The Cryosphere, 10, 3091–3105, 2016
www.the-cryosphere.net/10/3091/2016/
doi:10.5194/tc-10-3091-2016
© Author(s) 2016. CC Attribution 3.0 License.
Radiocarbon dating of glacier ice: overview,
optimisation, validation and potential
Chiara Uglietti1,2,3, Alexander Zapf1,2,3,†, Theo Manuel Jenk1,3, Michael Sigl1,3, Sönke Szidat2,3, Gary Salazar2,3, and
Margit Schwikowski1,2,3
1Laboratory of Environmental Chemistry, Paul Scherrer Institute, 5232 Villigen PSI, Switzerland
2Department of Chemistry and Biochemistry, University of Bern, 3012 Bern, Switzerland
3Oeschger Centre for Climate Change Research, University of Bern, 3012 Bern, Switzerland
†deceased
Correspondence to: Theo Manuel Jenk (theo.jenk@psi.ch)
Received: 8 July 2016 – Published in The Cryosphere Discuss.: 22 July 2016
Revised: 26 October 2016 – Accepted: 14 November 2016 – Published: 21 December 2016
Abstract. High-altitude glaciers and ice caps from midlatitudes
and tropical regions contain valuable signals of past
climatic and environmental conditions as well as human activities,
but for a meaningful interpretation this information
needs to be placed in a precise chronological context. For
dating the upper part of ice cores from such sites, several relatively
precise methods exist, but they fail in the older and
deeper parts, where plastic deformation of the ice results in
strong annual layer thinning and a non-linear age–depth relationship.
If sufficient organic matter such as plant, wood
or insect fragments were found, radiocarbon (14C) analysis
would have thus been the only option for a direct and absolute
dating of deeper ice core sections. However such fragments
are rarely found and, even then, they would not be very
likely to occur at the desired depth and resolution. About
10 years ago, a new, complementary dating tool was therefore
introduced by our group. It is based on extracting the
µg-amounts of the water-insoluble organic carbon (WIOC)
fraction of carbonaceous aerosols embedded in the ice matrix
for subsequent 14C dating. Since then this new approach
has been improved considerably by reducing the measurement
time and improving the overall precision. Samples with
∼ 10 µg WIOC mass can now be dated with reasonable uncertainty
of around 10–20 % (variable depending on sample
age). This requires about 300 to 800 g of ice for WIOC concentrations
typically found in midlatitude and low-latitude
glacier ice. Dating polar ice with satisfactory age precision
is still not possible since WIOC concentrations are around
1 order of magnitude lower. The accuracy of the WIOC 14C
method was validated by applying it to independently dated
ice. With this method, the deepest parts of the ice cores from
Colle Gnifetti and the Mt Ortles glacier in the European Alps,
Illimani glacier in the Bolivian Andes, Tsambagarav ice cap
in the Mongolian Altai, and Belukha glacier in the Siberian
Altai have been dated. In all cases a strong annual layer thinning
towards the bedrock was observed and the oldest ages
obtained were in the range of 10 000 years. WIOC 14C dating
was not only crucial for interpretation of the embedded
environmental and climatic histories, but additionally gave a
better insight into glacier flow dynamics close to the bedrock
and past glacier coverage. For this the availability of multiple
dating points in the deepest parts was essential, which is the
strength of the presented WIOC 14C dating method, allowing
determination of absolute ages from principally every piece
of ice.
1 Introduction
High-altitude glaciers and ice caps from midlatitudes and
tropical regions contain valuable signals of past climate and
atmospheric variability at regional and local scales and are
located in areas with large biological diversity and inhabited
by the majority of the world’s population. Midlatitude
glaciers, for instance in the European Alps or in the Himalayas,
are influenced, in particular, by the nearby anthroPublished
by Copernicus Publications on behalf of the European Geosciences Union.
3092 C. Uglietti et al.: Radiocarbon dating of glacier ice: overview, optimisation, validation and potential
pogenic pollution sources, thereby additionally preserving
the signature of human activities. This information can generally
be retrieved from glacier ice cores, but needs to be
placed in a precise chronological context to allow meaningful
interpretation with respect to environmental and climatic
changes.
Ice core dating is a sophisticated task and the most common
approach is annual layer counting, which relies on seasonally
fluctuating signals. A number of ice core parameters
such as the stable isotope ratio of hydrogen or oxygen
in the water (δ2H, δ18O), the concentration of trace components
(e.g. ammonium, mineral-dust-related trace elements,
black carbon) and the presence of melt layers may vary with
the seasons. To reduce uncertainty in layer counting, the
timescale is additionally anchored with reference horizons
like the radioactivity peak resulting from nuclear weapon
tests in the 1960s or tephra and aerosol layers caused by volcanic
eruptions (Thompson et al., 1998, 2013; Preunkert et
al., 2000; Schwikowski, 2004; Eichler et al., 2009; Moore et
al., 2012). An independent method is nuclear dating with the
naturally occurring radioisotope 210Pb. Determined by the
210Pb half-life of 22.3 years and its atmospheric concentration,
the time period accessible for dating is in the order of
a century (Gäggeler et al., 1983; Eichler et al., 2000; Herren
et al., 2013). All these dating techniques fail in the older
and deeper parts of glaciers, where plastic deformation of
the ice under the weight of the overlying mass results in horizontal
ice flow, stretching annual layers continuously with
increasing depth. Correspondingly, the age–depth relationship
of high-alpine glaciers is strongly non-linear (Jenk et al.,
2009) and annual layers and also volcanic signals become undetectable
below a certain depth with the current spatial resolution
of most analytical methods. Glacier flow modelling
can only give rough age estimates with large uncertainties
close to the bedrock of high-alpine glaciers (Lüthi and Funk,
2001). Radiocarbon (14C) analysis has been the only option
allowing direct and absolute dating of these deeper ice core
sections in the rare cases when sufficient organic matter such
as plant, wood or insect fragments were found (Thompson
et al., 1998, 2002). However, in glacier ice such findings are
not only very seldom, but even when they occur, they do not
allow for continuous or at least regular dating, which per se
limits the application of the 14C technique and its use for
deriving a complete chronology based on absolutely dated
layers. In the following we refer to the dating of ice with
macrofossils as conventional 14C dating.
A new, complementary dating tool was therefore introduced
by our group about 10 years ago, based on extracting
the µg-amounts of the water-insoluble organic carbon fraction
of carbonaceous aerosols embedded in the ice matrix
for 14C dating (Jenk et al., 2006, 2007). Carbonaceous compounds
represent a large, but highly variable fraction of the
atmospheric aerosol mass (Gelencsér, 2004; Hallquist et al.,
2009). Total organic carbon (TOC, also referred to as total
carbon, TC) is instrumentally divided into two subfractions
according to their refractory and optical properties. Elemental
carbon (EC) consists of highly polymerised substances
which are extremely refractory and light absorbent; therefore
this fraction is also called black carbon (BC) or soot (Gelencsér,
2004; Hallquist et al., 2009). EC is merely derived
from the incomplete combustion of fossil fuels and biomass.
Organic carbon (OC) is formed by weakly refractory hydrocarbons
of low to medium molecular weight. Whereas
EC is generally insoluble in water, OC is further subdivided
into water-soluble organic carbon (WSOC) and waterinsoluble
organic carbon (WIOC) (Szidat et al., 2004a). In
water samples the former is also known as dissolved organic
carbon (DOC) (Legrand et al., 2013; May et al., 2013).
