Introduction
The deposition of light-absorbing carbonaceous particles emitted by the
incomplete combustion of biomass and fossil fuel can decrease the albedo of
Arctic snow- and ice-covered surfaces, thereby amplifying high-latitude
warming driven by the buildup of greenhouse gas emissions (AMAP, 2011; Bond
et al., 2013). The widely used expression “black carbon” (BC) designates the
insoluble, refractory fraction of these aerosols that is largely made of
graphitic elemental carbon and strongly absorbs light at visible to
near-infrared wavelengths (Petzold et al., 2013). Along with sulfate
(SO42-), BC is one of the main short-lived climate pollutants
being targeted for mitigation and control under multinational legal
agreements (Quinn et al., 2008; AMAP, 2015).
In order to evaluate how past and future BC emissions have affected, and
will affect, climate forcing in the Arctic, global atmospheric climate
models can be used to simulate the transport and deposition of BC aerosols
in this region (Koch et al., 2011; Skeie et al., 2011; Lee et al., 2013;
Jiao and Flanner, 2016). At present, simulated BC dispersion suffers from
large biases, either positive or negative, compared with observational data
on BC in Arctic air and snow (Jiao et al., 2014). Validating model
simulations is difficult because of the scarcity of such observations across
the Arctic. Direct monitoring of atmospheric BC is currently limited to a few
decades and has only been carried out at a few stations (Hirdman et al., 2010; Gong et al.,
2010), and geographic surveys of BC in snow and ice are rare and difficult to conduct
over the vast Arctic region (e.g., Doherty et al., 2010).
Ice cores drilled from the accumulation area of glaciers and ice caps can be
used as surrogates for direct atmospheric observations, as they contain
archives of BC and other aerosol species deposited in snow over many
centuries (McConnell, 2010). At present, ice-core records of BC deposition
in the Arctic region are only available from Greenland (McConnell et al.,
2007; McConnell and Edwards, 2008; Zennaro et al., 2014; Sigl et al., 2015)
and from Svalbard, Norway (Ruppel et al., 2014). Here, for the first time, we
present a historical record of BC deposition in the Canadian Arctic,
developed from a core drilled on Devon Island ice cap, and roughly spanning the
years from 1735 to 1992. The Devon ice cap BC record presents some
striking differences from Greenland ice-core records of rBC concentrations
developed using the same methods. We discuss the possible reasons for these
differences, and consider the implications with respect to regional BC
transport and deposition patterns in the Arctic region.
Location map of the Canadian Arctic Archipelago (a),
with enlargement of Devon ice cap (b). The location of the various ice-core
sites mentioned in the text are shown. Sites A to E refer to the shallow
core array of Colgan and Sharp (2008). Elevation contours on Devon ice cap
are spaced at 100 m above sea level.
Materials and methods
Core sampling and analyses
The DV99.1 core was recovered in 0.4–1.1 m long increments (average 0.9 m),
with a diameter of 9.8 cm. The uppermost 2.8 m of the core was made of
crumbly firn, and was not preserved at the time of drilling. The
solid-state DC electrical conductivity (EC) of the core was measured continuously in the
field using a hand-held system with parallel electrodes
(Icefield Instruments Inc., Whitehorse, Canada), as described in Zheng et al. (1998).
The EC profiling started at a depth of 12.38 m, because sections
of cores above this were of brittle firn that provided inadequate electrode
contact for the hand-held instrument. The core was shipped and stored in
freezers at the GSC ice-core laboratory in Ottawa. There, it was sampled at
5–20 cm resolution for the subsequent determination of stable oxygen isotope ratios
(δ18O) by mass spectrometry at the University of Copenhagen.
Later, 57 discrete subsamples from depths below 29 m were analyzed for lead
(Pb) and other trace metals, as reported in Zheng et al. (2007). The
remaining cores were stored frozen (-20 ∘C) inside sealed
polyethylene bags, until archived core segments between 2.8 and 48 m depths
were selected for this study and shipped, still frozen, to Curtin University
in Australia for BC analyses. These combined core segments were estimated to
span > 250 years, as explained below.
