baffin Island, Arctic
Canada, for the past two millennia to allow reconstruction of periods
of ice expansion (Anderson et al., 2008; Miller et al., 2012,
2013; Margreth et al., 2014). Clusters of kill ages suggest a significant
expansion of local glaciers and ice caps on Baffin Island
occurred early in the first millennium CE, between CE 250 and 400,
with a second expansion interval between CE 800 and 950 (Fig. 2).
There is little evidence for ice advance between CE 950 and 1250,
but widespread ice expansion is recorded beginning about CE 1260,
with irregular but generally continuously colder summers from CE
1280 until 1450 (Fig. 2). The signal provided by entombed plants
vanishes after CE 1450 because ice cover was widespread. Historical
observations and distinct lichen trimlines indicate that the maxima
were achieved in the late CE 1800s (Miller et al., 2013). Margreth
et al. (2014) show a similar record from southern Cumberland
Peninsula, southeastern Baffin Island, with significant ice expansion
beginning early in the first millennium CE. Clusters of kill ages
provide evidence for expansion CE 400e500, and between CE 600
and 900, little evidence of ice expansion during the CE 11th-early
13th centuries, and an early onset of significant glacier advances
starting CE 1280 and a second pulse of advances around CE 1460. In
the 19th century the snowline in Baffin Island was almost 200 m
below its position of the end of the 20th century, while about two
millennia ago it was at least 200 m above the modern level (Miller
et al., 2013).
4.4. Greenland
Scarce data exist to define the extent of glaciers during the past
two millennia in Greenland. Consequently it is still difficult to
discuss potential synchronicity between the different regions, and
correspondence of glacier fluctuation patterns to the regional climatic
patterns at a multidecadal scale even for the relatively welldocumented
past century (Hall et al., 2008; Carlson et al., 2008;
Weidick, 2009; Hughes et al., 2012).
The reconstructions of glacier fluctuations in Greenland are
based on historical information (Weidick, 1958, 1968; Bennike and
Weidick, 2001), 14
C dating (Bennike, 2002; Kelly and Lowell, 2009;
Bennike and Sparrenbom, 2007; Knudsen et al., 2008; Kelly and
Lowell, 2009), TCN dating (Kelly et al., 2008; Lowell et al., 2013;
Levy et al., 2014; Winsor et al., 2014; Young et al., 2015), and proglacial-lake
and marine sediments (Lloyd, 2006; Briner et al., 2010,
2011, 2013; Young et al., 2011; Larsen et al., 2011; Kelley et al., 2012;
Balascio et al., 2015). The information was collected at the margins
of ice-sheet-outlet glaciers (both marine/land terminated), as well
as from locally sourced glaciers.
In northern Greenland (Washington Land) reworked shells in a
moraine of Humboldt Glacier attest to an advance around CE 1300
(Bennike, 2002). The glacier has receded since its maximum
advance around 100 years ago.
In central west Greenland marine cores just beyond the fjord
mouth suggest that Jakobshavn Isbræ reached late Holocene
maxima between CE 1200 and 1660 (Lloyd, 2006). Additional 14
C
ages from plant macrofossils and varve chronologies from glacial
lakes located close to Jakobshavn Isbræ show that both land-based
and marine-based ice margins advanced between CE 1500 and
1640 and reached a maximum at 1850 (Briner et al., 2010, 2011;
Young et al., 2011). This contrasts with other portions of the
western land- terminating ice sheet close (c. 40 km) to Jakobshavn
Isbræ, which achieved their maxima during the 20th century
(Kelley et al., 2012). Most probably the difference is explained by
the complex ice dynamic of marine-based and land-terminating
parts of the ice sheet.
In western Greenland (Nuuk region) Weidick et al. (2012) reported
asynchronous advances for land- and marine-terminating
outlet glaciers over the past centuries from historical observations
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e90 67
and lake sediment analysis. The land-terminating Qamanaarssup
Sermia and Kangaarssup Sermia may have reached their LIA maxima
in the mid CE 1800s and in the early 20th century, respectively.
During the past millennium marine- terminating Saqqap Sermersua
achieved its maximum in the early 2000s, while Nakkaasorsuap
Glacier reached its maximum in the middle of the 19th century
(Weidick et al., 2012). A maximum advance over the past centuries
of the Kangiata Nunaata Sermia ice sheet marine outlet glacier at CE
300 was deduced from the analysis of lake sediments (Weidick
et al., 2012).
In southwest Greenland, 10
Be ages the Narsarsuaq moraine of
land-terminating Kiagtut Sermia at CE 490 ± 110 (Winsor et al.,
2014). A second set of 10
Be ages from a more ice-proximal position
of Kiagtut Sermia shows that ice has been within or near its
present extent since CE 660 ± 150.
Inland ice and local glaciers in south and southwest Greenland
began to advance around CE 1600 and reached maximum positions
in CE 1750 and 1890e1900 (Weidick, 1968; Kelly and Lowell, 2009).
Greenland inland ice and glaciers have been significantly retreating
during the 20th century (e.g. Kjeldsen et al., 2015).
In southeastern Greenland at a nunatak located at the presentday
ELA of Mittivakkat Glacier, organic material dating back to CE
500e600 was found, indicating a warmer interval and a subsequent
glacier advance (Humlum and Christiansen, 2008; Knudsen et al.,
2008; Kelly and Lowell, 2009).
In northeastern Scoresby Sound, 10
Be ages show direct evidence
of glacier advances between CE 1100 and 1660 (Kelly et al., 2008;
Levy et al., 2014 e recalculated using Young et al., 2013 production
rate and Lal/Stone scaling scheme). In the same region, the
maximum Holocene advance of the local Istorvet Glacier occurred
between CE 1150 and 1660 based on radiocarbon and 10
Be ages
(Lowell et al., 2013).
4.5. Iceland
In Iceland the most prominent evidence for glacier advances
during the late Holocene is documented by glacier advances between
CE 1300 and 1900, although, moraines formed by small
glaciers also indicate fluctuations before CE 1300 (Caseldine, 1987;
Caseldine and St€otter, 1993; Stotter et al., 1999; Kirkbride and
Dugmore, 2001, 2006, 2008; Schomacker et al., 2003; Principato,
2008).
The most rigorously dated record of glacier advances in Iceland
over the past two millennia is found in Hvítarvatn, a glacial lake
adjacent to the second largest glacier of Iceland, Langj€okull. A
multibeam sonar bathymetry of the lake together with seismic
reflection analyses revealed multiple glacial advances, which have
been precisely dated with varve counts supported by the known
tephra layers found within the sediment (Black et al., 2004;
Geirsdottir et al., 2008; Larsen et al., 2011, 2013, 2015; Geirsdottir
et al., 2015). The evidence from Hvítarvatn demonstrates that the
two outlet glaciers of Langj€okull that drain into Hvítarvatn
(Suðurj€okull and Norðurj€okull) have response times of about 100
years (Black et al., 2004; Flowers et al., 2008; Larsen et al., 2015).
Langj€okull achieved its maximum Neoglacial extent between CE
1700 and 1930, when the two outlet glaciers advanced into the lake
and maintained active calving margins. Paleolimnological studies
of sediment cores from Hvítarvatn indicate glacier expansion in the
5th and 13th centuries. Norðurj€okull advanced into Hvítarvatn
around CE 1720, and remained at or near its maximum for most of
the 19th century, whereas Suðurj€okull underwent a quasi-periodic
series of eight surges between CE 1828 and 1930 (Larsen et al.,
2015). A very similar pattern of glacier expansion is seen around
small ice caps in northern Iceland where moraine features were
mapped and dated by the use of tephrochronology and 14
C ages
(Stotter et al., 1999).
The combined paleolimnology and glacier advance study from
Hvítarvatn, central Iceland, shows a stepwise intensification of
glacial activity until its culmination between CE 1300 and 1900
(Fig. 2; Larsen et al., 2011; Geirsdottir et al., 2013). Langdon et al.
(2011) identify particularly cold phases based on chironomid
studies in NW Iceland for CE 1683e1710, 1765e1780, and
1890e1917, with summer temperature decreases of 1.5e2.0 C
suggesting that the magnitude of summer temperature cooling
resulted in Icelandic glaciers reaching their maximum Holocene
extent at that time, as previously modelled for Langj€okull (Flowers
et al., 2007).
In south-central Iceland a combination of tephrochronology and
lichenometry was applied to date moraines, tills and meltwater
deposits. Three glacier advances occurred between CE 450 and 550,
and three more occurred between CE 900 and 1400. Five groups of
advances occurred between CE 1650 and 1930 (Kirkbride and
Dugmore, 2008). While glacier advances between CE 1600 and
1900 across Iceland appear to have been synchronous, the timing of
maxima differs between glacier type and region from the early 18th
to the late 19th century (e.g. Chenet et al., 2010).
4.6. Spitsbergen
Radiocarbon ages on organic material buried at or under the
modern glaciers (Baranowski and Karlen, 1976; Humlum et al.,
2005), sediment cores from proglacial lakes (Svendsen and
Mangerud, 1997; Røthe et al., 2015), lichenometry (Werner, 1993)
and 10
Be ages on moraines (Reusche et al., 2014) have all been used
to reconstruct glacier fluctuations of the past two millennia in
Spitsbergen.
Since the 1970s several authors reported the ages for bulk
organic material buried at glacier fronts in Spitsbergen and dating
back to the first millennium CE, claiming that the glaciers were
smaller or of equal size then compared to the end of the 20th
century (e.g. Baranowski and Karlen, 1976; Furrer, 1992; Furrer
et al., 1991). Humlum et al. (2005) found frozen soil and vegetation
in situ below the cold-based Longyearbreen Glacier. One soil
and ten bryophyte samples were dated and yielded ages between
CE 20 and 820, indicating that the glacier was 2 km shorter than in
the early 21st century for at least 800 years and possibly longer.
Snyder et al. (2000) demonstrated that the Linnevatnet cirque was
ice free until about CE 1400.
However there is also evidence of several glacial advances in
Spitsbergen in the first millennium CE. An advance at CE 350 ± 200
is documented by 10
Be dating of a moraine at Linnebreen in western
Svalbard (Reusche et al., 2014). The advance was close or even a
little larger than the past millennium maximum extent, that the
glacier occupied up to CE 1930 (Svendsen and Mangerud, 1997).
Somewhat earlier, an advance of Karlbreen around CE 250 is
recorded in lake sediments at Mitrahalvøya (Røthe et al., 2015).
Sediment records in Kongressvatnet point to possible minor glacial
events at CE 700e820 and CE 1160e1255 (Guilizzoni et al., 2006),
but this evidence needs replication and confirmation.
Glacial activity in Spitsbergen increased in the mid 14th century.
Continuous sequences of glacial varves in Kongressvatnet were
deposited between CE 1350 and 1880 (Guilizzoni et al., 2006). The
deposition of glacial sediments in Billefjorden indicates the increase
of activity of Nordenskioldbreen at CE 1520e1900, or
perhaps a little earlier (by CE 1425) (Szczucinski et al., 2009).
Lake sediments at Mitrahalvøya provide information that
Karlbreen was probably close to its Holocene maximum at CE 1725
and 1815 (Røthe et al., 2015). Many glaciers in Spitsbergen reached
their past millennium, and sometimes even Holocene, maxima very
recently, at the end of the 19th or early 20th century (e.g. Baeten
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9068
et al., 2010). Mangerud and Landvik (2007) demonstrated that the
outermost moraine of Scottbreen Glacier was deposited about CE
1900 and overlies a marine shoreline dated to ~9250 BCE. This indicates
that over the Holocene the glacier was never more extensive
than it was at the turn of the 20th century. Liestøl (1988)
estimated that the equilibrium-line altitude depression (DELA)
driving the recent maximum in Spitsbergen was about 100 m,
although the reconstruction based on lake sediment provides more
restricted estimates between 40 and 50 m (Røthe et al., 2015).
4.7. Scandinavia
Karlen (1988) and Karlen and Kuylenstierna (1996) recognized
several Holocene advances in Scandinavia around CE > 1, 100e400,
800e1000, and 1300e1800. This chronology was based on historical,
lichenometric, and dendrochronological moraine data, treeline
variations, and lacustrine sediments. The record shows that
in Swedish Lapland, glaciers most likely attained their greatest past
millennium extent between the 17th and early 18th centuries.