OC is emitted directly as primary aerosol from a vast variety
of sources and emission processes, including mobilisation
of plant debris, pollen, vegetation waxes, microorganisms,
spores, the organic fraction of soil as well as emissions
from biomass burning (e.g. forest fires) and anthropogenic
processes (biomass burning and fossil fuel combustion), but
it is also formed in the atmosphere by oxidation of gaseous
precursors, with the product referred to as secondary organic
aerosol (Gelencsér, 2004; Gelencsér et al., 2007; Hallquist et
al., 2009).
Carbonaceous aerosols are transported in the atmosphere
to high-alpine glaciers, where they may be deposited by both
wet and dry deposition processes and finally embedded in
glacier ice (Lavanchy et al., 1999; Jenk et al., 2006; Legrand
and Puxbaum, 2007; McConnell et al., 2007; Kaspari et
al., 2011). Consequently, using carbonaceous aerosols allows
any piece of ice to be dated, given that it contains sufficient
carbon mass. The WSOC fraction (i.e. DOC) would be ideal
for dating, since it has the highest concentrations in ice. However,
its extraction is complicated. It involves the outgassing
of aqueous atmospheric CO2, removal of dissolved carbonates,
wet oxidation of the organic compounds to CO2 under
inert gas and, finally, quantitative trapping of the evolved
CO2 (May et al., 2013). Since major contributors of DOC,
like light carboxylic acids, are ubiquitous in the air, all these
steps are prone to contamination. Therefore, and because of
the reasons summarised next, WIOC was selected from the
different carbonaceous particle fractions as the most promising
target for 14C dating. First, WIOC is mainly of biogenic
origin in pre-industrial times (Jenk et al., 2006) and, therefore,
supposed to contain a contemporary 14C signal representative
of the age of the ice (Jenk et al., 2006; Steier et
al., 2006). Second, the average WIOC concentration in ice
is higher than the respective EC concentration, allowing for
smaller ice samples and a potentially higher time resolution,
which consequently provides a better signal-to-noise ratio
(mainly determined by the overall blank) and smaller uncertainty
of the dating results. Third, OC has a lower probability
compared to EC for in-built reservoir ages, e.g. from the
burning of old trees or old organic matter (Gavin, 2001; Sigl
et al., 2009). Moreover OC is insensitive to potentially insufficiently
removed carbonates in mineral-dust-rich layers
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C. Uglietti et al.: Radiocarbon dating of glacier ice: overview, optimisation, validation and potential 3093
Table 1. Characteristics of the sites discussed and the respective dating approach. ALC stands for annual layer counting, RH for reference
horizons and 210Pb, 3H, and 14C for nuclear dating. 2p model (two-parameter model), MC (Monte Carlo simulation) and EF (exponential
fit) denote the applied approach to finally derive a continuous age–depth relationship (see Sect. 6 for details).
Site Coordinates Location Dating approach Time References
elevation span
(years)
Belukha 49.80◦ N, 86.55◦ E Altai Mountains, ALC, RH, 3H, 14C, ∼ 9100 Aizen et al. (2016)
4115 m a.s.l. Russia 2p model
Colle Gnifetti 45.93◦ N, 7.88◦ E Western Alps, ALC, RH, 3H, 210Pb, > 15 200 Jenk et al. (2009)
4450 m a.s.l. Swiss–Italian border 14C, 2p model
Juvfonne 61.68◦ N, 8.35◦ E Jotunheimen 14C of organic-rich layers ∼ 7600 Zapf et al. (2013),
1916 m a.s.l. Mountains, Norway and WIOC Ødegård et al. (2016)
Illimani 17.03◦ S, 68.28◦ W Andes, Bolivia ALC, RH, 3H, 210Pb, ∼ 12 700 Sigl et al. (2009),
6300 m a.s.l. 14C, 2p model Kellerhals et al. (2010)
Mt Ortles 46.51◦ N, 10.54◦ E Eastern Alps, ALC, RH, 3H, 210Pb, ∼ 6900 Gabrielli et al. (2016)
3905 m a.s.l. Italy 14C, MC
Quelccaya 13.93◦ S, 70.83◦ W Andes, Peru ALC, 14C ∼ 1800 Thompson et al. (2013)
5670 m a.s.l.
Tsambagarav 48.66◦ N, 90.86◦ E Altai Mountains, ALC, RH, 3H, 210Pb, ∼ 6100 Herren et al. (2013)
4130 m a.s.l. Mongolia 14C, EF
(e.g. Saharan dust), which may contribute to the EC fraction
because of the higher combustion temperature applied
to EC (Jenk et al., 2006). The extraction of WIOC from the
ice is straightforward as it can be collected by filtration of
the melted ice. Note that in previous publications (Sigl et al.,
2009; Zapf et al., 2013) the term POC was used for particulate
organic carbon (Drosg et al., 2007). Since POC can
be mistaken for primary organic carbon (Gelencsér, 2004;
Zhang et al., 2012), we adopted the term water-insoluble organic
carbon (WIOC) instead in this overview.
Our research group has a long history in 14C dating of
ice core using the aforementioned WIOC fraction of carbonaceous
particles. Lavanchy et al. (1999) introduced the
initial methods to determine the concentrations of carbonaceous
particles in ice from a European high-alpine glacier.
Next, the methodology was developed for source apportionment
of aerosols by 14C measurements in different carbonaceous
particle fractions (Szidat et al., 2004b). This was conducted
in close collaboration with the Laboratory of Ion
Beam Physics of the ETH Zurich, a well-established 14C dating
facility and a world-leading group in accelerator mass
spectrometry (AMS) technology, where the analytical aspect
of instrumentation was simultaneously and continuously improved
(Synal et al., 2000, 2007; Ruff et al., 2007, 2010).
The methodology of 14C analysis of the different carbonaceous
particle fractions was adopted to study the suitability
of WIOC for 14C dating of old ice, finding that it is of purely
biogenic origin prior to industrialisation (Jenk et al., 2006,
2007). Since then this novel 14C approach has been applied
for the dating of a number of ice cores from different highaltitude
mountain glaciers (Table 1) (Jenk et al., 2009; Sigl
et al., 2009; Kellerhals et al., 2010; Herren et al., 2013; Zapf
et al., 2013; Aizen et al., 2016). Meanwhile the method has
been further optimised and was additionally validated by determining
the age of independently dated ice. Here we give
an overview of the current status of the now routinely applied
WIOC 14C dating method for glacier ice, including an update
on recent optimisation and method validation. Uncertainties
and the potential of this novel approach are discussed and its
successful application to a number of ice cores is presented.
2 Sample preparation, OC/EC separation and 14C
analysis
The preparation of ice samples follows the procedure according
to Jenk et al. (2007). First, samples are decontaminated
in a cold room (−20 ◦C) by removing the outer layer (3 mm)
with a pre-cleaned stainless steel band saw (wiped 3 times
with acetone, followed by multiple cuts through a frozen
block of ultra-pure water, 18 M cm quality), followed by
rinsing the samples with ultra-pure water in a class 100 clean
bench. Around 20–30 % of the ice samples’ mass is lost during
these first steps, resulting in a final mass of about 200 to
500 g (initial mass of around 300–800 g of ice). The samples
are then transferred and stored in a freezer at −20 ◦C
in pre-cleaned (soaked and rinsed for 3 days with daily exchanged
ultra-pure water) 1 L-containers (Semadeni, PETG)
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until they are melted at room temperature directly before filtration.