Sample preparation and analysis was conducted between 6 and 11 December 2012 at
the Trace Research Clean Environmental facility at Curtin University. The
facility consists of a large class 100 space containing multiple class 10
laboratory modules including a -20 ∘C walk-in freezer within
a general lab space (also class 10). The space was specifically designed for
trace metal and particle work on ice cores (e.g., Burn-Nunes et al., 2011,
Ellis et al., 2015, 2016; Tuohy et al., 2015; Vallelonga et al., 2017). The
DV99.1 core sections were cut into subsamples with a ∼2.5×2.5 cm cross section, which were processed in an ice-core melter
coupled to a continuous flow analysis (CFA) system (see Supplement, Fig. S1). Ice-core preparation was carried out in the walk-in freezer, while
processing in the CFA system was conducted in the general lab class 100
space. The CFA melter system was similar to that described by McConnell et al. (2002)
with the exception that the ice-core melter head was made from
aluminum. The method used to quantify BC in the ice core was the same as that
used for the analysis of Greenland and Antarctic cores (see Bisiaux et
al., 2012; McConnell et al., 2007; McConnell and Edwards, 2008; Zennaro et
al., 2014). Meltwater from the CFA system was aerosolized and desolvated
with a U5000AT ultrasonic nebulizer (CETAC Technologies, Omaha, NE, USA) and
injected into a single-particle intracavity laser-induced incandescence
photometer (Schwarz et al., 2010; SP2, Droplet Measurement Technologies,
Boulder, CO), which measured the mass concentration of BC particles in the
meltwater flow. Instrumental settings are given in the Supplement (Table S1). Following Petzold et al. (2013), we refer to the BC fraction measured
by this method as “refractory BC” (rBC), reported here in mass concentration units of
ng g-1.
Details of the three Devon ice cap cores used in this
study. MAAT refers to mean annual surface air temperature. See text for specific
references to published data.
Core
Lat.
Lon.
Max. depth
Approx. elevation
MAAT
Annual accum.
(N)
(W)
(m)
(m a.s.l.)
(∘C)
(m H2O)
Parameters measured
DV98.3
75.34∘
82.14∘
302
1930
-12
0.25–0.28
δ18O, radioactivity,
major ions, and trace metals
DV99.1
75.32∘
81.64∘
169
1903
no data
0.16
δ18O, melt features, EC, and rBC
DV2000
75.34∘
82.14∘
64
1930
-12
0.25–0.28
δ18O, radioactivity, and trace metals
On each day of the analysis, a log journal was created. The length of every
piece of the DV99.1 core was carefully measured prior to analysis. During
CFA, the time of each break between two ice-core pieces was recorded, making
it possible to reconcile the rBC record of each piece based on the
time–depth log. The flow rate of the CFA to the nebulizer was controlled by
oversupplying a <1 mL debubbling vessel with excess water, which allowed
the instrument to maintain a very constant flow rate. External calibration
of the SP2 nebulizer system was achieved using eight standards of 100 %
carbon black pigment (MIS Ink Supply, Eboni-6K; Fig. S2) spanning a
concentration range of 0–20 ng g-1. The standards were analyzed each
day before and after ice-core analysis and the results were compared to
assess the stability, reproducibility, and measurement uncertainty of the
SP2. Additional details and calibration curves (Figs. S3–S6) are provided in
the Supplement, and potential sources of uncertainty in the results are
discussed in Sect. 3.3.
To compare the DV99.1 record of rBC with that of other deposited aerosol
species, we used glaciochemical data obtained from two other cores drilled
from the summit area of Devon ice cap in 1998 (core DV98.3) and 2000 (core
DV2000) (Fig. 1; Table 1). The DV98.3 core was sampled continuously and
analyzed for eight major ionic species by ion chromatography, as described
in Kinnard et al. (2006). In this study, we used SO42-, sodium
(Na+), calcium (Ca2+), potassium (K+), and ammonium
(NH4+) data obtained from the top 85 m of the core, which had been
sampled at 3–12 cm resolution. The non-sea salt fraction of sulfur
(nssS) was estimated from Na2+ using the mean surface seawater
composition of Pilson (2012), whilst the biomass burning fraction (BB) of
K+ was estimated from Na2+ and Ca2+ as follows:
[K+]BB=[K+]-(0.038×[Na+])-(0.04×[Ca2+]), following Legrand et al. (2016). The DV2000 core was drilled
at the same site as the DV98.3 core, and was analyzed for Pb and other
metals, as reported in Shotyk et al. (2005). However, the remaining archived volume
from cores DV98.3 and DV2000 was insufficient to carry out rBC
analyses, which is why core DV99.1 was used for this purpose.