Rosqvist et al. (2004) used sediments in Lake Vuolep, Allakasjaure
to suggest periods of advanced glacier positions at CE 200, and
during the past 1300 years, with a maximum in CE 1700e1800. In
southern Norway, Matthews and Dresser (2008) indicated that the
most prominent Neoglacial maxima during the past two millennia
were centered at intervals CE 100e400, 600e700, 900e1000,
1200e1300, and 1600e1800.
Data from glacier-fed lakes provide continuous histories of
glacier variations during the past two millennia. Records of glacial
activity from upstream catchments, have been compiled from
northern Folgefonna (Bakke et al., 2005b), Grovabreen (Seierstad
et al., 2002), Jostedalsbreen (Nesje et al., 1991, 2001; Vasskog
et al., 2012), Spørteggbreen (Nesje et al., 2006), Breheimen
(Shakesby et al., 2007), Jotunheimen (Matthews et al., 2000;
Matthews and Dresser, 2008), Austre Okstindbreen (Bakke et al.,
2010), Lyngen (Bakke et al., 2005a), Langfjordjøkelen (Wittmeier
et al., 2015), and northern Sweden (Rosqvist et al., 2004). The
composite record (Fig. 2) indicates that glaciers in Scandinavia were
in an advanced state between CE 200e300, 600e700, 800e1000,
and 1500e1800, whereas they were in a retracted state during CE
1e200 (Roman Period), 300e600, 700e800, 1000e1500, and during
the 19the20th centuries.
Lichenometric dating and historical evidence indicate that the
timing of the past millennium maximum position of outlet glaciers
for ice caps and valley glaciers in southern Norway varied considerably,
ranging from the early 18th to the late 19th century
(Bickerton and Matthews, 1993). Differences in glacier hypsometry,
frontal time lag, and responses to climatic parameters (e.g. winter
precipitation and summer temperature) may explain the differences
in the timing (e.g. Aa, 1996; Nesje, 2005; Nesje et al., 2008).
The majority of dated terminal moraines in southern Norway
cluster around CE 1740e50, 1780e90, 1860e70, and 1920e40
(Nesje et al., 2008). However, on decadal scale these variations in
southern Norway do not show a consistent regional pattern
(Bickerton and Matthews, 1993; Winkler et al., 2003; Matthews,
2005). This is because glacier mass-balance measurements show
that the annual (net) mass balance of maritime glaciers in western
Norway is mostly controlled by winter precipitation, while continental
glaciers best correlate with the summer temperature (e.g.
Nesje et al., 1995a,b; Mernild et al., 2014; Trachsel and Nesje, 2015).
4.8. Russian Arctic
4.8.1. Novaya Zemlya
Marine sediment cores from Russkaya Gavan’ (NW Novaya
Zemlya) where the outlet of Shokal’ski Glacier terminates show
that between ~CE 1170 and ~CE 1400 Shokal’ski Glacier was in a
contracted position and was contributing limited sediment to the
fjord (Murdmaa et al., 2004; Polyak et al., 2004). An advance
occurred around CE 1400, and between CE ~1470 and ~1600 the
glacier front was relatively stable before major glacier retreat
started. Based on the grain sizes of marine sediments Zeeberg et al.
(2003) suggested another advance between CE 1700 and 2000,
most likely in the 19th century.
No local high-resolution records are available in Novaya Zemlya;
however, the most prominent advance around CE 1400 occurred
during an interval of increased winter precipitation and higher
cyclonic activity over the North Atlantic as interpreted from ion
content in the GISP-2 ice-core record (Zeeberg and Forman, 2000;
Meeker and Mayewski, 2002; Murdmaa et al., 2004). The advance
in the 19th century may also have been forced by an increase in
winter precipitation. This later advance also corresponds with low
summer temperatures as indicated by tree-ring reconstruction of
June-July temperature from northwest Siberia (Briffa et al., 2013)
(Fig. 2).
4.8.2. Franz Josef Land
Our understanding of the fluctuations of glaciers on Franz Josef
Land is based on a large collection of 14
C ages (details are provided
in the SM Table 1) constraining positions of 16 glaciers relative to
1991e1995 (Lubinski et al., 1999). The 14
C age on in situ moss at the
Glacier B margin (Southern Hall Island) indicates an advance at 409
BCE e CE 223. These data also indicate that the glacier was never
less extensive than today over the past 2000 years. At Sonklar
Glacier (Southern Hall Island) a lateral moraine is dated to CE
718e885 by a marine shell either overlain by or incorporated inside
the moraine. Yuri Glacier advanced 1.5 km beyond its 1990 margin
around CE 778e1021 and receded prior to CE 1191e1285. At Cape
Lagerny (Northbrook Island) the glacier was beyond the present
margin until at least CE 545e671. Moss found in situ at the present
margins and killed by advances of glaciers dates to CE 1054e1256,
1519e1950, 1646e1950, and 1684e1928 at Sedov Ice Cap (Western
Hooker Island), CE 975e1145 at Cape Lagerny Glacier, CE
1487e1641 at Leight Smith Island, and to CE 1651e1950 at Cape
Flora Glacier. Based on these 14
C ages and the geomorphology
Lubinski et al. (1999) identified glacier advances during the ~10th,
and 12th centuries, at CE 1400, 1600, and in the early 20th century.
Between major advances some glaciers were smaller than at the
end of the 20th century. The ice expansion that occurred around CE
1000 was the most prominent in the past two millennia (Lubinski
et al., 1999; Forman and Weihe, 2000).
An ice core from Windy Dome provides information on accumulation
and summer temperature (melt features) for CE
1225e1997 (Henderson, 2002) (Fig. 2). The advance that started at
the end of the 19th century coincides with high accumulation rates.
Unfortunately the radiocarbon ages are not precise enough for
direct comparison of the time of other advances with the highresolution
ice-core records.
4.9. Northern Asia
4.9.1. Kamchatka
Many glaciers in Kamchatka are located on active volcanoes and
therefore are strongly impacted by volcanic activity (Vinogradov
et al., 1985; Barr and Solomina, 2014), which has to be taken into
account when interpreting Kamchatka glacier fluctuations. Historical
data, tephrochronology, lichenometry and, to a limited
extent, tree rings were used to date moraines in Kamchatka (Barr
and Solomina, 2014). Several glaciers record advances of up to
4 km in length in the first millennium CE, but tephrochronological
dating of moraines is still very broad. Bilchenok Glacier (surging in
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e90 69
the 1960s) advanced beyond its present limit around 2000 and
1000 years ago. These estimates are constrained by ages of tills
found between tephra layers ~600 BCE and ~CE 200 and between
~CE 700 and ~CE 1100, respectively (Yamagata et al., 2002). According
to tephrochronology, the West Ichinsky Glacier advanced
between CE 200 and 600. Two moraines of Koryto Glacier were
formed between CE 700 and 1864, most likely around CE 1000
(Yamagata et al., 2000, 2002) and at Kozelsky glacier one moraine
was deposited shortly before CE 1737 (Solomina et al., 1995).
In many cases the most prominent moraines in Kamchatka are
ice cored and were deposited during, and often towards the end of,
the 19th century. These advances broadly correspond to the
greatest accumulation values at CE 1810e1860 recorded in the
Ushkovsky ice core (Solomina et al., 2007; Sato et al., 2013).
Tephrochronology, lichenometry, tree-ring dating of moraines
and instrumental records of Koryto Glacier in the maritime region
of Kronotsky Peninsula indicate advances at CE 1610, 1720,
1770e1780, 1805e1820, 1830e1845, 1855e1875, 1890e1906,
1912e1926, 1954e1957, 1970e1976, and in the mid-1980s, which
generally correspond to periods of low summer temperature
reconstructed from tree-rings (Solomina and Calkin, 2003; Dolezal
et al., 2014). However, the role of winter precipitation for mass
balance of glaciers in Kamchatka is also very important
(Vinogradov et al., 1985). As it was shown by numerical experiments
(Yamaguchi et al., 2008), the retreat of Koryto Glacier since
the mid-20th century was likely due to decreasing winter
precipitation.
4.9.2. Altay
The collection of 14
C ages on wood at glacier forefields in the
Altay Mountains is very large, though most were not in situ and the
relationship between the wood samples and glacier fluctuations is
difficult to interpret. Agatova et al. (2012) suggested that glacier
advances similar in extent to those of the 17the19th centuries
occurred between 300 BCE and CE 300, although the evidence for
these advances is not strong. Nazarov (personal communication,
see SM Table 1) reports 14
C ages on wood related to advances of
Masshey (CE 433e595) and Mensu (CE 435e767) glaciers roughly
corresponding to the strong “Late Antiquity Little Ice Age” cooling
occurring from CE 536 to around CE 660 in the Altay and the European
Alps (Büntgen et al., 2016).
Wood found above the upper modern tree line of Aktru, Maashey
and Shavla glacier valleys indicates that glaciers in the Altay
were probably in retracted positions between the 8th and 12th
centuries (Agatova et al., 2012).
A large number of trees were damaged by glacier advances
between ~CE 1200 and ~CE 1800, and ages cluster in the following
periods: CE 1200e1260 (3 ages from 3 valleys), CE 1320e1380 (4
ages from 3 valleys), CE 1440e1510 (5 ages from 3 valleys), and CE
1640e1740 (8 ages from 5 valleys) (Nazarov and Agatova, 2008;
Nazarov et al., 2012; Agatova et al., 2012) (see also Table 1 in SM).
Moraines deposited between CE 1640 and 1740 mark the maximum
advance of the past millennium in the Altay region. Since then less
prominent re-advances occurred shortly before CE 1835 (Belukha
region), in the mid and late 19th century and the early 20th century.
A 2367-year-long tree-ring chronology sensitive to early summer
temperature from the Altay region (Myglan et al., 2012a,b;
Büntgen et al., 2016) shows that glacier advances of the middle of
the first millennium and from the 15th to 19th centuries generally
coincide with cool summers (Fig. 2).
4.10. European Alps
The glacier record for the European Alps over the past two
millennia is mainly based on calendar-dated tree-ring series for the
early periods and on historical documents from CE 1600 onwards.
Additionally, for many glaciers there are continuous length measurements
available for the past 130 years.
The first centuries CE were characterized by retracted glacier
termini. Some evidence is available that at least some glaciers (e.g.
Great Aletsch and Steinlimni in the Swiss Alps) were as small as
they were during the late 20th century (Holzhauser et al., 2005;
Joerin et al., 2006). Sediments in glacial meltwater-fed Blanc
Huez (western French Alps) indicate reduced glacier activity in its
catchment (Simonneau et al., 2014) in the first centuries CE.
Archeological artefacts, collected at the Schnidejoch pass, a glacier
pass in the Bernese Alps (Switzerland) that could hardly be crossed
in periods of large ice extent, span much of the first millennium CE.
They show the highest activity during the Roman period and
indicate the relatively intensive use of this crossing (Grosjean et al.,
2007; Hafner, 2012). Glacier advances in the Alps are recorded as
early as CE 270 and 335 (Nicolussi and Patzelt, 2001; Holzhauser
et al., 2005). Ice during the 5th century was as extensive as the
late 20th century for Glacier du Trient (Hormes et al., 2001) and
around CE 410 for Great Aletsch Glacier (Holzhauser et al., 2005).
Significant advances in the 6th century at Great Aletsch, Lower
Grindelwald, Gorner, and Mer de Glace glaciers (Swiss and French
Alps) culminating at the end of that century or in the early 7th
century brought these ice masses to their maxima in the western
and central Alps during the first millennium CE (Holzhauser et al.,
2005; Le Roy et al., 2015). A persistent advance phase during this
period, from the 5th to the 9th century, is also reconstructed for the
Miage Glacier (Italian Alps) (Deline and Orombelli, 2005). However,
evidence for the 6th century advance is less clear in the eastern Alps
(Gepatsch Ferner and Sulden Ferner) and does not indicate extents
comparable to the 19th century (Nicolussi and Patzelt, 2001;
Nicolussi et al., 2006). Sulden Ferner (Italy) was shorter than in
the 19th century between CE 400 and 800. At Unteraar Glacier
(Switzerland), the ice extent during the 8th century was as small as
during the late 20th century (Hormes et al., 2001). An advance
phase during the 9th century is evident at Gepatsch Ferner and
Lower Grindelwald glaciers, but is less well constrained at Great
Aletsch and Gorner glaciers (Nicolussi and Patzelt, 2001;
Holzhauser et al., 2005). At the Sulden Ferner, an advance at CE
835 nearly reached the 17the19th centuries’ margins (Nicolussi
et al., 2006). In the eastern Alps, the CE 835 advance was probably
the most far reaching during the first millennium CE. General
retreat of glacier termini can be deduced for the late 9th to 11th
centuries. Reduced glacier activity at this time is also suggested by
the sediment record of Lake Blanc Huez (Simonneau et al., 2014),
and the youngest artefacts from Schnidejoch indicating that it was
possible to cross the pass at that time (Hafner, 2012).