To ensure that carbonates potentially present in the
ice are completely dissolved, ∼ 20 mL of 1 M HCl (30 %
Suprapure, Merck) are added to the melted samples (Cao et
al., 2013), resulting in a pH of < 2, before being sonicated
for 5 min. Subsequently, the insoluble carbonaceous particles
are filtered onto preheated (5 h at 800 ◦C) quartz-fibre filters
(Pallflex Tissuquartz, 2500QAO-UP) using a dedicated glass
filtration unit, also carefully pre-cleaned by being rinsed with
ultra-pure water and baking the glass at 450 ◦C for 3 h. As
a second carbonate removal step, the filters are acidified 3
times with a total amount of 50 µL 0.2 M HCl (Jenk et al.,
2007). Afterwards the filters are left in a class 100 clean
bench for 1 h to allow potentially present carbonates to be
transformed into CO2 through reaction with HCl, then they
are rinsed with 5 mL ultra-pure water to entirely remove remaining
HCl. The filters are left again for 1 h to reach complete
dryness, packed in aluminium foil and kept frozen until
analysis, when the filters are taken out of the freezer to
let them reach ambient temperature (at least half an hour).
Details regarding OC and EC separation, AMS 14C analysis
and improvements achieved since the first applications will
be discussed in Sects. 3 and 4.
3 Recent optimisation in OC/EC separation and AMS
analysis
In previous ice core dating applications using 14C of WIOC
(Jenk et al., 2009; Sigl et al., 2009; Kellerhals et al., 2010;
Herren et al., 2013; Zapf et al., 2013), the OC and EC combustion
was performed with the two-step heating system for
the EC/OC determination of radiocarbon in the environment
(THEODORE), developed for aerosol applications (Szidat et
al., 2004b). The combustion was conducted in a stream of
oxygen for the controlled separation of OC and EC fractions.
The temperature for OC separation was set at 340 ◦C, while
for the recovery of EC the temperature was then increased
to 650 ◦C. The CO2 produced by oxidation during the combustion
was cryogenically trapped, manometrically quantified
and sealed in glass ampoules (Szidat et al., 2004b). In the
earliest application described by Jenk et al. (2006) the CO2
subsequently had to be transformed to filamentous carbon
(graphitisation) using manganese granules and cobalt powder
prior to the final AMS 14C analysis. This was initially performed
at the ETH AMS facility (TANDY, 500 kV pelletron
compact AMS system) (Synal et al., 2000). Since 2006, the
200 kV compact AMS (mini radiocarbon dating system, MICADAS)
has been in operation at the ETH (Synal et al.,
2007). The MICADAS is equipped with a gas ion source and
a Gas Interface System (GIS) (Ruff et al., 2007; Synal et al.,
2007), allowing measurements of 14C directly in CO2 with
an uncertainty level as low as 1 % (Ruff et al., 2010). The
GIS includes a gas-tight syringe for injecting the CO2 into
the ion source of the AMS (Ruff et al., 2010). The syringe
has a maximum capacity of 1.3 mL of CO2 (equivalent to
100 µg of carbon). The position of the syringe plunger is automatically
adjusted according to the sample size as well as
the helium flow carrying the sample to the ion source. With
this, the transformation of gaseous CO2 to solid graphite targets
became needless (Sigl et al., 2009). Instead, the glass
ampoules sealed after the combustion of the filters with the
THEODORE system were opened in a designated cracker, an
integral part of the GIS (Ruff et al., 2007), and the resulting
CO2-He mixture could directly be fed into the MICADAS
ion source.
The main advantages of switching from solid to gaseous
targets were (1) a decrease in the number of necessary preparation
steps and the associated risk of lost samples from
incomplete graphitisation, (2) a higher sample throughput,
(3) a reduction in the variability and overall blank contribution
as well as (4) the elimination of the correction applied
to account for fractionation during the graphitisation
step, which contributed around 10 % to the overall uncertainty
(Jenk et al., 2007). As will be discussed in Sect. 4, a
precision increase is one of the main challenges for improving
the method.
Since spring 2013, the 14C analysis has been performed
with a MICADAS installed at the Laboratory for the Analysis
of Radiocarbon with AMS (LARA laboratory) of the University
of Bern, also equipped with a GIS interface (Szidat et
al., 2014). There, an improvement was recently achieved by
replacing THEODORE with a commercial combustion system,
which is a thermo-optical OC/EC analyser (Model4L,
Sunset Laboratory Inc., USA), normally used for aerosol
OC/EC separation and source apportionment studies (Zhang
et al., 2012, 2013, 2014; Zotter et al., 2014). Similarly to the
THEODORE system, the carbonaceous particles are combusted
in a stream of pure oxygen. The Sunset instrument
is specially equipped with a non-dispersive infrared (NDIR)
cell to quantify the CO2 produced during the combustion.
The combustion process in the Sunset system follows a wellestablished
protocol (Swiss 4S) for the thermal separation
of OC and EC fractions under controlled conditions (Zhang
et al., 2012). To avoid potential damage to the infrared cell
detector by residual HCl, the final rinsing of the filters after
adding HCl for carbonates removal was introduced (see
Sect. 2). Recently the Sunset instrument was directly coupled
to the zeolite trap of the GIS (Ruff et al., 2010), which allows
online 14C measurements of the carbonaceous fractions
separated in the Sunset system (Agrios et al., 2015). When
combusted, the gaseous carbonaceous species pass through
a MnO2 bed heated to 850 ◦C for completing the oxidation
to CO2, which is further transported by helium to the zeolite
trap. This trap is then heated up to 500 ◦C to release the CO2
to the gas-tight syringe for the final injection into the AMS
ion source (Ruff et al., 2007; Synal et al., 2007).
The newly coupled Sunset-GIS-AMS system has major
advantages compared to the old set-up. The OC/EC separation
in THEODORE was relatively time consuming and only
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C. Uglietti et al.: Radiocarbon dating of glacier ice: overview, optimisation, validation and potential 3095
Table 2. Samples analysed for the comparability test for OC/EC separation using the THEODORE apparatus and the Sunset OC/EC analyser
directly coupled to the AMS, with WIOC masses and concentrations. Calibrated ages (cal BP) denote the 1σ range.