Age models
Annual layers are not easily resolved in cores from Canadian Arctic ice
caps, partly owing to relatively low A˙, but also due to the effects of
wind and/or summer surface melt. Therefore, age models developed for these
cores are commonly based on a variety of alternative methods. For the DV98.3
and DV99.1 cores, an ice-flow model (Dansgaard and Johnsen, 1969) was used,
constrained by the total ice thickness obtained from ice-radar measurements
or from borehole depths, and by the estimated A˙ at each coring site.
For the DV98.3 core, the age model was further constrained by approximate
layer counting using δ18O and glaciochemical data at shallow
depths. At greater depths, reference horizons from bomb
radioactive fallout (1963; Pinglot et al., 2003) and historical volcanic
eruptions were used, including the Laki volcanic eruption in Iceland, in 1783 (All given dates are
C.E.), which is one of most recognizable historical volcanic signals
recorded in EC and/or SO42- records of other Canadian Arctic ice
caps (e.g., Zheng et al., 1998; Goto-Azuma et al., 2002). The age model in
the upper 48 m of the DV99.1 core was constrained using a reference horizon
provided by a large EC (acidity) spike at a depth of 42.60 m (29.56 m ice
equivalent), which was attributed to the 1783 Laki eruption (Fig. 2). This
model gives an estimated maximum age of 1735 for the section of the DV99.1
core used in the present study, and the last year in the record is 1992. The
age model also gave an acceptable agreement between profiles of various
measured parameters in the DV98.3 and DV99.1 cores (Figs. S7–S8). The DV2000
core was drilled at the same site as the DV98.3 core and used the same age
model. The two cores were correlated using measurements in the DV2000 that
allowed for the identification of radioactive layers dated to 1958 (16.5 m depth)
and 1963 (13.5 m depth) (Krachler et al., 2005). The DV2000 core was
estimated to extend back to 1842.
Age models for parts of the DV98.3 and DV99.1 cores from
Devon ice cap. The error bars on the age curve relative to ice-equivalent
depths (red) bracket the 95 % confidence interval on the estimated age
for discrete depths, as established from Monte Carlo simulations (see text).
Using the Laki 1783 reference layer, the estimated A˙ at the DV99.1
site is 0.14 m ice a-1 (0.16 m H2O a-1) which is lower than
at the ice cap summit (∼0.25–0.28 m H2O a-1) or at
sites elsewhere in the Devon ice cap accumulation zone (0.17–0.25 m H2O a-1; Colgan and Sharp, 2008).
The most likely explanation for this is partial
scouring of winter snow layers by downslope winds at the DV99.1 site, as
also observed on parts of the Agassiz ice cap (Fisher et al., 1983). This is
supported by a comparison of the δ18O measurements in the
DV99.1 and DV98.3 cores, which shows that δ18O variations in
the DV99.1 core are truncated of their most negative (“coldest”) values
relative to the DV98.3 core (Fig. S9). An estimate of the amount of snow
lost by wind scouring at the DV99.1 site can be made from the difference in
the amplitude of the δ18O data at the DV98.3 and DV99.1 sites,
and from A˙ at the DV98.3 site, following Fisher and Koerner (1988).
The calculation suggests that ∼40 %–45 % of the annual snow
accumulation is removed by wind at this site, compared to the summit of
Devon ice cap.
Quantifying uncertainties in the rBC record
Analyses of rBC in the DV99.1 core were performed at high depth resolution,
producing ∼55–80 data points per meter over most of the
core's length. The data were subsequently averaged over discrete depth
increments equivalent to ∼1-year and ∼10-year
intervals, respectively, based on the core's age model. In this paper,
annually averaged figures are used for illustrative purposes only, as
individual years can not be confidently resolved in the DV99.1 core.
Down-core variations of rBC in the ice core are the result of a combination
of processes, including temporal changes in atmospheric deposition rates
(fluxes, abbreviated F), spatial variations of deposition of aerosols in
snow, and post-depositional modifications (e.g., by wind scouring or summer
surface melt). Additional uncertainties in the rBC data come from the age
model of the ice core (Fig. 2) and from limitations of the analytical
method.