A subsequent glacier advance occurred in the 12th century, e.g.,
Gepatschferner, Gorner and Mer de Glace (Nicolussi and Patzelt,
2001; Holzhauser et al., 2005; Le Roy et al., 2015). After a short
retreat phase, a general advance is documented in the late 13th
century that culminated between ~CE 1350 and 1385 at Great
Aletsch, Gorner, Mer de Glace, and Vernagtferner (Austria) glaciers
(Holzhauser et al., 2005; Patzelt, 2013; Le Roy et al., 2015). This
advance was less extensive at Gepatschferner and Pasterze (Austrian
Alps) (Nicolussi and Patzelt, 2001). Glaciers retreated by CE
1400 but further advances are recorded later in the 15th
(Gepatschferner, Stein, and Tsidjiore Nouve glaciers; Austrian and
Swiss Alps) and early 16th centuries (Great Aletsch and Tsidjiore
Nouve glaciers) (Nicolussi and Patzelt, 2001; Holzhauser et al.,
2005; Schimmelpfennig et al., 2012, 2014).
The most widespread phase of glacier advances in the Alps (~CE
1600 to 1860) is characterized by a series of advances during which
generally similar extents were reached: 1600, 1640, 1680, 1720,
1775, 1820 and 1855/60 (e.g. Zumbühl et al., 1983; Nicolussi and
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9070
Patzelt, 2001; Grove, 2004; Deline and Orombelli, 2005;
Holzhauser et al., 2005; Nicolussi et al., 2006; Nussbaumer et al.,
2007; Ivy-Ochs et al., 2009; Imhof, 2010; Nussbaumer and Zumbühl,
2012; Schimmelpfennig et al., 2012, 2014; Patzelt, 2013; Le
Roy et al., 2015). However, the number of maxima, the ages of the
advances and the similarity of extents vary and were controlled
mainly by the dynamics of each glacier system.
General retreat began after CE 1860 and was interrupted by
readvances around 1890, 1920 and 1980. However, some glaciers,
e.g., Great Aletsch Glacier (Holzhauser et al., 2005), have retreated
continuously since their mid-19th century maxima (Zemp et al.,
2008).
Mass balance variability in the Alps mainly correlates with
summer temperature, while winter precipitation usually plays a
less important role (e.g. Vincent et al., 2005; Steiner et al., 2008). A
summer temperature reconstruction (Büntgen et al., 2011) based
on tree-ring-width data from the eastern Alps (Nicolussi et al.,
2009) is in very good agreement with the high-resolution glacier
records supporting summer temperature as the primary control on
glacier fluctuations.
4.11. Caucasus
There is no evidence of glacier variations in the first millennium
in the Caucasus. Glaciers were probably small in the beginning of
the past millennium, but the available evidence is only indirect
(archeological, biostratigraphic and pedological data) (Solomina
et al., 2016). Serebryanny et al. (1984) reported a minimum 14
C
age for a moraine in the Bezengi Valley (CE 1245e1428). However,
taking into consideration the location of this moraine relative to the
modern glacier (6 km) the age obtained from a peat bog within the
moraine complex is probably too young to closely constrain the age
of the moraine. At Bol’shoy Azau Glacier a radiocarbon age on a soil
horizon buried between two glacial units dates to CE 1465e1642
(Baume and Marcinek, 1998). Trees growing on this surface are
more than 400 years old, suggesting that the age of the upper
glacial unit and of the corresponding glacial advance is at least CE
1600.
According to tree-ring minimum ages, Bol’shoy Azau Glacier
advanced prior to CE 1598 and Kashkatash Glacier before CE 1600
(both are located in the Elbrus area). A younger advance of both
glaciers in the mid-19th century reached almost the same extent
(Solomina et al., 2016). Two other major advance phases in the
northern Caucasus date back to the second half of 17th century and
the first half of the 19th century (minimum tree-ring ages at Tsei
Glacier) (Solomina et al., 2016). General glacier retreat started in
the late 1840s. Four to five minor readvances occurred in the period
between CE 1860s and 1880s and three readvances or still stands
took place in 1910s, 1920s and 1970se1980s. A reconstruction of
June-September temperature in the central Caucasus based on blue
intensity (equivalent of maximum density) from conifers spans the
period CE 1569e2008 (Dolgova, 2016), and the glacial advance of
the 1840s coincides with a reconstructed cold period (CE
1825e1867).
4.12. Central Asia
4.12.1. Central Asia, semi-arid area
Compiling and evaluating 10
Be ages, Dortch et al. (2013) recognized
late Holocene advances in the semi-arid western end of the
Himalayan-Tibetan orogen at CE 400 ± 300 and CE 1600 ± 100. At
Abamova Glacier (Pamir-Alai Mountains) a radiocarbon age from a
shallow soil horizon below a fresh till, CE 903-1188, indicates a
glacier advance shortly after this time (Zech et al., 2000). Wood
found in a buried soil horizon at Raigorodskogo Glacier (Pamir-Alai
Mountains) provides evidence of a warm period and glacier retreat
shortly before CE 255e527 when the glacier expanded (Narama,
2002). The youngest advance is dated by a 14
C age (CE
1521e1646) on a trunk of Juniperus turkestanica broken by a glacier
advance. These two advances were of similar magnitude, although
the former was slightly larger. Historically, advances of Raigorodskogo
Glacier are recorded back to CE 1908e1934, and CE
1960e1977 (Narama, 2002). In the Urumchi valley (Tien Shan)
samples of whewellite coating on glacial boulders from the outermost
moraines indicate that the glacier retreated from its
maximum extent before CE 1535 ± 120.
4.12.2. Central Asia, monsoon area
R€othlisberger and Geyh (1985) undertook some of the first
studies using radiocarbon ages from glaciers in Pakistan, India, and
Nepal to show that glaciers advanced at CE 250e550, 650e1050,
1150e1400 and 1450e1850. Using 10
Be, Murari et al. (2014) identified
three advances in the monsoon-influenced Himalaya-Tibetan
area at CE 500 ± 200,1300 ± 100 and 1600 ± 100. For the same area,
Yang et al. (2008) established a chronology for the past two
millennia based on multi-proxy data and recognized three main
periods of glacier advance at around CE 200e600, 800e1150, and
1400e1920, with the most widespread glacier advance occurring at
about CE 400e600 (Fig. 2). They suggested that these advances
were synchronous with those in the southern Himalaya during CE
380e600, 870e1100, 1400e1430, and 1550e1850, and that the
glacier advance around CE 1000 also occurred in the central
Himalaya. Xu and Yi (2014) reported that glaciers retreated from
their maxima during the 16th to early 18th, late 14th to early 15th
and early 16th centuries in the southern, northwestern, and
northeastern regions, respectively. In addition, they suggested that
the periods of glacier advance during the late 18th to early 19th
centuries and the retreat period of the late-19th century are common
throughout Tibet.
Dendrochronological dating provides higher-precision ages for
glacier advances in Tibet, which occurred before the 15th century,
before CE 1662, between 1746 and 1785 (maximum extent), early
19th, late 19th to early 20th, and during the mid-20th centuries
(Br€auning, 2006; Zhu et al., 2013; Xu et al., 2012; Hochreuther et al.,
2015; Loibl et al., 2015). On a centennial (Yang et al., 2003, 2008; Xu
and Yi, 2014) and decadal (Wang et al., 2015) timescale, temperature
changes are the main controlling factor for glacier fluctuations
rather than precipitation changes caused by variations in the south
Asian summer monsoon. Tree-ring studies for the western Himalaya
show that the 18th and 19th centuries were the coldest interval
of the past millennium coinciding with the expansion of glaciers in
the western Himalaya (Yadav et al., 2011).
The DELA for Zepu Glacier (S-E Tibetan Plateau) estimated from
moraines deposited at CE 200e600, 800e1150, 1400e1650, and the
19th century was 160 m, 110 m, 60 m, and 10 m, respectively, which
is equivalent to temperature lowering of 1.0 C, 0.7 C, 0.4 C and
0.1 C compared to 1989 (Yang et al., 2008). In most areas of central
Asia, glaciers began and continue to retreat since the beginning of
the 20th century (Mayewski and Jeschke, 1979; Mayewski et al.,
1980; Kutuzov and Shahgedanova, 2009; Owen and Dortch, 2014).
4.13. Low latitudes
Ninety-five percent of tropical glaciers are located in the Andes.
A collection of 14
C ages constraining mostly minimum or maximum
ages of glacier advances in the tropics of South America is available
(e.g. Mercer and Palacios, 1977; Clapperton, 1983; Goodman et al.,
2001; Rodbell et al., 2009), though only recent development of
TCN dating made it possible to describe the late Holocene history of
some glaciers (mostly in the outer tropics) in sufficient detail
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(Licciardi et al., 2009; Hall et al., 2009; Jomelli et al., 2011, 2014;
Stroup et al., 2014, 2015 e recalculated in this study using the
regional mean production rate from Martin et al., 2015 and Lal/
Stone scaling scheme). Lake-sediment studies (Abbott et al., 2000;
Polissar et al., 2006; Rodbell et al., 2008; Stansell et al., 2014)
provided a continuous context useful for the interpretation of the
moraine ages.
In the outer tropics accumulation occurs only during the wet
period, whereas in the inner tropics it lasts the whole year (Kaser
and Osmaston, 2001). Here we consider the two regions separately.
4.13.1. Inner tropics
According to lake-sediment records from the Venezuelan Andes
(Polissar et al., 2006), glaciers in the Mucubají watershed were
absent between CE 500 and 1100 and since the 1820s, but four
glacial advances occurred in the second millennium (CE
1180e1350, 1450e1590, 1640e1730, and 1800e1820). In the same
region Stansell et al. (2014) documented glacier changes in two
other catchments. At low elevation (<4400 m asl), lake sediments
do not show any signal of glacier advance during the past centuries
while the increase in clastic sedimentation in a catchment located
at a higher elevation (4750 m) indicate increased glacial activity
during the period between CE 1350 to 1850.
On the southern slope of the Cordillera del Cocuy (Colombia)
two moraines of Ritacuba Negro Glacier were dated by TCN to CE
900 ± 90 and CE 1750 ± 30 (recalculated from Jomelli et al., 2014).
Using historical data, Hastenrath (1981) demonstrated that glaciation
in the Ecuadorian Andes, near Quito, was more extensive at
the time of the Spanish conquest in the 16th century than in the
late 20th century.
In the inner African tropics Karlen et al. (1999) identified periods
of glacier advances at Mount Kenya from lake sediments for CE
400e750, around 1350 and around 1550, while retreat phases
occurred between 150 BCE and CE 400 and in CE 750e850. In the
1890s the glaciers on Mt Kenya were still close to their prominent
moraines. Comparing the glacier records with the paleoclimatic
information Karlen et al. (1999) attributed the glacier advances
primarily to changes in temperature.
4.13.2. Outer tropics
Rodbell (1992) identified the culmination of clastic sediment
flux in Laguna Queshque, Cordillera Blanca, at ~CE 500 corresponding
to the 14
C date of a moraine deposited in a nearby cirque
in CE ~450 ± 190. Abbot et al. (2000) demonstrated that the glacial
activity in Lago Taypi Chaka Kkota, Cordillera Real, Bolivia, was
gradually increasing from ~300 BCE and reached the modern level
around ~ CE 500e600. Stansell et al. (2013) reported that clastic
sediment flux was low in the deposits of proglacial lakes in
southern Peru from ~4.0 ka to 14th century CE, indicating that the
glaciers were generally in retreated positions.