Sample ID AMS Lab. no. WIOC WIOC F14C 14C age (BP) Cal age
mass concentration (cal BP)
(µg) µg kg−1 ice
1_THEODORE (JUV 3) ETH 42845.1.1 44 176 1.134 ± 0.017 −1010 ± 120 −46 – −7
ETH 42847.1.1
ETH 42849.1.1
ETH 43446.1.1
1_Sunset (JUV 3) BE 3683.1.1 46 119 1.157 ± 0.014 −1171 ± 97 −42 – −8
BE 3701.1.1
2_THEODORE (JUV 1) ETH 43555.1.1 18 60 0.743 ± 0.029 2386 ± 314 2011–2783
ETH 43557.1.1
2_Sunset (JUV 1) BE 3679.1.1 9 33 0.744 ± 0.021 2376 ± 225 2158–2737
3_THEODORE (BEL 1) ETH 42841.1.1 18 63 0.771 ± 0.017 2089 ± 177 1886–2310
3_Sunset (BEL 1) BE 4282.1.1 15 61 0.725 ± 0.022 2587 ± 233 2353–2924
4_THEODORE (BEL 2) ETH 43448.1.1 15 47 0.402 ± 0.022 7320 ± 440 7686–8588
4_Sunset (BEL 2) BE 4175.1.1 18 48 0.387 ± 0.022 7626 ± 457 7999–9011
four ice samples could be processed per day. Two more days
were needed to produce all the standards and blanks required
for AMS calibration and for quality control and graphitisation
(Jenk et al., 2007). Besides the disadvantages of solid
graphite targets described before, there is also a risk of losing
samples during the delicate phase of flame-sealing the
ampoules and later on when scratching them to allow a clean
break in the automated GIS cracker. With the online coupling
of the Sunset, this risk is completely removed. Further the
preparation and measurement time is significantly reduced
because there is no need for offline combustion, resulting in
a total measurement time of approximately 35 min per sample
only. In addition, it not only allows for an automated
protocol of standard injection for AMS calibration, but also
offers the possibility for easy and regular (daily) survey of
the 14C background in the entire process line (Sunset-GISAMS)
by analysis of variably sized standards and blanks if
required (Agrios et al., 2015). Finally, the Sunset system enables
continuous monitoring of the combustion process, reducing
a potential bias due to charring, and the standardised
and automated combustion protocol (Swiss 4S) ensures high
reproducibility increasing the overall precision.
With the current set-up, the 14C / 12C ratio of the samples
is background subtracted, normalised and corrected for
mass fractionation by using fossil sodium acetate (14C free,
NaOAc, p.a., Merck, Germany), the reference material NIST
standard oxalic acid II (modern, SRM 4990C) and the δ13C
simultaneously measured in the AMS, respectively (Wacker
et al., 2010). All results are expressed as fraction modern
(F14C), which is the 14C / 12C ratio of the sample divided
by the same ratio of the modern standard. Further corrections
are subsequently applied to the F14C values considering isotopic
mass balance (e.g. Jenk et al., 2007) to account for constant
contamination, cross contamination and for the procedural
blank contribution introduced by the preparation of ice
samples (for details see Sect. 4). 14C ages (before present
– BP, i.e. before 1950) are calibrated using OxCal v4.2.4
(Bronk Ramsey and Lee, 2013) with the Northern (IntCal13)
or Southern Hemisphere (ShCal13) calibration curves (Hogg
et al., 2013; Reimer et al., 2013), depending on the sample
site location. Calibrated dates are given in years before
present (cal BP) with 1σ uncertainty range (Stuiver and Polach,
1977; Mook and van der Plicht, 1999). For simplicity
the ages discussed in the text are given as the mean of this
range ±1σ. See Sect. 4 for further details regarding the applied
corrections, 14C calibration and discussion of uncer-
tainties.
To ensure comparability between previous data and the
newly derived results using the above-described improved
set-up configuration, 14C analysis was conducted on remaining
pieces of samples, which were previously processed with
the THEODORE set-up. Two samples (JUV 1 and JUV 3)
from the Juvfonne ice patch in Norway (Zapf et al., 2013)
and two samples (BEL 1 and BEL 2) from an ice core drilled
at Belukha glacier in the Siberian Altai (Aizen et al., 2016)
were used, covering an age range from modern to more than
8000 cal BP. The OC masses were above 10 µg carbon, except
for sample Juv2_Sunset with a carbon mass of 9 µg (Table
2), still resulting in more than 4500 14C counts with a
corresponding uncertainty of the F14C of 2 %, which we
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consider sufficiently low for this comparison. At first, the
obtained WIOC concentrations are discussed, which are assumed
to agree as indicated by a carbon quantification test
carried out on homogeneous aerosol filters using both combustion
instruments (Zotter et al., 2014). As expected, a good
consistency was found for the WIOC concentrations in the
Belukha ice core (Table 2), whereas a discrepancy was observed
for the Juvfonne samples, probably related to the natural
inhomogeneity of particles in this small-scale ice patch
with a distinct ice accumulation behaviour (see below). Concerning
the 14C ages, a very good agreement is shown between
all parallel samples (Fig. 1). This is also true for the
procedural blanks, both in terms of carbon amount and F14C.
The WIOC procedural blank measured and used for correction
in this comparison experiment was 1.41 ± 0.69 µg of carbon
with an F14C of 0.64 ± 0.12, and 1.21 ± 0.51 µg of carbon
with an F14C of 0.73 ± 0.13 for THEODORE and the
coupled Sunset set-up, respectively (additional details can be
found in Sect. 4). In summary, we conclude that dating results
obtained with the previously used THEODORE combustion
set-up (Jenk et al., 2009; Sigl et al., 2009; Herren
et al., 2013; Zapf et al., 2013) and the improved coupled
Sunset-GIS-AMS system are in good agreement.
4 Radiocarbon dating uncertainties
First of all, the signal-to-noise ratio of the AMS measurement
is defined by counting statistics. Generally, the smaller
the sample, the shorter the measurement time and the higher
the uncertainty. For defining the contamination contribution
of the overall instrument set-up (constant contamination) and
the memory effect between subsequent samples of very different
14C content and carbon mass (cross contamination), a
series with varying amounts of solid grains of fossil NaOAc
and the modern reference material oxalic acid II was combusted
with the Sunset and measured for its 14C content. The
constant contaminant mass was estimated as 0.4 ± 0.2 µg carbon
with a F14C of 0.8 ± 0.4 and for the cross contamination
0.5 ± 0.4 % of the carbon of the previous sample was found
to mix with the next injection (Agrios et al., 2015).
The total carbon amounts in ice cores are rather low, in the
µg kg−1-range. Because of that, each step of sample preparation
implies a potential risk of contamination with either
modern or fossil carbon. Thus a large contribution to the final
overall uncertainty on the age is induced by the procedural
blank correction, especially for small size samples. It
is therefore crucial that cutting, melting and filtrating the ice
results in the lowest possible procedural blank with a stable
F14C value to ensure a high and stable signal-to-blank
ratio for obtaining reliable results with the smallest possible
uncertainties. Procedural blanks were estimated using artificial
ice blocks of frozen ultra-pure water, treated in the
same way as real ice samples (Jenk et al., 2007). Blanks
were usually prepared together with samples and their anal-
1_THEODORE
1_Sunset
2_THEODORE
2_Sunset
3_THEODORE
3_Sunset
4_THEODORE
4_Sunset
02000400060008000100001200014000
Calibrated date (calBP)
Figure 1. OxCal output for the comparability test for OC/EC separation
using the THEODORE apparatus and the Sunset OC/EC
analyser directly coupled to the AMS. Bars below the age distributions
indicate the 1σ range. See Table 2 for details of the samples.
ysis was performed during every AMS measurement session
(Sunset combustion and AMS analysis). The mean of the
overall procedural blank (WIOC) used to correct all samples
is 1.34 ± 0.62 µg of carbon with a F14C of 0.69 ± 0.13
(100 and 54 measurements, respectively, performed over a
10-year period). This includes all values obtained with both
the THEODORE and Sunset systems. We decided to use this
combined value, since the ice sample preparation step is the
by far the largest contribution to the blank and is system independent.