The largest uncertainty with regards to the rBC analysis is due to the
nebulization/desolvation step before the SP2 analysis. At the time of this
study we had adopted nebulizer/desolvation systems used as a front end to
inductively coupled plasma mass spectrometers (ICP-MS). These systems are
designed to deliver appropriate aerosol size distributions for analysis in
the ICP-MS. Schwarz et al. (2012) and Wendl et al. (2014) report rBC
size-dependent losses during nebulization/desolvation for several types of
nebulizer desolvation systems. The study found that the system used in this
investigation had a poor transport efficiency for rBC particles with a
volume equivalent diameter > 500 nm. Hence, rBC data from the
DV99.1 core should be considered with this limitation in mind (see Sect. 4.2 for a
discussion). Other published ice-core data sets from Greenland (e.g.,
Mc Connell et al., 2007) are also subject to this limitation, but are at least
comparable. Further research is required to assess the true size
distribution of rBC deposition to the Devon ice cap and other Arctic sites.
Uncertainties in the DV99.1 age model are primarily due to the potential
identification error of the Laki 1783 layer in the EC profile, and to
interannual variations in A˙ at the ice-coring site. The relationship
between true depth and ice-equivalent depth is nearly linear in the DV99.1
core down to 48 m, which suggests a steady firn densification rate over the
corresponding time interval, with no signs of dynamically induced changes in
the vertical strain rate. For the 1783 layer, we conservatively assumed a
possible dating error of ±5 years, corresponding to a depth
registration error of ∼±1 m at the 42.6 m EC peak. The
interannual variability in A˙ was estimated from an array of shallow
cores (Colgan and Sharp, 2008) and from winter mass balance measurements
since 1961 (data available through the World Glacier Monitoring Service).
This information was used in a Monte Carlo simulation in
MATLAB™ with 1000 realizations to compute confidence limits
(CL) on the decadally averaged rBC data. Briefly, a constrained random walk
algorithm was used to estimate the probabilistic distribution of the true
age at any depth in the core from the surface down to the Laki 1783 layer
(Kinnard et al., 2006). Interannual variations in A˙ were considered
to behave as a stationary, autoregressive blue noise process with a lag-one
serial autocorrelation coefficient of -0.5 to -0.3, based on empirical data
presented by Fisher et al. (1985). Thus, a population of 1000 alternative age
models was generated. From each of these, 10-year averages of the rBC
data were computed, and 95 % CL were calculated for the geometric mean
rBC concentration in each decade (Fig. S10). Expressed as a coefficient of
variation (CV), the estimated uncertainty on the decadally averaged rBC
concentrations that arise from age model errors varies from 3 %–23 %
(median 6 %), depending on the decade considered.
Profiles of physical properties and rBC in the top
48–50 m of the DV99.1 ice core. (a) Firn density and estimated mean
annual layer thickness. (b) Frequency of discrete ice layers
(>3 mm thick) per core section. (c) Solid-state electrical conductivity (EC)
profile of the core from 12.8 to 50 m depth, smoothed to a vertical
resolution of ∼1 cm. The EC peak attributed to acidic fallout
from the Laki 1783 eruption is labeled. (d, e) rBC concentrations
plotted on linear and log scales. The bold red line is a 50-point
(∼1 m) moving average.
Environmental changes on Devon ice cap, 1740–1999,
recorded in three cores from the summit region (DV98.3, DV99.1, and DV2000).
(a) rBC concentrations in the DV99.1 core; (b) Pb concentrations in the
DV99.1 core (∼1740–1840) and DV2000 core (1840–2000); (c)
SO42- in the DV98.3 core; and (d) volumetric percentage of icy melt features
in the DV99.1 core due to surface summer melt. Data are presented in
∼1-year, 5-year and/or 10-year averages. For panels (a) to (c),
10-year geometric mean values of the data are also plotted in red on
separate scales (right). The shaded grey bar identifies the Laki 1783
isochron used to correlate the different cores. The width of the bar denotes
the maximum dating uncertainty at the corresponding depths in these cores.