At some glaciers the peak of glacial activity occurred in the late
17th through 18th century. In Cordillera Real, Bolivia, the TCN
samples from Telata Glacier show an advance at CE 1760 ± 130
(Jomelli et al., 2011). Using the TCN Hall et al. (2009) dated an
advance of Miticocha Glacier to CE 1740 ± 30 in Cordillera Huayhuash.
Licciardi et al. (2009) identifed the moraines deposited
around the same time at three glaciers (Salcantay e CE 1680 ± 70,
Tucarhuay e CE 1710 ± 50, and Sisaypampa e CE 1710 ± 40).
The most robust reconstruction of glacier and climate variations
in the outer tropics comes from the Quelccaya Ice Cap (Peru). Here
there is evidence of earlier glacier advances. Five successive moraines
at Qori Kalis Glacier were formed at CE 1500 ± 20, 1635 ± 20,
1685 ± 20, 1705 ± 10, and 1795 ± 10 (recalculated from Stroup et al.,
2014, 2015). Comparing the time of Qori Kalis Glacier advances and
the Quelccaya Summit Dome ice core records (Thompson et al.,
2013) Stroup et al. (2014) deduced that temperature rather than
accumulation was the main driver of glacier changes. In the 20th
century tropical glaciers were retreating at faster rates than in other
mountain regions of the World (Rabatel et al., 2013a).
4.14. Dry Andes
Extremely limited information on glacier fluctuations during the
past two millennia is available in the Dry Andes between 17 and
36 S. No ages are available for the first millennium, but 14
C-dated
peat samples and paleosols date maximum expansion of the past
millennium to the 16the17th centuries (Espizua and Pitte, 2009;
Masiokas et al., 2009a; SM Table 1 and Fig. 2). Additional limited
evidence indicates subsequent readvances during the 19th century.
During the general retreat of the past decades, numerous frontal
readvances and surging events have been identified, mostly in the
past 30e40 years (Masiokas et al., 2009a).
4.15. Southern South America (Patagonia, Tierra del Fuego)
In the north Patagonian Andes (~36e45S) no ages are available
for glacier advances during the first millennium. Maximum
extents of the past millennium occurred between the 16th and 19th
centuries as indicated by tree-ring-dated moraines and 14
C dating
of in situ stumps affected by advancing glaciers (e.g. Villalba et al.,
1990; Masiokas et al., 2009a; Ruiz et al., 2012; see SM Table 1).
Further south at the latitudes of the northern Patagonian Icefield
(NPI; 46e47S), in situ roots of trees killed by the advancing
Soler Glacier were 14
C dated to between the 13th and 14th centuries
and provide one of the best indicators of early glacier activity in this
region (Glasser et al., 2002). Tree-ring and lichenometric dating of
trimlines and moraines of several NPI outlet glaciers indicate that
they probably reached their maxima during the 19th century
(Winchester and Harrison, 1996; Harrison and Winchester, 1998,
2000; Winchester et al., 2001; Harrison et al., 2007, 2008; Glasser
et al., 2011).
Radiocarbon dating of in situ and reworked stumps and 10
Be
dating of boulders on moraine crests from several sites around the
southern Patagonian Icefield (SPI; 49e51S) show a period of
glacier advances between the 4th and 9th centuries (e.g. Mercer,
1965; R€othlisberger, 1986; Aniya, 2013; Strelin et al., 2014).
Numerous in situ trees killed by Ameghino Glacier (an eastern
outlet of the SPI) and Glaciar Jorge Montt (northern tip of the SPI)
during its main advance between the 15th and 17th centuries were
up to 300 years old indicating smaller glaciers and likely warmer
conditions; however, the size of the glaciers at that time is not
known (Masiokas et al., 2009a). There is, however, a growing body
of evidence for glacial advances between the 13th and 15th centuries
(e.g. Mercer, 1970; R€othlisberger, 1986; Strelin et al., 2014). In
this portion of the Andes the maximum expansion has been dated
between the 17th and 19th centuries largely based on dendrochronological
determinations and 14
C dating of in situ stumps and
paleosoils (e.g. Mercer,1965,1970; R€othlisberger,1986; Aniya,1995,
2013; Aravena, 2007; Masiokas et al., 2009b; Aniya and Skvarca,
2012; Strelin et al., 2014; Glasser et al., 2011). Despite several
readvances identified in the past century, glacier recession has been
clearly evident both in northern and southern Patagonia and it has
apparently accelerated in recent decades (Rignot et al., 2003;
Masiokas et al., 2009a; Glasser et al., 2011).
The evidence in the Magallanes-Tierra del Fuego region indicates
a similar pattern to that observed around the SPI, with
glacier advances around the 8th century, around the 14th century,
and maxima between the 17th and 19th centuries (Kuylenstierna
et al., 1996; Koch and Kilian, 2005; Aravena, 2007; Strelin and
Iturraspe, 2007; Strelin et al., 2008). The earliest ages are based
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9072
on 14
C dating of reworked stumps whereas the later events are
largely based on tree-ring dating of deposits. Glacier readvances
during the 19th and 20th centuries have also been identified and
dated using tree rings and direct observations (Kuylenstierna et al.,
1996; Koch and Kilian, 2005; Aravena, 2007; Strelin et al., 2008).
Glaciers in the Cordillera Darwin show a contrasting behaviour in
the late 20th century, with glaciers reaching positions at or close to
their Holocene maxima in the southern and western flanks and
showing marked frontal retreats on the northern and eastern flanks
of the Cordillera (Holmlund and Fuenzalida, 1995; Porter and
Santana, 2003). Glaciers in the Cordillera Fueguina Oriental (at
almost 55S) probably reached their past millennium maximum
between the 16th and 17th centuries and receded slowly
until ~ 60e70 years ago, when the rate of glacier retreat accelerated
(Strelin and Iturraspe, 2007).
Moreno et al. (2014) reconstructed centennial changes in
southern South American climate over the past three millennia
based on a stratigraphic record collected around 51S. This reconstruction
is particularly relevant because it reproduces well the
patterns of the Southern Annular Mode (SAM), the main driver of
atmospheric variability in the mid- to high-latitudes of the
Southern Hemisphere. In their reconstruction, Moreno et al. (2014)
identified two periods of cold/wet conditions in southern Patagonia
peaking at around CE 250e350 and CE 550e850. Although the
evidence for glacier advances in this region is limited for the first
millennium, these cold/wet intervals seem to correspond with the
few events identified during this period in Patagonia and Tierra del
Fuego (Fig. 2). Glasser et al. (2004) claimed that the advance of Soler
in CE 1220e1340 and of four other glaciers in southern Patagoina
corresponds to the increase of winter precipitation. However, the
period of extensive glacier advances between the 13th and 19th
centuries coincides with an extended cold/wet interval. This
extended period of glacier advances is also correlative with a longterm
cool/wet interval identified in marine sediment records at
around 44S along the north Patagonian coast (Sepulveda et al.,
2009). Several studies have noted a general warming in southern
South America during the 20th century (Villalba et al., 2003;
Masiokas et al., 2008; Neukom et al., 2010), and Moreno et al.
(2014) corroborate these previous studies showing a dramatic
change towards warm/dry conditions starting in the late 19th and
early 20th centuries (Fig. 2).
4.16. New Zealand
Recent glacier recession in the Southern Alps of New Zealand
has led to the exposure of soils and wood buried within steep
lateral-moraine walls. Radiocarbon ages on organics associated
with these soils afford age constraints on underlying and/or overlaying
tills (Burrows,1973; R€othlisberger,1986 ; Gellatly et al.,1988;
Grove, 2004) (see SM Table 1). According to these data the most
extensive advances in New Zealand occurred around 1000 and 600
years ago and in the 19th century (Gellatly et al., 1988). In addition,
Wardle (1973) counted rings of trees living on moraine ridges to
place minimum-limiting ages on late Holocene moraine deposition
by glaciers draining the western flank of the Southern Alps. He
determined that 18 Westland glaciers constructed moraines during,
or prior to, the mid- to late-18th, the early- to mid-19th, and the
late 19th centuries.
Recently, longer and more detailed moraine chronologies from
four glacier systems have been constructed using 10
Be surfaceexposure
dating at Mueller, Tasman, Hooker Glaciers near Aoraki/
Mount Cook (Schaefer et al., 2009) and Cameron Glacier in the
Arrowsmith Range (Putnam et al., 2012) in the central Southern
Alps. These results compare remarkably well with previously obtained
and recalibrated radiocarbon ages (arithmetic means with
1s uncertainty from the 14
C ages, when referenced to the same year
as exposure ages).
At Mueller Glacier, the outermost late Holocene lateral moraine
was deposited in 1410 ± 150 BCE, and inboard moraines were
constructed at 380 ± 100 BCE, 135 ± 100 BCE, CE 150 ± 100, CE
1390 ± 80. Results reported here are those of Schaefer et al. (2009)
as recalculated by Putnam et al. (2012) using the local New Zealand
production rate of Putnam et al. (2010). The most prominent
moraine of the past millennium was constructed by CE
1540 ± 80 yrs, with moraines being constructed at positions slightly
inboard at CE 1715 ± 60, CE 1754 ± 8, and CE 1824 ± 31.
At Hooker Glacier, the sequence of lateral-moraine ridges were
constructed by CE 430 ± 220 and CE 860 ± 100. At Tasman Glacier,
the distal moraine was deposited CE 146 ± 100, whereas a proximal
moraine was formed contemporaneously with the one at the
Hooker Glacier ~ CE 860 (Schaefer et al., 2009; Putnam et al., 2012).
At Cameron Glacier moraines deposited during the past two
millennia are fewer, with moraines dated to CE 1298 ± 27 and CE
1427 ± 61 that are adjacent to much older (4940 BCE ± 190)
moraine surfaces. Progressive retreat of the glacier terminus since
the 15th century was interrupted by stillstands or minor readvances
around CE 1770, 1864, and 1930 (Putnam et al., 2012).
Since the 1860s, the glaciers of New Zealand have been
retreating with minor readvances in 1888, 1908e1909, early 1920s,
1947e1950, 1965e1967 and 1980e2005. There has been net
recession since the 1940s (Purdie et al., 2014); however, a period of
stable mass balance and glacier readvance/stillstand registered
widely across Southern Alps glaciers interrupted this trend between
CE 1980 and CE 2005 (Chinn et al., 2005; Purdie et al., 2014).
This general pause in ice retreat and snowline rise has been suggested
to relate to a switch in the phase of the Interdecadal Pacific
Oscillation (IPO) that occurred around CE 1978, which involved an
increase in the proportion of cold air-mass incursions to the
Southern Alps from southerly quarters during the ablation season
(Tyson et al., 1997). Atmospheric warming and renewed glacier
retreat beginning shortly after CE 2005 may relate to the most
recent switch of the IPO (Purdie et al., 2014). Currently, less than a
third of the maximum ice volume from around CE 1850 remains
(Chinn et al., 2012).
The magnitude of advances in the past two millennia in the
Southern Alps of New Zealand varied from glacier to glacier, but in
general it was decreasing over the past millennium. However, all
moraines formed since the mid Holocene are located very close to
one other and to the moraines of 19th century. Due to this similarity
in magnitude of glacial advances some moraines were destroyed
and overlain by later advances, leading to different compositions of
moraine complexes of individual glaciers.
5. Discussion
5.1. What was the magnitude of glacier fluctuations in the past two
millennia and is it comparable in scale with contemporary glacier
retreat?
Determining the size of former glaciers that were less extensive
than later glacial advances is a challenge due to the paucity of these
records. However, there is much evidence that during certain periods
of the first millennium glaciers were smaller (or of equal
sizes) than they were at the end of the 20th to early 21st centuries
in many regions including the Arctic (e.g. Iceland, Spitsbergen,
Alaska), temperate zone (e.g. Scandinavia, Altay, Alps, Tibet) and
the tropics (e.g. Peru, Bolivia) (see references in Table 1). The
extremely rare occurrence of moraines deposited in the first millennium
in the Southern Hemisphere and in the tropics also indirectly
indicates that climate conditions were less favourable for
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e90 73
glaciers than over most of the second millennium excluding the
past one to two centuries. In some regions such as southern Alaska,
the European Alps and the Altay Mountains, the glaciers also
retreated from approximately the 8th to the 13th centuries, but the
magnitude of their retreat during the MCA is unknown. We found
only one clear indication (Røthe et al., 2015) of glacier retreat to
close or within their present limit after the 13th century (Spitsbergen
around CE 1400 and between CE 1720 and 1810) (see Fig. 3).