This mean value is consistent with previously reported
results (Jenk et al., 2007; Sigl et al., 2009), indicating
the long-term stability of the procedural blanks.
In summary, all the corrections have the strongest effect
on low carbon mass samples, resulting in the largest dating
uncertainties. Further, such small samples can only be measured
for a short period of time, with reduced stability of the
12C current, additionally worsening the signal-to-noise ratio.
Old low carbon mass samples are most strongly affected by
these two effects because for them the contained number of
14C is particularly low due to the radioactive decay. Among
all uncertainties described, the correction for the procedural
blank typically contributes around 60 %. As an example, for
hypothetical samples with a WIOC mass of 5 or 10 µg, the resulting
uncertainty of the final calibrated ages for 1000-year
old ice would be around ±600 or ±250 yrs and for 8000year
old ice around ±1600 or ±700 yrs, respectively. Hence
by doubling the mass, the uncertainty is reduced by more
than 50 %. We therefore generally discuss dating results only
for sample masses larger than ∼ 10 µg WIOC, which have an
acceptable age uncertainty in the range of 10–20 %.
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C. Uglietti et al.: Radiocarbon dating of glacier ice: overview, optimisation, validation and potential 3097
Figure 2. World map showing the sites from which ice samples were analysed with the 14C method (grey stars): Edziza, Canada, 57.71◦ N
130.63◦ W; GRIP, Greenland, 72.59◦ N, 37.65◦ W, 3230 m a.s.l.; Juvfonne, Norway, 61.68◦ N, 8.35◦ E; Colle Gnifetti, Switzerland, 45.93◦ N,
7.87◦ E; Mt Ortles, Italy, 46.51◦ N, 10.54◦ E; Belukha, Russia, 49.80◦ N, 86.55◦ E; Tsambagarav, Mongolia, 48.66◦ N, 90.86◦ E; Naimonanji,
China 30.45◦ N, 81.54◦ E; Kilimanjaro, Tanzania, 3.06◦ S 37.34◦ E; Quelccaya, Peru, 13.93◦ S, 70.83◦ W; Nevado Illimani, Bolivia,
16.03◦ S, 67.28◦ W; Mercedario, Argentina, 31.97◦ S, 70.12◦ W; Scharffenbergbotnen, Antartica, 74.00◦ S, 11.00◦ W. The average WIOC
concentration in µg kg−1 at each site is indicated with green bubbles.
While calibrating the ages with the OxCal, a sequence
constraint can be applied based on the assumption of a
monotonous increase of age with depth (Bronk Ramsey,
2008). This approach often leads to a reduction of the final
uncertainty, which, however, strongly depends on the sample
resolution with depth; see example in Jenk et al. (2009).
5 Validation of the dating accuracy
5.1 First attempts
Validating the accuracy of the approach described here for
14C dating of ice is a challenging task since it requires ice
samples with known ages, preferentially covering a large age
range. The first validation attempts using ice from Greenland
with the age independently determined by annual layer
counting failed, because WIOC concentrations are an order
of magnitude lower compared to ice from glaciers located
closer to biogenic emission sources (Fig. 2). Large ice
samples were thus needed; nevertheless they still resulted in
small amounts of carbon. Our preparation method is not optimised
for such sample sizes, and the required pooling of
several pieces of ice may have induced a higher procedural
blank. As a result 14C ages tended to be biased by the procedural
blank value (Sigl et al., 2009). 14C ages of the Fiescherhorn
ice core (Jenk et al., 2006) ranged from modern
values to 1000 years, thus reasonably matching the age of
the ice older than AD 1800 obtained by annual layer counting.
For the ice core from Mercedario (31.98◦ S, 70.13◦ W;
6100 m a.s.l.) the deepest core sections show ages of < 550
and 320–1120 cal BP, well in line with a tentative chronology
based on annual layer counting (Sigl et al., 2009). However,
considering the relatively large uncertainty of our method
compared to conventional 14C dating typically derived from
samples with much larger carbon masses and the flatness of
the 14C calibration curve between around 500 and 0 cal BP,
such young samples are not ideal for precise validation. Two
samples from the Illimani ice core, bracketing the AD 1258
volcanic eruption time marker resulted in a combined calibrated
age of AD 1050 ± 70 (1σ) overestimating the expected
age by around 200 years. This would be an acceptable
accuracy if applicable to ice that is several thousand years old
(Sigl et al. 2009).
Overall, these were first indications that the 14C method
gives reliable ages. Since then we have had access to independently
dated ice from the Juvfonne ice patch and the
Quelccaya ice cap, we dated a fly which we discovered in the
Tsambagarav ice core and dated ice cores from the Mt Ortles
glacier, in which a larch leaf was found, altogether allowing
a more robust validation as outlined in the following.
5.2 Recent validation
Juvfonne is a small perennial ice patch in the Jotunheimen
Mountains in central southern Norway (61.68◦ N, 8.35◦ E).
In May 2010, a 30 m-long ice tunnel was excavated, revealing
several dark organic-rich layers up to 5 cm thick containing
organic remains, which were interpreted as previous
ice-patch surfaces and conventionally 14C dated (Nesje
et al., 2012). We received two samples of clear ice adjacent
to the organic-rich layers and a surface sample (JUV 1,
JUV 2, JUV 3, Table 3). The results derived using WIOC
agreed well with the corresponding conventionally dated 14C
ages, ranging between modern and 2900 cal BP (Zapf et al.,
2013). In summer 2015 we collected additional clear ice samwww.the-cryosphere.net/10/3091/2016/
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3098 C. Uglietti et al.: Radiocarbon dating of glacier ice: overview, optimisation, validation and potential
Table 3. Juvfonne samples analysed for method validation. JUV 1, JUV 2 and JUV 3 were ice blocks collected from the 2010 tunnel
(Zapf et al., 2013; Ødegård et al., 2016) and JUV 0 from the 2012 tunnel (Ødegård et al., 2016). To visualise the expected increase in age
with increasing depth of the ice patch, samples are listed in stratigraphic order from top to bottom. Sample JUV 1 was collected between
two separated organic-rich layers (Poz-56952 and Poz-36460). For comparison, an age range between these two layers was calculated
(∗, range between the lower and upper 2σ boundary, respectively). The results from subsamples of the individual ice blocks were averaged to
derive the combined values shown. Uncertainties (1σ) were calculated by error propagation of all analytical uncertainties for the individual
measurements and for the combined values denote the standard error of the unbiased standard deviation. For a graphic display of the
comparison see Fig. 3.