The time of the most prominent advances in the past two
millennia varies not only from region to region, but also from
glacier to glacier. As described in section 4 (see references
therein) in the Alps the maximum extent of the glaciers was
reached in the 17the19th centuries, but advances in the 6th
century in the Swiss and French Alps and in the 9th century in
the eastern Alps were of nearly similar magnitude. A maximum
glacier advance also occurred in the Alps in the 14th century. In
Alaska outboard moraines were deposited in the 1800s, but there
are also places where they were formed in the first millennium or
about CE 1300. In the high latitudes the maximum glacier
advance often occurred quite late (in the 19th to early 20th
centuries), although recent discoveries indicate advances of
similar magnitude dating to the first millennium (e.g. in Spitsbergen
in CE 350 ± 90; in Greenland in ~ CE 300 and in ~ CE 490).
Maxima of glacial advances in the 17th to 18th centuries occurred
in the Swedish Lapland, the Altay Mountains and in some regions
of the Canadian Cordillera; in the 18th to early 20th century they
occurred in Iceland, the Canadian Rockies, and southern Patagonia.
In central Asia and in the tropics, many glaciers reached their
maxima in the 15the16th centuries. Major advances around ~ CE
1000 occurred in Franz Josef Land and at some glaciers in central
Asia. Mueller Glacier in New Zealand reached its maximum at the
beginning of the first millennium, while the maximum advance of
Hooker Glacier occurred a few centuries later in the ~9th century,
and Cameron Glacier reached its maximum extent even later in
the 13the15th centuries.
During the maximum advances of the past millennium the ELA
was lower than at the end of the 20th century in the Alps (Grove,
2004), Norway (Nesje, 2009), the Caucasus, the Altay, Kamchatka
(Solomina, 1999), and in the central Andes of Argentina (Espizua
and Pitte, 2009) by 100e150 m. In New Zealand the DELA from
the maximum last millennium extent is about 140 m (Putnam et al.,
2012). The DELA was much smaller (10e20 m) in arid central Asia,
while in the tropics of the Peruvian Andes it reached as low as
300 m (Jomelli et al., 2011) and up to 300e500 m in Venezuela
(Polissar et al., 2006). On Baffin Island the magnitude of the
snowline variations in the past two millennia is estimated to be up
to 400 m (difference between the beginning of CE and the 19th
century) (Miller et al., 2013). Unfortunately for most regions we
cannot assess the full range of ELA variations over the past two
millennia primarily due to uncertainties related to the magnitude
of glacier retreat during warm periods.
Globally, the number of advances recorded before the 11th
century is much smaller than in the second millennium, but it is
still extensive enough to state that in many regions glacier advances
occurred during the first millennium as well, and whereas
later advances did not destroy some of these moraines, the
magnitude of these first millennium advances was rather large and
most likely comparable with those of the second millennium.
Table 1
Periods of glacier retreats (glaciers as extensive as in the end of 20th-early 21st centuries or smaller).
Region, glacier Ages, CE Reference point Reference
Coastal and southern Alaska 950 Close to their modern margins Wiles et al., 2008; Barclay et al., 2009a,b
Southern Alaska 10thÀ13th centuries Less extensive than during the late 20th
century
Wiles et al., 2008
Western Cordillera Before 500 Equal or smaller than in the late 20th
century
Allen and Smith, 2007; Clague et al.,
2010
British Columbia (Llewellyn Glacier) Before 300e500 No more extensive than today Clague et al., 2010
SE Greenland (Kangerlussuaq) Until 200e250 Near the modern margins Young and Briner, 2015
West Greenland (Jakobshavn Isbræ) Until 200e250 Near the modern margins Lloyd, 2006
NW Greenland (Upernavik Isstrøm) Until ~900e1200 Less extensive than now Briner et al., 2013
SW Greenland (Kiagtut Sermia) 660 ± 150 Within or near its present extent since Winsor et al., 2014
Iceland First millennium Smaller than in the past millennium
and on some occasions smaller than
those of today
Geirsdottir et al., 2009
Spitsbergen (Kongressvatnet) Before 8th century, probably up to
before 1370e1400
Smaller than today (the presence of a
glacier signal is not detected in the
sediments)
Guilizzoni et al., 2006
Spitsbergen (Longyearbreen) 20e820 2 km shorter than in early 20th century Humlum et al., 2005
Spitsbergen 2nde6th centuries Smaller than during the past
millennium
Reusche et al., 2014
Spitsbergen (Karlbreen) Before 50, 550e750, 800e1100, ~1400,
between 1720s and 1810s
Close to or beyond its present limit Røthe et al., 2015
Norway (Juvfonne ice patch) 250e530 and 800e900 Strong glacier retreat Nesje et al., 2012
Franz Josef Land (Cape Lagerny,
Northbrook Island)
Until at least ~550e670 Beyond the present margin Lubinski et al., 1999
Altay 8the12th centuries Retreat of unknown magnitude Agatova et al., 2012
Alps Before 270, in the 5th and in the 8th
centuries
As small as in the late 20th century Holzhauser et al., 2005; Hormes et al.,
2001; Joerin et al., 2006
Alps Late 9th to 11th centuries In retreated positions relative to the
17the19th centuries
Holzhauser et al., 2005; Simonneau
et al., 2014
Alps (Unteraar, Steinlimi and Lago di
Musella glaciers)
Until 780 Less extensive than present Hormes et al., 2006
Swiss Alps (Glacier du Trient) 5th century As extensive as the late 20th century Hormes et al., 2001
Tropical Andes (southern dome of the
Quelccaya ice cap)
Until 344e539 Smaller than today Mercer and Palacios, 1977
Cordillera Real, Bolivia (Lago Taypi
Chaka Kkota)
Until ~500e600 Smaller than today Abbott et al., 2000
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9074
5.2. Ablation-season temperature or accumulation-season
precipitation? Climate signal in glacial records of the past 2
millennia
Glacier mass-balance measurements over the historical period
demonstrate that for most regions summer temperature is the
dominant control on annual mass balance (Koerner, 2005;
Bj€ornsson et al., 2013). Some exceptions when glacier advances
were forced primarily by high winter precipitation are recorded in
coastal areas of Scandinavia, SE Alaska, Kamchatka and New Zealand
(e.g. Lemke et al., 2007). Similarly, over the past two millennia
for those regions where summer temperature or both summer
temperature and winter precipitation reconstructions are available,
and the accuracy of the glacier records allows a resolution of a few
decades, lower summer temperatures coincide with many of the
recorded ice advances (e.g. Alaska: Wiles et al., 2004, 2008; Barclay
et al., 2013; Anchukaitis et al., 2013; Canadian Rockies: Luckman,
2000; Alps: Holzhauser et al., 2005; McCormick et al., 2012;
Lüthi, 2014; Le Roy et al., 2015; Monsoon Asia: Yang et al., 2008;
New Zealand: Schaefer et al., 2009; South America: Neukom et al.,
2010; Leclercq et al., 2012; Tropics: Jomelli et al., 2011; Malone
et al., 2015) (see also Fig. 2). In a few cases increased snowfall
was suggested as the primary driver of former advances, e.g., in
western Norway (Nesje et al., 2007; Bakke et al., 2008; Rasmussen
et al., 2010) in the Canadian Rockies (Luckman and Villalba, 2001),
during the MCA (Koch and Clague, 2011), in the Brooks Range ~CE
1600 (Boldt et al., 2015), and in the central Himalaya in the 19th
century (Yang et al., 2008).
At a more general level there is some agreement between the
temperature reconstructions (PAGES 2k Consortium, 2013) for the
Arctic and Europe, and with glacier fluctuations in the high and
mid latitudes (Fig. 4). This suggests that maximum advances
correspond to cool periods. Glacier advances dominated from the
mid 13th century, consistent with the transition to cooler summer
temperatures (Fig. 4). The advances to the past millennium
maximum glacier extents occurred in the 17the18th and 19th
centuries and generally correspond to the most prominent cooling,
at least in the mid and high latitudes of the Northern Hemisphere
as reconstructed by PAGES 2k Consortium (2013). The number of
glacier advances fits well with the Northern Hemisphere temperature
reconstruction based on high- and low-frequency proxy
data (Moberg et al., 2005), especially for the period since the 8th
century (Fig. 5 and 6). The number of recorded glacier advances
continued to increase after the peak of cooling in the 17th century
partly due to the well-preserved recessional moraines deposited
after the most prominent advances in the 17th, 18th and 19th
centuries.
Modelling based on current mass-balance observations and
paleoclimatic proxies was used to estimate the climate conditions
responsible for the advances over the past two millennia in several
regions. Regional estimates of ELA changes in the Alps between the
maximum LIA extent in the mid-19th century and ~1970, based on
ablation-to-accumulation area calculations, suggests an ELA rise of
~69 m and ~94 m for the Swiss (Maisch et al., 2000) and Austrian
Alps (Patzelt and Aellen, 1990), respectively. A modelled ELA rise of
75 m calculated on the bases of the observed increase in summer
temperature between 1850 and 1971/90 of 0.6 C meets these
values (Zemp et al., 2007). A further rise of the ELA between 1984
and 2010 by about 170 m was estimated for a group of glaciers in
Fig. 3. Selected time series for glacier size spanning the past two millennia. 1. Spitsbergen.
Reconstructed ELA (5 pt. running mean) from Karlbreen (Røthe et al., 2015). 2.
Alaska. Time-distance diagram for Sheridan Glacier. Horizontal green lines show intervals
of forest growth on forefields (Barclay et al., 2013). 3. Western Canada.
Compilation of glacier extents (Koch, this study). 4. Iceland.. Sediment flux into glacial
lake Hvítarvatn e proxy for glacial activity (Larsen et al., 2013). 5. Southern Scandinavia.
Reconstructed ELA (compilation from Nesje et al., 2001;Bakke et al., 2005a,b;
Vasskog et al., 2012). 6. The Alps. Fluctuations of Great Aletsch and Lower Grindelwald
Glacier (Holzhauser et al., 2005), and Mer de Glace Glacier (Le Roy et al., 2015). 7. Altay,
southern Sibiria. Glacier fluctuations based on 14
C and tree-ring dating of wood buried
by glacial deposits (Nazarov et al., 2012, modified). 8. Glacier fluctuations in western
Tibet. Compilation from tree-ring, optically stimulated luminescence and 14
C dating of
moraines (from Yang et al., 2003, 2008; R€othlisberger and Geyh, 1985, Br€auning, 2006;
Zhu et al., 2013; Xu et al., 2012; Hochreuther et al., 2015; Loibl et al., 2015). 9. Magnetic
susceptibility from fjord sediments, Golfo Elefantes, as a proxy for Gualas Glacier
fluctuations, northern Patagonia, Chile (Bertrand et al., 2012). 10. Glacier fluctuations in
New Zealand based on 14
C and 10
Be ages (Schaefer et al., 2009; Putnam et al., 2012,
modified). The colors in the background highlight the well-dated advances (orange)
and retreat (gray) during the MCA and early 1st millennium in the Alps to enable the
inter-regional comparison. Dashed horizontal lines represent the modern ELA or front
positions in the end of 20th-early 21st centuries.
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e90 75
the western Alps (Rabatel et al., 2013a,b). This ELA increase in the
Alps since the mid-19th century was mainly a result of summer
climate change, i.e., mass-balance reconstructions proved no
significant trend for winter balances (Huss et al., 2008). Statistical
analyses of mass-balance series from the Alps covering the past ~60
years also prove the dominance of summer climate variability, i.e.,
Fig. 4. Glacial extent: yellow e glacier(s) smaller than now (end of 20th-early 21st centuries); pale yellow e glacier(s) smaller than maximum extent, but size is generally not well
known (included here only if this status is clearly indicated in the original publications, mostly based on the ages of wood incorporated into till or overrun by glaciers); light green e
glaciers present in the watersheds (only for lake-sediment records); light blue e glacier(s) advancing or expanded; dark blue e glacier(s) close to or at their maximum extent of the
past 2 ka. Although a few glaciers experienced small-scale, intermittent advances during recent decades they were too minor to represent on this summary diagram. The empty cells
indicate the absence of information on the glacier status. The descriptions of all series are listed in SM Table 1. The sequence of the series corresponds to the sequence of the records
in the SM Table 1. Temperature reconstructions gridded at 50 years (50 year mean) (PAGES 2k Consortium, 2013).
Fig. 5. Stacked time series of the number of advances for each region documented in this review. The summary is based on 275 time series (Table 1 SM), with age control based on
of the 14
C, TCNs, tree rings, and historical dates for ages of moraines, and periods of glacial activity recorded in lake and marine sediments settings, and multi-proxy reconstructions.
Considering the heterogeneity of evidence used to infer former glacier extent, and because what constitutes an “advance” differs among authors, the number of advances (y-axis)
should be regarded as approximate only. Nonetheless, several clear trends emerge. The series are of different nature and their number cannot be considered as numerical records.
They show the number of cases when the signal of certain glacial activity is preserved either as geomorphic features or as sediment layers. A small number of glacier advances
occurred in many regions between the 1st and 12th centuries CE, including some that were close in magnitude to those occurring later in the second millennium CE, although most
advances were probably of smaller extent than during the 2nd millennium and their traces are not preserved (see Section 4 “Regional description”). An abrupt increase of the
number of recorded advances starting in the 13th century and lasting through to the 19th century indicates both widespread glacier growth and greater preservation because these
moraines have not been obliterated by larger advances, i.e., the advances of this time were of larger magnitude. After these maxima glaciers started to recede with readvances of
smaller extent therefore the number of records after the maximum continue to increase (generally in the 19th to early 20th centuries).
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9076
summer temperature and related variables, on the evolution of
glaciers (e.g. Sch€oner and B€ohm, 2007).
In the tropical Andes, Rabatel et al. (2008) suggested that the
glacier maximum in the 17th century may have been due to a
20e30% increase in precipitation and a 1.1e1.2 C decrease in
temperature compared with the present conditions. Malone et al.
(2015) demonstrated that a 1 C cooling compared to today
(without changes in precipitation) is necessary to explain Qori
Kalis’ maximum extent at CE 1500 ± 20 (reaculated from Stroup
et al., 2014). In Venezuela, Polissar et al. (2006) determined that
a 300e500 m lowering of the ELA occurring between CE 1250 and
CE 1820 is equivalent to a 2.6e4.3 C drop in temperature,
assuming constant annual precipitation. By combining these results
with a biome migration index, they advocated for a
3.2 ± 1.4 C cooling and a 22% increase of precipitation. Leclercq
et al. (2012) calculated that the retreat of Frias Glacier in northern
Patagonia in CE 1639e2009 could be best explained by a 1.2 C
warming.
Finally, it has to be stressed that glaciers react in a complex
manner to both temperature and precipitation, and in many cases
the data are still too scarce to assess the respective role of temperature
and precipitation changes on glacier fluctuations. The
relative contribution of each parameter to mass-balance changes is
specific for each glacier.
5.3. What were the periods of major glacier advances and retreats
in key mountain regions over the last 2 millennia, and how
synchronous were they regionally and globally?
Below we summarize the information on glacier advances
from regional reviews as well as from the best-dated individual
series of glacier variations in chronological order (see also
Table 2). Along with these data we provide information on
possible decadal-scale forcings (volcanic and solar) for global
coolings that might have induced glacier advances. Please note
that we report here, as well as in the regional chapters (Fig. 1 SM),
the recalculated 10
Be ages from original publications by Schaefer
et al. (2009), Jomelli et al. (2014), and Stroup et al. (2014) (see
details in Table 1 SM).
5.3.1. CE 1e200s
The evidence of glacier advances in the first two centuries CE is
rare. This time is known in Europe as the warm Roman period
(Wanner et al., 2008). Glaciers were smaller or close to the late 20th
century extent not only in Europe (the Alps and southern Scandinavia),
but also in Altay, Tibet, eastern Canada and Canadian Arctic
(see details in section 4). The forelands of some glaciers were
covered by forest (e.g. Llewellyn Glacier in British Columbia; Clague
et al., 2010). In contrast an advance between 250 BCE and CE 150
Fig. 6. Number of glacial advances, temperature reconstruction and climatic forcings. (A) Total solar irradiance anomalies (Steinhilber et al., 2009). (B) Reconstructed global volcanic
aerosol forcing from sulfate composite records from tropical (bipolar) eruptions (Sigl et al., 2015). (C) Multi-proxy temperature reconstruction of the Northern Hemisphere (CE
1e1979 (red) with its 80-yr component (blue) and the instrumental record (green)) (Moberg et al., 2005). (D) Number of glacial advances normalized and inverted (avereged for the
high and mid latitudes of the Northern and Southern Hemispheres and from the tropics). See also caption of Fig. 5.
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e90 77
probably occurred in Iceland (Kirkbride and Winkler, 2012).
Schaefer et al. (2009) dated three of the most prominent advances
in the past two millennia in New Zealand between CE 5 and CE 180
(Schaefer et al., 2009).
5.3.2. CE 200se300s
Advances during this period are recorded in many regions in
both hemispheres. In the Arctic, an extensive glacial expansion
throughout Baffin Island occurred between CE 250 and CE 400
Table 2.
Periods of glacier advances.
Region Centuries CE 1 2 3 4 5 6 7 8 9 10 11
Southern Alaska 200e320 550e720
Southern Alaska 200 400 600 800
Western Canada 400e600
Franklin Glacier, Mt. Waddington, BC 380e470
Baffin Island and the Queen Elizabeth Islands 250e400 400e500 600e900? 800e950
Greenland 300 490 660
SE Greenland ~650
Iceland >200 5th 600 9th 11th
Iceland
Spitsbergen 700e820
Spitsbergen 350 ± 200
Spitsbergen ~250
Scandinavia 200e300 600e700 800e1000
Southern Norway 100e400 200e300 600e700 900e1000
Novaya Zemlya
Franz Josef Land 427e766 11th
Northern Asia 200e600 700e1100 11th
Altay 400e600 700 early 11th
Alps 270 335 410 500s 835
Alps 312e337 485e606
Semi-arid Asia 400
Monsoon Asia 250e500 650e1050
Monsoon Asia 500 ± 200
SE Tibet 200e400 400e600 800e1150
Southern Himalaya 380e600 870e1100
Tibet
Qori Kalis outlet glacier, Quelccaya Ice Cap
Venezuelian Andes
Inner tropics 760e940
Outer tropics
Mt Kenia, Africa 400e750
Andes, Venezuela
Desert-central Andes, Pe~non, Azufre, Damas glaciers
Northern Patagonia, Tronador & Esperanza glaciers
Northern Patagonian Icefield
Southern Patagonian Icefield 330e770
Small glaciers, southern Patagonia 640e780
Tierra del Fuego 650e900
New Zealand 5e180 (3) 430 1005
New Zealand, Cameron Glacier
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9078
(Miller et al., 2012, 2013) and at CE 350 ± 200 in western Svalbard
(Reusche et al., 2014). In Alaska an advance occurred between CE
200 and CE 320 (R€othlisberger, 1986). Llewellyn Glacier in British
Columbia advanced into mature forest between CE 300 and CE 500
(Clague et al., 2010). Detailed records in the European Alps
(Nicolussi and Patzelt, 2001; Holzhauser et al., 2005; Le Roy et al.,
2015) prove first advances after the warm Roman period from the
late 3rd to early 5th century (e.g. a synchronous advance recorded
12 13 14 15 16 17 18 19 References
1180se1320s 1540se1710s 1600 1810se1880s Barclay et al., 2009a,b, this
paper
1180e1300 1600e1715 1820e1900 Wiles et al., 2008
1200e1500 1690e1730 1830e1850,
1860e1890
This paper
1150e1270 1340e1390 1380e1440 mid 19e
early 20th
Mood and Smith, 2015
1260e1280 1460 late19th Anderson et al., 2008;
Miller et al., 2012, 2013;
Margreth et al., 2014
1150e1660 1600 1750 1890e1900 Winsor et al., 2014, Kelly,
Lowell, 2009
1250e1300 1450e1630 1700e1930 Balascio et al., 2015
12th 13th 1400e1550 1700e1930 1720 1828e1930s This paper
1400e1550 1680e1890 Larsen et al., 2013
1160e1255 1350e1880 Guilizzoni et al., 2006
Reusche et al., 2014
~1725 ~1815 Røthe et al., 2015
1500e1800 Nesje, compilation from
lake sediments
1200e1300 1600e1800 1740e1750,
1780e1790
1860e1870 Matthews and Dresser,
2008, Nesje et al., 2008
1170e1400 1300 1400 19th Forman et al., 1999; Polyak
et al., 2004
>1200 >1400 mid 1600s late 19th Lubinski et al., 1999
>1737 19th This paper
12th cent. 1200e1260 1340e1380 1470e1480 1560 1640se1730s 19th (3) Nazarov et al., 2012
12th cent. late 13th 1350e1385 late 15th early 16th 1600, 1640, 1680 1720, 1775 1820 1855/60
1890
This paper
1120e1178 1248 1278/1296 >1352 Le Roy et al., 2015
1450 1640 19th This paper
1150e1400 1450e1850 R€othlisberger and Geyh,
1985
1300 ± 100 1600 ± 100 19th Murari et al., 2014
1400e1920 19th Yang et al., 2008
1400e1430 1550e1850 19th Yang et al., 2008
~15th >1662 1746e1785 early-19th,
late-19the20th
Br€auning, 2006
~1505 ~1635, ~1685 ~1705, ~1795 Stroup et al., 2014
1180e1350 1450e1590 1640e1730 1800e1820 Polissar et al., 2006
1720e1780 This paper
1480e1520 1610e1700 1690e1800 (~1730) This paper
~1350 1550 Karlen et al., 1999
1180e1350 1450e1590 1640e1730 1800e1820 Polissar et al., 2006
1460e1660 Espizua and Pitte 2009
>1230 1650s 1720e1790 1850s Villalba et al., 1990;
Masiokas et al., 2009a,
2010; Ruiz et al., 2012
1220e1350 1850e1890 Glasser et al., 2002;
Winchester et al., 2001;
Harrison et al., 2008;
Masiokas et al., 2009a
1130e1440 1450e1670 Mercer 1965;
R€othlisberger, 1986; Aniya,
2011; Masiokas et al.,
2009a; Strelin et al., 2014
1270e1400 1500e1650 1740e1800 Mercer 1965;
R€othlisberger, 1986;
Masiokas et al., 2009ab;
Strelin et al., 2014
1210e1440 1450e1680 Strelin et al., 2008;
Masiokas et al., 2009a
1155 1205 1435 1605 (2) 1735e1895 Schaefer et al., 2009,
recalculated from Putnam
et al., 2012
1350 1480 1770 1864, 1930 Putnam et al., 2012
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e90 79
between CE 312 and CE 337 at Mer de Glace and around CE 335 CE
at Gepatschferner). Two separate advances at ~CE 270 and ~CE 410
were identified at Great Aletsch Glacier. Some of these advances
may match a climatically effective eruption at CE 266 (Sigl et al.,
2015), although the dating uncertainties of moraines are generally
still too large for more definitive conclusions.
5.3.3. CE 400se700s
The advances around CE 600 are identified as a major regional
event in northwestern North America (Reyes et al., 2006). Advances
occurred at CE 550e720 in British Columbia and Alaska (Barclay
et al., 2009a,b; 2013; Zander et al., 2013). They are recorded in
southwest Greenland at CE 490 ± 110 (Winsor et al., 2014), in Iceland
with three advances between CE 450 and 550 (Kirkbride and
Dugmore, 2008), in Scandinavia in CE 600e700 (Matthews and
Dresser, 2008), in southeastern Tibetan Plateau in CE 400e600
(Yang et al., 2008), in New Zealand in ~CE 430 (Schaefer et al.,
2009), in the tropics in Africa at Mt. Kenya between ~ CE 400 and
750 (Karlen et al., 1999) and in South America at ~CE 450 ± 190
(Rodbell, 1992). In southern Patagonia and Tierra del Fuego advances
are recorded between the 4th and 9th centuries (Strelin
et al., 2008; Masiokas et al., 2009a). In some regions glaciers
were close to their past millennium maxima, e.g., Great Aletsch
Glacier and Mer de Glace in the Alps (Holzhauser et al., 2005; Le Roy
et al., 2015). Unfortunately the dating is still too poor to assess the
inter-regional synchronicity or distinguish individual stages.