Sample ID AMS Lab. no. WIOC g of F14C 14C age (BP) Cal age
(µg) ice (cal BP)
JUV 3_1 ETH 42845.1.1 55 292 1.124 ± 0.013 −939 ± 93
JUV 3_2 ETH 42847.1.1 43 268 1.094 ± 0.015 −722 ± 110
JUV 3_3 ETH 42849.1.1 47 325 1.155 ± 0.015 −1158 ± 104
JUV 3_4 ETH 43446.1.1 43 208 1.164 ± 0.017 −1220 ± 117
JUV 3 (surface 2010) 1.134 ± 0.017 −1010 ± 120 modern
Organic remains, Poz-37877 0.873 ± 0.003 1095 ± 30 963–1052
JUV 2_1 ETH 43443.1.1 27 215 0.881 ± 0.023 1018 ± 210
JUV 2_2 ETH 43445.1.1 9 171 0.792 ± 0.066 1873 ± 669
JUV 2_3 ETH 43559.1.1 17 257 0.870 ± 0.035 1119 ± 323
JUV 2_4 ETH 45109.1.1 19 219 0.869 ± 0.031 1128 ± 287
JUV 2 (2010) 0.853 ± 0.022 1277 ± 207 965–1368
Organic remains, Poz-56952∗ 0.777 ± 0.003 2027 ± 31
JUV 1_3 ETH 43555.1.1 20 281 0.766 ± 0.029 2141 ± 304
JUV 1_4 ETH 43557.1.1 9 214 0.719 ± 0.064 2650 ± 715
Organic remains, Poz-36460∗ 0.692 ± 0.003 2958 ± 35
JUV 1 (2010) 0.743 ± 0.029 2386 ± 314 2011–2783
Organic remains, age range between the two layers∗ 0.735 ± 0.037 2473 ± 404 2005–3004
Organic remains (plant fragment), Poz-56955 0.486 ± 0.002 5796 ± 33 6561–6656
JUV 0_1 BE 4184.1.1 393 283 0.479 ± 0.015 5913 ± 252
JUV 0_2 BE 4380.1.1 246 298 0.457 ± 0.008 6290 ± 141
JUV 0-A (2015) 0.468 ± 0.014 6099 ± 240 6720–7256
JUV 0_3 BE 4185.1.1 219 208 0.445 ± 0.012 6504 ± 217
JUV 0_4 BE 4381.1.1 182 188 0.442 ± 0.007 6559 ± 127
JUV 0_5 BE 4186.1.1 238 227 0.403 ± 0.012 7301 ± 239
JUV 0_6 BE 4382.1.1 36 184 0.438 ± 0.011 6632 ± 202
JUV 0_7 BE 4187.1.1 262 200 0.404 ± 0.011 7281 ± 219
JUV 0_8 BE 4383.1.1 203 214 0.451 ± 0.013 6397 ± 232
JUV 0-B (2015) 0.431 ± 0.009 6761 ± 168 7476–7785
ples adjacent to a conventionally 14C dated plant fragment
found in an organic-rich layer at the base of a new tunnel
excavated in 2012 and extending deeper into the ice patch
(Ødegård et al., 2016). Four ice blocks were collected and
afterwards subdivided into two subsamples each. Ice block 1
(JUV 0_1 and JUV 0_2) was taken adjacent to the plant
fragment layer, ice block 2 (JUV 0_3 and JUV 0_4), ice
block 3 (JUV 0_5 and JUV 0_6) and ice block 4 (JUV 0_7
and JUV 0_8) at the bottom of the wall, a few centimetres
below the plant fragment layer. JUV 0_1 and JUV 0_2
yielded an average age of 6966 ± 264 cal BP, which is in
good agreement with the age of the plant fragment layer of
6595 ± 47 cal BP, considering the observed increase in age
with increasing depth. Accordingly, the other six samples
collected even further below this organic-rich layer were significantly
older (7630 ± 150 cal BP, Table 3).
Three sections of the ice core from the Quelccaya Summit
Dome drilled in 2003 (QSD, Peruvian Andes, 168.68 m,
13◦56 S, 70◦50 W, 5670 m a.s.l.) were kindly provided by
Lonnie Thompson, Ohio State University. The entire ice core
was dated by annual layer counting, indicating an age of
1800 years at the bottom (Thompson et al., 2013). We intentionally
received the samples without knowing their ages
or depths in order to have the opportunity to perform a “blind
test”. The three sections were not decontaminated as usual,
only rinsed with ultra-pure water, because the amount was
not large enough to remove the outer layer mechanically. As
shown in Fig. 3 (see also Table 4 for the results), the resultThe
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C. Uglietti et al.: Radiocarbon dating of glacier ice: overview, optimisation, validation and potential 3099
0
1000
2000
3000
4000
5000
6000
7000
8000
0 1000 2000 3000 4000 5000 6000 7000 8000
Agesfrom14CWIOCdating(calBP)
Ages from independent dating (cal BP)
Juvfonne
Quellcaya
Tsambagarav
1
2
3
Figure 3. Scatter plot showing the ages obtained with the WIOC
14C method for independently dated ice, including the conventionally
14C dated Juvfonne organic-rich layers (Ødegård et al., 2016),
the 14C dated fly found in the Tsambagarav ice core and the Quelccaya
ice dated by annual layer counting (Thompson et al., 2013).
Error bars denote the 1σ uncertainty. Note that the Juvfonne WIOC
samples and the organic-rich layers were not sampled from the exact
same depth, but adjacent to each other. For the youngest (1)
and oldest (2, containing the plant fragment) the ice for WIOC 14C
analysis was sampled below the layers whereas the third sample (3)
was bracketed by two layers. For (3) the arrow thus indicates the age
range between the lower and upper 2σ boundary of these two layers,
respectively. For (2) the open circle indicates an estimated age for
the according WIOC ice sampling depth based on a fit through all
the conventionally dated organic-rich layers, presented in Ødegård
et al. (2016).
ing calibrated ages agree very well with the ages based on
annual layer counting (L. Thompson, personal communication,
2015).
Recently a number of core segments of the previously
dated Tsambagarav ice core (Herren et al., 2013) were resampled.
In segment 102 a tiny insect (Fig. 4) was found
and immediately separated from the ice matrix. Since it was
small, a conventional 14C analysis was not suitable and instead
the Sunset-AMS system was deployed. The ice section
containing the fly was melted, possible contamination from
carbonates and humic acids were removed by an acid-based
treatment at 40 ◦C (Szidat et al., 2014), the fly was dried,
placed onto a quartz-fibre filter and combusted in the Sunset,
resulting in 13 µg of carbon. The age of 3442 ± 191 cal BP
(BE-5013.1.1) is in perfect agreement with the age of WIOC
from this ice segment of 3495 ± 225 cal BP (Herren et al.,
2013) (Fig. 3).
Additionally, we dated three sections from a set of ice
cores drilled in 2011 on Mt Ortles (see Table 1 for location)
for which a preliminary age of 2612 ± 101 cal BP was
derived by conventional 14C dating of a larch leaf found
at 73.2 m depth (59.60 m w.e., ∼ 1.5 m above the bedrock)
(Gabrielli et al., 2016). Every section was horizontally diFigure
4. Photo of the fly found in segment 102 of the Tsambagarav
ice core. The age of the fly was 3442 ± 191 cal BP, while the
surrounding ice yielded an age of 3495 ± 225 cal BP (photo by Sandra
Brügger).
vided into three subsamples (top, middle and bottom). For
the section at 68.61 m depth (55.08 m w.e., core #1) and the
section at 71.25 m depth (57.94 m w.e., core #3), the ages
obtained for the subsamples were not significantly different
from each other, especially when accounting for the expected
thinning of annual layer thickness at these depths (Fig. 5).
Accordingly the results of the respective subsamples were
combined to derive the most accurate ages for the mid-depths
of these two sections (mean F14C with the estimated 1σ uncertainty
being the standard error of the unbiased standard
deviation). On the contrary the ages of the three subsamples
from the deepest section at 74.13 m (60.54 m w.e., core #3)
significantly increased with depth, implying strong glacier
thinning close to the bedrock (see also Gabrielli et al., 2016).