Recently, Büntgen et al. (2016) identified a long-lasting cooling
between CE 536 and ~CE 660 in Eurasia after a set of large volcanic
eruptions in CE 536, 540 and 547 and named this period the “Late
Antique Little Ice Age”. Grand Solar Minimum (GSM), which
occurred at the end of the 7th century (Steinhilber et al., 2009), also
contributed to this cooling and probably to the glacier advances
around this time.
5.3.4. CE 800se900s
In the European Alps the advance around CE 835 was almost
equal to the past millennium maximum extent at some glaciers
(e.g. Suldenferner) (Nicolussi et al., 2006). The advance between CE
800 and 1150 was the largest in the past 2 ka in Gongga Shan (Yang
et al., 2008). An advance is recorded on Baffin Island between about
CE 850 and 950 (Miller et al., 2012) and in British Columbia between
CE 865 and 940 (Clague et al., 2010). Advances probably
occurred also in Spitsbergen at CE 700e820 (Guilizzoni et al., 2006)
and in Franz Josef Land at CE 720e880 (Lubinski et al., 1999). No
clear evidence of correspondence with volcanic eruptions is available
for this period. Volcanic events of moderate magnitude
occurred in CE 764 and CE 853 (Sigl et al., 2014).
5.3.5. CE 900e1000
Glaciers were retreating in the European Alps during the late 9th
to 11th centuries (Grove, 2004; Holzhauser et al., 2005; Simonneau
et al., 2014; Hafner, 2012), around CE 950 in southern coastal Alaska
(Wiles et al., 2014), and between CE 900 and 1100 in British
Columbia (Clague et al., 2010). Glacier advances in the 10th century
were rare, perhaps because this was a time when climate was
relatively unperturbed by anomalous solar or volcanic forcing
(Bradley et al., 2016). An advance around CE 1000 occurred in the
central Himalaya and, the southeastern Tibetan Plateau (Yang et al.,
2008), in Colombia (Jomelli et al., 2014), and New Zealand (Schaefer
et al., 2009). In western Canada, some glaciers advanced between
CE 1030 and 1210 (Clague et al., 2010), and in Baffin Bay an advance
is dated to between ~CE 1050 and 1220 (Young et al., 2015). One
strong high-latitude volcanic event is reported for CE 939 (Sigl
et al., 2015), but we do not find any clear evidence of potential
glacier advance around this time.
5.3.6. CE 1100s
Although the 12th century is considered the peak of the warm
period (MCA) in some places (Lamb, 1965), advances in the CE
1120se1180s are reported in many regions (Grove, 2004) (see
Table 2), e.g., western Canada (Luckman, 1986, 1996; Koch and
Clague, 2011), the European Alps (Le Roy et al., 2015), monsoon
Asia (R€othlisberger and Geyh, 1985), Greenland (Young et al., 2015),
Franz Josef Land (Lubinski et al., 1999), and New Zealand (Schaefer
et al., 2009). Two volcanic eruptions in CE 1108 and 1170 (Sigl et al.,
2015) may be responsible for these advances, but this is only
speculation based on coincidence of the ages.
5.3.7. CE 1200s
Miller et al. (2012) dated advances of glaciers on Baffin Island to
CE 1260e1280. Glaciers in eastern (Young et al., 2015) and southeastern
Greenland (Balascio et al., 2015) advanced around the same
time in CE 1250e1300. In the European Alps, Mer de Glace
advanced between ~CE 1248 and 1278e1296 (Le Roy et al., 2015),
and Gepatschferner advanced around CE 1286 (Nicolussi and
Patzelt, 2001). Nazarov et al. (2012) found evidence of trees killed
by advancing glaciers in the Altay in the CE 1200se1250s, and Soler
Glacier of the northern Patagonian Icefield advanced between CE
1222 and 1342 (Glasser et al., 2002). Advances in the 13th century
are also described in Alaska, western Canada, Scandinavia, Iceland,
and the Russian Arctic (regional descriptions in Section 4). The
advances might be related to the volcanic eruptions that occurred
in CE 1230, 1258, and 1286 (Sigl et al., 2015). The CE 1258 volcanic
event was the strongest of the past two millennia, and the period
between CE 1251 and 1310 is identified as one of the “volcanic-solar
downturns” (PAGES 2k Consortium, 2013) characterized by strong
cooling.
5.3.8. CE 1300s
A period of glacial activity, likely of global significance, experienced
its first culmination in the mid/late 14th century: e.g., Mer de
Glace in the Alps after CE 1352 (Le Roy et al., 2015); Franklin Glacier
in western Canada in CE 1340e1390 (Mood and Smith, 2015);
Spitsbergen after CE 1350 (Guilizzoni et al., 2006); southern Patagonia
after CE 1350 (Masiokas et al., 2009a); and New Zealand
around CE 1355 (Schaefer et al., 2009). Glaciers in the Peruvian
Andes also began to advance around the mid 14th century after a
long period of retreat (Stansell et al., 2013). The glacial activity
seems to have persisted at least until the end of this century. A
strong low-latitude eruption with bipolar climatic effects that
occurred in CE 1345 (Sigl et al., 2013) was possibly responsible for
the glacier advances recorded in the mid 14th century.
5.3.9. CE 1400se1500s
An advance around CE 1460s on Baffin Island (Miller et al., 2012)
is synchronous with the beginning of advances in southeast
Greenland (CE 1450e1630) (Balascio et al., 2015), in the Altay (CE
1440se1510s) (Nazarov et al., 2012), and in central Asia (CE
1450e1850) (R€othlisberger and Geyh, 1985). In the tropical Andes,
Stroup et al. (2014, 2015) report an advance of Qori Kalis in CE
1500 ± 20, and in New Zealand the advance around CE 1435 was
one of the most prominent (Schaefer et al., 2009; Putnam et al.,
2012). A strong eruption in the Pacific Ocean (Kuwae/Vanuatu)
that consisted probably of two events in CE 1452 and 1458 (Sigl
et al., 2013, 2015) and the Sp€orer solar minimum (~CE 1450)
roughly coincide with the beginning of above mentioned glacier
advances. Glacial advances in 16th century were generally less
common (see Table 2).
5.3.10. CE 1600s
Glaciers advanced between CE 1640 and ~1740 in the Altay
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9080
(Nazarov et al., 2012), in CE 1640se1670s in south-central Alaska
(Barclay et al., 2009a; Daigle and Kaufman, 2009), and around CE
1600 in Greenland (Briner et al., 2010, 2011). Two events are
recorded in New Zealand around CE 1605 (Schaefer et al., 2009). In
the southern Patagonian Andes, advances occurred between the
late 16th century and the mid 17th century (Glasser et al., 2004,
2011; Koch and Kilian, 2005; Aravena, 2007; Masiokas et al.,
2009a). An advance between CE 1640 and 1730 occurred in the
tropical Andes (Polissar et al., 2006), in CE 1638e1639 in northern
Patagonia (Frias Glacier) (Masiokas et al., 2009a), and in the middle
of the 17th century on Franz Josef Land (Lubinski et al., 1999). The
most-detailed and best-dated glacier advances are reported in the
Alps in the CE 1600s,1640s, and 1680s (Zumbühl et al.,1983; Grove,
2004; Nicolussi et al., 2006; Nussbaumer et al., 2007; Ivy-Ochs
et al., 2009; Le Roy et al., 2015). They closely correlate with eruptions
that occurred in CE 1601, 1641 and 1695 (Sigl et al., 2015).
Volcanic-solar downturns in CE 1581e1610 and 1641e1700 were
identified in the PAGES 2 ka temperature reconstruction (PAGES 2k
Consortium, 2013). In many regions the advances of the 17th century
were of a major magnitude (see details in Section 5.2).
5.3.11. CE 1700s
Two advances are recorded in the European Alps around CE
1720 and 1775 (Zumbühl et al., 1983; Nicolussi and Patzelt, 2001;
Nussbaumer et al., 2007), around CE 1725 in Spitsbergen (Røthe
et al., 2015), around ~CE 1720 in Iceland (Larsen et al., 2015),
shortly before 1737 in Kamchatka (Solomina et al., 1995), around
1735 and around 1770 in New Zealand (Schaefer et al., 2009;
Putnam et al., 2012), and around ~CE 1700, 1730 and 1780 in the
tropical Andes (Rabatel et al., 2005, 2008: Jomelli et al., 2014;
Stroup et al., 2014). Tree-ring-dated moraines in southern Patagonia
were deposited at ~CE 1740s (Masiokas et al., 2009a,b). Advances
between CE 1746 and 1785 occurred in the area of
monsoonal temperate glaciers in Asia (Br€auning, 2006). In southern
Norway glaciers advanced in CE 1740e1750 and CE 1780e1790
(Matthews and Dresser, 2008; Nesje, 2009). One group of these
advances (around CE 1770e1780) clusters around the CE 1783 Laky
eruption in Iceland; however, in the Alps, where the dating is more
precise the culmination of the late 18th century advance occurred
earlier than this eruption.
5.3.12. CE 1800s
Numerous glacier advances in the 19th century occurred and
were directly observed in many mountain regions. They are usually
recognized by the presence of well-preserved moraines of one of
the last millennium maxima and of subsequent readvances of
retreating glaciers. In many regions, such as the European Alps,
Scandinavia, Altay, and Patagonia, glacial records of the 19th century
are accurately dated, described and analyzed (e.g. Grove, 2004;
Wanner et al., 2000; Nicolussi and Patzelt, 2001; Holzhauser et al.,
2005; Luckman, 2000; Zumbühl et al., 2008; Masiokas et al., 2009a;
WGMS, 2008 etc.). Two periods of major advances in the European
Alps at around CE 1820 and 1855e1860 were followed by minor
readvances in the 1890s. In most regions the retreat from the 19th
century maximum began in the 1860s, although in some Arctic
regions the glaciers were still close to their maxima at the beginning
of 20th century. During the first three decades of the 19th
century e a period of several strong advances e a covariance between
a Grand Solar Minimum (Dalton) and several large tropical
volcanic eruptions is observed. The climatic effects of the Tambora
eruption in CE 1815 and another major low-latitude eruption in CE
1809 have been well studied (e.g. Harington, 1992; Grove, 2004).
5.3.13. 1900s CE to the present
In 20th and early 21st centuries the global trend of glacier
retreat with some episodes of regionally variable minor readvances
or still-stands is well documented, e.g., in the 1920s, 1970s and
1990s (Lemke et al., 2007; WGMS, 2008; Vaughan et al., 2013).
We summed the number of glacier advances for 50-year-long
intervals spanning the time period from CE 1 to CE 2000 for all
regions considered here and analyzed the agreement of the
resulting time series. If the time (the age) of advance covered more
than 50 years it was included in subsequent 50-year-long interval.
Although this approach has serious limitations it still provides a
general idea of the coherence of the time of advances across the
globe. The number of glacial advances for all regions is significantly
inter-correlated. This coherence is clear for both individual regions
(Fig. 5) and for cumulative curves of glacier advances for high, mid
and low latitudes (Fig. 6). The correlation is partly generated by the
shared millennial-scale trends towards both enhancing glacial activity
and increasing sample density (Table 2 SM). Removing the
long-term trend results in fewer significant correlations between
regions (Table 2 SM), though many regions still do demonstrate
century-scale co-variability. Based on this review we identified
potential synchronicity of glacier advances during the Common Era
that cluster in the CE 200se300s, 400se600s, 800s, around CE 1000
as well as around CE 1150s,1280s,1350s,1400s,1450s,1600s,1640s,
1680s, 1720s, 1770s, 1820s, 1850se1860s, 1890s, 1900se1920s,
1970se1980s and 1990s. Synchronicity of some glacier advances
over the past one to two millennia was advocated before by
R€othlisberger (1986), Grove (2004), Matthews (2007), Stroup et al.
(2014) and Solomina et al. (2015). However, the gaps inherent in
terrestrial glacial records, the scarcity of data for certain times,
together with the limited accuracy associated with many of the
proxy records and the complexity of glacier sensitivity to climate
parameters (Paterson, 1994; Oerlemans, 2001) all set limits on
identifying the global synchronicity of glacier variations.