Our WIOC 14C ages obtained for the Mt Ortles ice core agree
well with the age of the larch leaf, assuming an exponential
increase of age with depth (Fig. 5).
The scatter plot in Fig. 3 summarises the different validation
experiments described above. The results for the Mt Ortles
ice core were not included because larch leaf and WIOC
samples were extracted from depths with significantly different
ages. As shown, within the uncertainties, the 14C ages fall
onto the 1 : 1 line in the age range from ∼ 700–3500 cal BP,
convincingly demonstrating good accuracy of our method.
All validation experiments were performed on low-dust samples,
thus avoiding potential dating bias due to the presence
of dust (Hoffmann, 2016).
6 Applications and current potential of the 14C method
for dating glacier ice
Over the last 10 years the deepest parts of several ice
cores have been dated by applying the presented WIOC
14C method. To illustrate the current potential of the
method with respect to the time period accessible we compiled
five ice core chronologies in Fig. 6. The sites difwww.the-cryosphere.net/10/3091/2016/
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3100 C. Uglietti et al.: Radiocarbon dating of glacier ice: overview, optimisation, validation and potential
Table 4. Quelccaya samples analysed for method validation. Calibrated ages (cal BP) denote the 1σ range. ALC stands for annual layer
counting.
Sample Depth (m) AMS Lab. WIOC F14C 14C age Cal age ALC (yr BP)
no. (µg) (BP) (cal BP)
139–140 144.7–146.8 BE 4336.1.1 15 0.888 ± 0.026 954 ± 237 675–1036 730–788
149–150 155.2–157.3 BE 4335.1.1 24 0.859 ± 0.018 1216 ± 171 1005–1300 1072–1157
157–158 163.9–166.1 BE 4337.1.1 14 0.803 ± 0.025 1761 ± 246 1414–1957 1439–1543
0 1000 2000 3000 4000 5000 6000 7000 8000
62
60
58
56
54
Larch needle
WIOC
Depth(mw.e.)
Calibrated age (cal BP)
Figure 5. Dating of the bottom part of the Ortles ice core. Circles
indicate the ages derived with the WIOC 14C method and the triangle
shows the age of the conventionally 14C dated larch leaf found
in the ice core (Gabrielli et al., 2016). Light grey circles show the
ages obtained for the subsamples. Errors bars represent the 1σ un-
certainty.
fer in recent net annual snow accumulation and ice thickness
(in brackets): Tsambagarav ice cap in the Mongolian
Altai 0.33 m w.e. (72 m) (Herren et al., 2013), Belukha
glacier in the Siberian Altai 0.34 m w.e. (172 m) (Aizen
et al., 2016), Colle Gnifetti glacier in the European Alps
0.46 m w.e. (80 m) (Jenk et al., 2009), Illimani glacier in
the Bolivian Andes 0.58 m w.e. (138.7 m) (Kellerhals et al.,
2010), Mt Ortles glacier 0.85 m w.e. (75 m) (Gabrielli et al.,
2016). All of these are cold glaciers and frozen to the bedrock
with the exception of the Mt Ortles glacier, which is polythermal
and experienced a recent acceleration of glacier flow
due to sustained atmospheric warming over the past decades
(Gabrielli et al., 2016). To derive a continuous age–depth
relationship, a two-parameter flow model (Bolzan, 1985;
Thompson et al., 1990) was applied for Colle Gniffetti (Jenk
et al., 2009), Illimani (Kellerhals et at., 2010) and here also
for the core from Belukka using the data presented in Aizen
et al. (2016). A different approach, discussed below, was implemented
for the ice cores from the Tsambagarav ice cap
(Herren et al., 2013) and the glacier on Mt Ortles (see also
Gabrielli et al., 2016). The two-parameter model is based on
1 10 100 1000 10 000
160
140
120
100
80
60
40
20
0
RH
14
C WIOC
Depth(mw.e.)*
Ortles*
Colle Gnifetti
Illimani
Tsambagarav
Belukha
Depth(mw.e.)
Calibrated ages before year 2000
140
120
100
80
60
40
20
0
-20
Figure 6. Compilation of age–depth relationships for five different
ice cores, highlighting the importance of the WIOC 14C dating for
obtaining continuous chronologies and constraining the very specific
glaciological conditions and settings of each site. For simplicity
only reference horizons and 14C dates were included. Grey triangles
indicate reference horizons (RH) and red circles the 14C-WIOC
ages both plotted with 1σ uncertainties (smaller than the symbol
size in some cases). For the Mt Ortles core, 210Pb dated horizons
with a larger uncertainty were used as RH due to the lack of absolute
time markers prior to 1958; the grey triangle at 57.8 m w.e. depth
is the conventional 14C age of the larch leaf. Grey shaded areas
represent the 1σ range of the respective fit for retrieving a continuous
age–depth relationship. For sample details and the fitting
approaches applied, see main text and Table 1. References to the
original data are summarised in Table 1. Note that for better visibility
(avoiding overlap with Tsambagarav and Colle Gnifetti) the
curve for the Mt Ortles glacier was shifted down by 20 m and refers
to the right-hand y axis (∗).
a simple analytical expression for the decrease of the annual
layer thickness L(z) (m w.e.) with depth:
L(z) = b 1 −
z
H
p+1
,
where z is depth (m w.e.), H the glacier thickness (m w.e.),
b the annual accumulation (m w.e.) and p a thinning parameter
(dimensionless). The age T (z) as a function of depth can
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C. Uglietti et al.: Radiocarbon dating of glacier ice: overview, optimisation, validation and potential 3101
be calculated when the inverse layer thickness is integrated
over depth:
T(z) =
dz
L(z)
=
1
b
1 −
z
H
−p−1
dz.
Solving the integral and setting the age at the surface to be
T (0) = 0, the final age–depth relation is obtained:
T(z) =
H
bp
1 −
z
H
−p
− 1 .
The thinning rate (vertical strain rate) is the first derivative of
the layer thickness:
L(z) =
dL(z)
dz
= −
b(p + 1)
H
1 −
z
H
p
.
The model has two degrees of freedom, the net annual accumulation
rate b and the thinning parameter p, both of which
are assumed to be constant over time. This allows us to fit
the model by a least squares approach through the available
reference horizons if the glacier thickness H is known (if
drilled to the bedrock) or can be reasonably well estimated
(e.g. from radar sounding). In order to not overweigh the
data from the deepest horizons, the model is fitted using the
logarithms of the age values. For the ice cores from Colle
Gnifetti (Jenk et al., 2009), Illimani (Kellerhals et al., 2010)
and Belukha (Aizen et al., 2016), these ages were based on
annual layer counting, identification of reference horizons
(radioactive fallout and well-known volcanic eruptions) and
14C dates. The data are summarised in Table 1. In Fig. 6, only
reference horizons and 14C dates were included for simplifi-
cation.
In summary, a reasonable fit was achieved for these three
glaciers and the derived annual net accumulations (Colle
Gnifetti 0.45 ± 0.03 m w.e., Belukha 0.36 ± 0.03 m w.e., Illimani
0.57 ± 0.13 m w.e.) are comparable with the values
previously published (see above), which were determined either
by surface measurements or were estimated based on
ALC or/and the uppermost age horizons only (e.g. nuclear
fallout peak), thereby accounting for the (slight) layer thinning
occurring in these uppermost few metres (Nye, 1963).