5.4. Possible external forcings of glacier variations in the past 2 ka
Circulation dynamics and their related energy and mass fluxes
react in a complex manner to different forcing factors as well as to
internal (chaotic) variability and account for past global glacier
advances and retreats on different time scales. It is challenging to
translate forcings and internal (chaotic) variability into circulation
dynamics; precise regional or local temperature and precipitation
values with seasonal or higher resolution for the past two millennia
are also difficult to generate, thus we can only speculate about their
significance for past glacier advances and retreats. On the longer
(millennial) timescale changes in orbital forcing, mainly during the
summer season, played a major role, causing the latitudinal
displacement of the global circulation belts (Wanner et al., 2008).
On centennial to decadal timescales additional factors contributed
to glacier fluctuations: such as large tropical volcanic eruptions,
solar irradiance variability and greenhouse gas concentrations as
well as large-scale internal variability, mainly El Ni~no Southern
Oscillation (ENSO), Pacific Decadal Variability (PDV), North Atlantic
Oscillation (NAO) and Atlantic Multidecadal Oscillation (AMO)
(PAGES 2k, 2013). However, the fairly coherent response seen in the
overall record of glacier advances (Fig. 6) suggests that regional
circulation patterns were not the primary factors influencing
glacier dynamics.
Generally, the opposite orbital trends in the Northern and
Southern Hemispheres in the past two millennia would result in a
gradual increase of glacier sizes in the Northern Hemisphere and in
the decrease of those in the Southern Hemisphere. The orbital
signal is indeed well manifested in the Holocene glacier variations
(e.g. Solomina et al., 2015); however, in the past two millennia it is
masked by the higher-frequency perturbations such as solar
minima and sets of consequent volcanic eruptions reversing the
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e90 81
pattern of glacier variations expected from orbital forcing. This is
the case in some regions of the Southern Hemisphere where glacial
activity had increased by the second half of the second millennium
despite the increase of solar insolation in summer induced by the
orbital signal. Several hypotheses have been suggested to explain
this phenomenon (Broecker, 2000; Denton and Broecker, 2008;
Schaefer et al., 2009), but this problem is not yet completely
resolved. Several authors have drawn attention to the coincidence
of well-constrained glacial chronologies of past centuries with solar
activity (Luckman, 2000; Wiles et al., 2004, 2008; Holzhauser et al.,
2005; Hormes et al., 2006; Nussbaumer et al., 2011a,b; MacGregor
et al., 2011). Hormes et al. (2006) demonstrated the high correlation
of glacier length changes with total solar irradiance (up to
r2
¼ 0.79 with 9e14 year lags). However, neither the magnitude of
total solar irradiance (TSI) changes, nor the effects of such changes
on the climate system, are well understood (Lockwood, 2012) and
so a simple temporal correlation with available TSI proxies must be
treated with caution.
Porter (1981, 1986) suggested that volcanic eruptions are the
primary forcing of glacier fluctuations on a decadal timescale;
however, it was difficult to explain how the short-term coolings
induced by the volcanic eruptions can result in glacier advances
that last sometimes for decades. Recently, Sigl et al. (2015) showed
that the cooling may persist at least ten years after the largest
eruptions. Notably, some of these large eruptions occurred at the
same time as solar minima (“volcanic-solar downturns”; PAGES 2k,
2013). The combination of solar minima and volcanic forcing may
have been a decisive factor in driving widespread potentially
globally synchronous glacier advances of the past. Miller et al.
(2012) suggested an additional mechanism to maintain the shortlived
cooling periods that follow explosive eruptions: positive
feedbacks related to the expansion of Arctic sea ice. They noted that
abrupt summer cooling and expansion of Canadian Arctic ice caps
between CE 1275 and 1300, followed by even stronger cooling and
further advances at CE 1430e1455, coincided with major volcanic
eruptions, at times of decreased irradiance during Grand Solar
Minima (Steinhilber et al., 2009).
Unfortunately, as we showed in this overview, the precision of
the majority of the ages for glacier advances is still not high enough
to firmly connect the decadal-scale external forcing and glacier
advances, especially in the first millennium, though there is a
certain coincidence between glacier advances and large volcanic
eruptions that occurred in CE 1258, 1345, 1458, 1601, 1641, 1695,
1783 and 1815 (Sigl et al., 2015; see details in 5.3 and in Table 2).
Along with the global-scale forcing factors, modes of internal
climate system variability, such as the NAO (Nesje, 2009; Reichert
et al., 2001; Six et al., 2001; Imhof et al., 2012; Marzeion and
Nesje, 2012; Larsen et al., 2013; Young et al., 2015), ENSO (Koch
and Clague, 2011), AMO (Lüthi, 2014), and the IPO (Schaefer et al.,
2009) modulate regional climates, and may thereby affect glacier
variations in some areas. Such circulation modes may result in
asynchronous or sometimes antiphased cycles of glacier advances
and retreats (e.g. Scandinavia and the Alps in the early 18th century
(Dahl et al., 2003), Iceland and the Alps between CE 1550 and 1680
(Larsen et al., 2013). Koch and Clague (2011) reported that several
glaciers in western North America advanced up to their LIA limits in
Medieval time, possibly due to increased winter precipitation
forced by La Ni~na-like conditions.
Denton and Broecker (2008) suggested that the Atlantic Multidecadal
Oscillation (AMO) could have had thermal effects across
the Northern Hemisphere, and thus provide an explanation for the
synchronicity of glacier variations in the European Alps and in
western North America. Due to the bipolar seesaw effect, fluctuations
of glaciers in the Southern Hemisphere would be expected to
show the opposite behaviour at the decadal and multi-decadal
scale. However, as shown in Section 5.3 of this paper several advances
reported for the Southern Hemisphere in Patagonia and
New Zealand, e.g., ~CE 1150, 1350, 1430, 1600, 1735, 1770, are synchronous
with those of the Northern Hemisphere within the given
accuracy of the ages (see Table 2). Furthermore, Schaefer et al.
(2009) claim that their 10
Be ages on Holocene moraines support
neither synchronicity nor rhythmic asynchrony in glacier fluctuations
in the Northern and Southern Hemispheres. They suggest that
IPO may have played a crucial role in climate and glacier fluctuations
in New Zealand.
Anthropogenic warming is a global signal that is currently
resulting in retreat of glaciers in both hemispheres (Lemke et al.,
2007; Vaughan et al., 2013). We did not find evidence of any
other globally uniform period of glacier retreat in the past two
millennia except for the modern one, although in the 1ste2nd and
10th centuries, the number of advances seems to have been quite
low. However, our knowledge of the number and extent of glaciers
in retreat in the past is also limited, especially in the tropics and in
the Southern Hemisphere. In general the fast and global retreat of
mountain glaciers in the 20th to 21st centuries appears highly
unusual in the context of the two preceding millennia as well as for
the entire Holocene (Solomina et al., 2015) and is strongly influenced
by the rising temperature induced by anthropogenic forcing
(Oerlemans, 1994).
6. Conclusions
6.1. Synchroneity
1. Considering the global database reviewed in this paper, a
number of advances in the past two millennia cluster in CE
200se300s, 400se600s, 800s, CE 1000 as well as around CE
1150s, 1280s, 1350s, 1400s, 1450s, 1600s, 1640s, 1680s, 1720s,
1770s, 1820s, 1850se1860s, 1890s, 1900se1920s, 1970se1980s
and 1990s. However, in some regions the data are too scarce and
the precision of the dating is too low for a definite conclusion
about the global synchrony of glacier advances. Advances
occurred at one or more locations in every century of the first
millennium CE, but it is difficult to determine if these were of
more than regional significance. Very few cases of advancing
glaciers are reported for the 1st to 2nd and the 10th centuries
CE.
2. There is a coherency in the number of glacier advances in the
past two millennia in both hemispheres, including the glaciers
at low latitudes; the most extensive and geographically widespread
were in the 13th to 19th centuries and indicate that the
LIA was of global significance (Grove, 2001, 2004).
6.2. Magnitude of changes
1. The full range of glacier variations in the past two millennia is
unknown. During the maximum advances DELA ranged from 10
to 20 m in arid central Asia, between 100 and 150 m in the mid
latitudes and up to 500 m in the tropics. During the 1st and 2nd
centuries, ELAs were generally higher than during the 21st
century by 20e25 m in Scandinavia, up to 100 m in Spitsbergen,
and at least 200 m in the eastern Canadian Arctic.
2. The number of recorded glacial advances in the 1st millennium
CE is much smaller than in the 2nd millennium, probably not
only due to less numerous advances, but more importantly
because they were of generally smaller magnitude, hence the
evidence was subsequently overridden. Some advances (e.g. in
6th to 7th centuries in Arctic Canada, west Greenland, Svalbard,
O.N. Solomina et al. / Quaternary Science Reviews 149 (2016) 61e9082
the Alps) were of similar magnitude to those of the “classical”
LIA and, thus required similarly strong external forcing.
6.3. External forcings
1. In most cases considered here summer temperature rather than
winter precipitation was the primary driver of decadal to century
scale glacier variations over the past two millennia. In some
areas, such as the maritime regions in the northern North
Atlantic, western Canada, Patagonia, and New Zealand, glacier
advances during the past two millennia were also dependent on
the winter precipitation regime, driven by prevailing circulation
modes.
2. Higher levels of explosive volcanic activity, perhaps aided by low
solar irradiance forcing and amplified by internal climate system
feedbacks, e.g., sea-ice expansion, led to the expansion of glaciers
between the beginning of the 13th to the early 19th centuries.
Southern Hemisphere glacier advances occurred in the
same period despite summer insolation increases, and neither
sea ice nor AMOC provides a major direct trigger for glacier
growth in the Southern Hemisphere in the past two millennia.
Thus the linkages across the hemispheres and identifications of
forcings remain largely unresolved.
3. Multiple forcings are responsible for glacier advances. Several
advances, such as those occurring in the 6th to 7th centuries CE,
in the late 13th, in the 14th, in the early and mid 15th, as well as
in the 17th to early 19th centuries generally coincide with the
timing of large explosive low-latitude volcanic eruptions and
minima of solar activity.
6.4. 20th century
Glaciers were as small as today (late 20th e early 21st centuries)
in some regions and likely even smaller several times in the past
two millennia. However, the current globally widespread glacier
retreat is unusual in the context of the past two millennia and,
indeed, for the whole Holocene. Contemporary glacier retreat
breaks a long-term trend of increased glacier activity that
dominated the past several millennia. The trend of glacier
retreat is global, and the rate of this retreat has increased in the
past few decades. The observed widespread glacial retreat in the
past 100e150 years requires additional forcing outside the
realm of natural changes for their explanation. With a high
probability anthropogenic greenhouse gas emissions are the
most likely reason for the prevalent, ongoing global ice loss.
Acknowledgments
We are grateful to our collegues Zhu Haifeng, Philipp
Hochreuther, David Loibl, Andrey Nazarov, Vladimir Myglan, Irina
Bushueva, Liudmila Lazukova, Vladimir Mikhalenko, Mikhail Alexandrin,
Elya Zozovskaya, Leo Martin for their valuable help and
additional material that we requested for this paper. M. Masiokas
acknowledges support from Consejo Nacional de Investigaciones
Científicas y Tecnicas (CONICET) and Agencia Nacional de
Promocion Científica y Tecnica (grants PICT 2007-0379 and PICT
2010-1438. Financial support for V.Jomelli was provided by the LMI
GREAT ICE, the ANR El Paso programme no.10-blan-68-01 and CEPS
Green Greenland project (grant ANR-10-CEPL-0008). LDEO contribution
no. 8033 G. Wiles was supported by the U.S. National Science
Foundation under grants ATM-0902799, 1202218, and AGS-
1502186. O. Solomina was supported by the scientific programms N
4 and N 15 of Russian Academy of Sciences Kurt Nicolussi has been
supported by the Austrian Science Fund FWF (project I 1183-N19).
We very much appreciated the comments of Stepan Harrison
and anonymous reviewer of the first versions of this paper and the
understanding of the QSR team that was very supportive.
Appendix A. Supplementary data
Supplementary data related to this article can be found at http://
dx.doi.org/10.1016/j.quascirev.2016.04.008.
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