Since the assumption of constant accumulation (b) and a constant
thinning parameter (p) over time/with depth is likely
only true in a first-order approximation, it is thus no surprise
that the two-parameter model may fail to result in a
reasonable fit within the derived age uncertainties. In such a
case these two underlying assumptions should then be investigated
more thoroughly, e.g. as done for the ice cores
from Tsambagarav and Mt Ortles. Whereas Tsambagarav is
also a cold glacier, Mt Ortles is polythermal. For Tsambagarav,
a good fit can be achieved if an additional degree of
freedom is given to account for variations in the net accumulation
rate while p is fixed to the initially derived value,
suggesting significant changes in the accumulation rate over
time. This is supported by the fact that the resulting strong
variation in net accumulation is consistent with precipitation
changes in the Altai derived from lake sediment studies (Herren
et al., 2013). In contrast, a reasonable fit for the Mt Ortles
ice core can only be obtained if the thinning parameter p is
allowed to increase with depth, while the annual net accumulation
is assumed to be constant over time (i.e. b fixed
to the value defined by the stake measurements and surface
layers). This points to an exceptionally strong thinning. The
Mt Ortles glacier is polythermal with temperate conditions in
the upper part and relatively warm ice with −2.8 ◦C near the
bedrock. We hypothesise that the faster horizontal velocity
of the warm ice causes exceptional horizontal stress (internal
horizontal deformation) on the ice frozen to the bedrock, resulting
in stronger thinning. In both cases, a purely empirical
approach of fitting the age horizons was chosen to establish
the age–depth relationship. Note that, due to the lack of absolute
time markers prior to 1958, 210Pb dated horizons were
used for Mt Ortles. For Tsambagarav a combination of different
polynomial functions was used (Herren et al., 2013),
whereas a slightly more sophisticated approach by means of
a Monte Carlo simulation was applied for Mt Ortles, allowing
an objective uncertainty estimate for each depth defined
by the density of dating horizons and their individual uncertainty
(Gabrielli et al., 2016). These purely empirical approaches
are justified given the high confidence assigned to
the determined ages for the dated horizons.
As shown in Fig. 6, the time period dated with 14C,
ranges from 200 to more than 10 000 years. Due to their
uncertainty, 14C ages derived by our method cannot compete
with the conventional methods for dating ice that is
only a few centuries old. The strength of 14C dating using
WIOC is that it allows us to obtain absolute ages from basically
every piece of ice from cold and polythermal ice bodies.
This is especially valuable for glaciers not containing
the last glacial/interglacial transition, such as Tsambagarav
and Mt Ortles, since in such cases not even climate wiggle
matching of the transition signal with other dated archives
is possible. In any case an absolute dating method is preferable
to wiggle matching, which is not necessarily reliable.
For example, a depletion in δ18O, presumably indicating the
Last Glacial Maximum (LGM)/Holocene transition, might
not always be a true atmospheric signal, but can be caused
by unknown mechanisms potentially happening close to the
bedrock (Jenk et al., 2009; Wagenbach et al., 2012). All of
the five presented examples show strong thinning towards
the bedrock with the oldest ages having been obtained in the
range of 10 000 years. Because of the strong thinning, the 14C
ages of the deepest samples represent a strongly mixed age of
ice with a large age distribution. In these cases, the age limit
was thus not determined by the 14C half-life of 5730 years
(Godwin, 1962), but by the achievable spatial depth resolution
since some hundred grams of ice are required.
Since an absolute WIOC mass of ∼ 10 µg is needed to
achieve a 14C dating with reasonably low uncertainty, the
overall applicability of the method essentially depends on
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3102 C. Uglietti et al.: Radiocarbon dating of glacier ice: overview, optimisation, validation and potential
the WIOC concentration in the ice and the ice mass available.
Figure 2 summarises WIOC concentrations determined in ice
from various locations around the globe. In general, midlatitude
and low-latitude glaciers contain sufficient WIOC from
21 to 295 µg kg−1, allowing dating with less than 1 kg of
ice. The highest concentration was found at the Juvfonne ice
patch, which is small, located at a low elevation and, therefore,
by far the closest to biogenic emission sources. WIOC
concentrations might be further elevated due to meltwater
and superimposed ice formation, enriching water-insoluble
particles present in the surface layer. The lowest concentrations
of only 2 to 15 µg kg−1 WIOC were observed in polar
snow and ice from Greenland and Antarctica. For this concentration
range a reliable dating is impossible with the current
method capability.
7 Conclusions
Since the introduction about 10 years ago of the 14C dating
technique for glacier ice, in which the WIOC fraction of
carbonaceous aerosol particles embedded in the ice matrix
was used, major improvements in separating the OC from
the EC fraction and in AMS technology have been achieved.
The new configuration with direct coupling of a commercial
thermo-optical OC/EC analyser to the gas ion source of the
MICADAS AMS via its gas introduction interface has two
major advantages. First, the measurement time was significantly
reduced to approximately 35 min per sample. Second,
the implemented automated protocol allows for a controlled
routine analysis with high reproducibility and a stable blank,
thereby increasing the overall precision.
The presented WIOC 14C dating method was validated by
determining the age of independently dated ice samples. It
principally allows absolute and accurate dating of any piece
of ice containing sufficient WIOC. With the current set-up,
the age of samples with a minimum of ∼ 10 µg WIOC can be
determined a with satisfying precision of about 10 to 20 %,
depending on the age. This requires about 300 to 800 g of ice
considering a mass loss of 20–30 % during surface decontamination
and the WIOC concentrations typically found in
midlatitude and low-latitude glaciers. Dating polar ice with
satisfactory age uncertainties is still not possible since WIOC
concentrations are around 1 order of magnitude lower. This
would require further reduction of the procedural blank for
such samples requiring larger ice volumes which potentially
could be achieved by an additional, specifically designed
sample preparation set-up for such kind of samples.
The 14C method is suitable for dating ice with ages from
200 to more than 10 000 years. Whereas for ice that is a few
centuries old, the conventional dating methods are typically
higher in precision and the WIOC 14C method presents the
only option for obtaining reliable continuous timescales for
the older and deeper ice core sections of mountain glaciers.
This is not only crucial for interpreting the embedded environmental
and climatic history, but gives additional insight
into glacier flow dynamics close to the bedrock as demonstrated
by the age–depth scales derived from 14C dating of
ice cores from various midlatitude and low-latitude glaciers.
It can also reveal information about the time of glacier for-
mation.
Author contributions. Manuscript written by Chiara Uglietti,
Theo Manuel Jenk and Margit Schwikowski with editing by
Sönke Szidat. Sample preparation and 14C measurements performed
by Chiara Uglietti, Alexander Zapf and Michael Sigl
with expert supervision from Gary Salazar, Sönke Szidat and
Theo Manuel Jenk.
Acknowledgements. This work was supported by the Swiss National
Science Foundation (200020_144388) and by the Oeschger
Centre for Climate Change Research of the University of Bern. We
thank Sandra Brügger and Edith Vogel for sample preparation of
the fly extracted from the Tsambagarav ice core.
Edited by: S. Hou
Reviewed by: three anonymous referees
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