The Pb data are from Shotyk et al. (2005) and Zheng et al. (2007), the
SO42- data are from Kinnard et al. (2006), and the melt feature data are from Fisher et al. (2012).
The spatial variability of BC deposition on Canadian Arctic ice caps is
unknown. An estimate for Devon ice cap can be made from major ion analyses
on shallow cores (Colgan and Sharp, 2008; Fig. 1). In these cores, the
spatial CV on the annual SO42- deposition averages 42 %
(range 17 %–100 %) over a period of ∼40 years. Here, we make
the assumption that the deposition of BC on Devon ice cap shares the same
spatial variability as SO42-, an aerosol species which, like BC
but unlike others such as nitrate (NO3-), is not subject to
re-emission from snow to air. While the spatial variability may be large on
an annual basis, Monte Carlo simulation results show that averaging the rBC
data over 10-year intervals reduces its effect on the geometric mean rBC
uncertainty to a few percent (CV) in any decade (Fig. S10). The potential
impact of post-depositional modifications in the rBC record is discussed
in Sect. 4.2.
Results and discussion
The DV99.1 record of rBC
The depth profile of rBC measured in the DV99.1 core is shown in Fig. 3. The
probability distribution of rBC concentrations is approximately log-normal
(Fig. S11); therefore, we use both the arithmetic and geometric means
(μ, μg), as descriptive metrics for these data. Over the
entire core length, rBC concentrations average 1.8±3.9 ng g-1
(μg=0.8 ng g-1) with a maximum of 74.0 ng g-1. The
mean rBC concentration is approximately constant between 42 and 15 m depths,
and decreases gradually at shallower depths to reach ∼1.0 ng g-1 (μg=0.5 ng g-1) in the uppermost meter of core.
Concentrations below 42 m show a comparatively larger variability and a
greater range of values (Fig. S12).
In Greenland cores, rBC deposition rose in the 1880s, peaked in the
1910s–1920s, and decreased thereafter (McConnell et al., 2007), in step with
historical changes in coal burning BC emissions from North America and
Europe (Novakov et al., 2003; Bond et al., 2007; Lamarque et al., 2010). In
south-central Greenland, the early 20th century rise in rBC and nssS was
also accompanied by increased deposition of Pb and other trace metals
(McConnell and Edwards, 2008). Measurements from the DV98.3 and DV2000 ice
cores (Fig. 4) show that Devon ice cap also experienced increased
atmospheric deposition of SO42- and Pb and during the 20th
century, peaking between the 1960s and 1980s. This was followed by a decline,
consistent with trends in midlatitude anthropogenic emissions from fossil
fuel combustion. However, unlike in Greenland, the DV99.1 core shows no
large, sustained increase in rBC concentration concomitant with that of
SO42- or Pb. There is a modest rise in mean rBC concentrations
from the early 1800s to the mid-20th century, but it is much more gradual
and of lesser magnitude than the rBC rise observed in ice-core records from
Greenland; the relative timing and the magnitude of these increases also
differ between core sites (Figs. 5 and 6). In the DV99.1 ice core, the
highest mean rBC concentrations for the 20th century occur in the decade from
1960 to 70 (μ=4.7 ng g-1, μg=1.7 ng g-1), but
these are not unprecedented, and comparable mean concentrations occur in the
earliest part of the record, in the decade from 1780 to 1790 (Fig. S12).
The record of atmospheric rBC and non-sea salt sulfur
(nssS) deposition on Devon ice cap over the period from 1800 to 2000 compared with
similar records developed at various sites in Greenland by identical or
almost identical methods. Full lines are rBC; stippled lines are nssS. Data
from Summit, D4, ACT2, and Humboldt are sourced from McConnell et al. (2007) and Koch et al. (2011), and data from NEEM are sourced from Zennaro et al. (2014) and Sigl et al. (2015).
Also shown is the location of the ice-core record of elemental carbon
deposition developed from Holtedahlfonna, Svalbard, by Ruppel et al. (2014),
in addition to other sites (Alert, Dye 2) mentioned in the text.
(a) The DV99.1 record of atmospheric rBC deposition since
1800 compared with other records developed from sites in Greenland
identified in Fig. 5. All records are presented in 1-year averages. (b) As
in (a) but for records of non-sea salt sulfur (nssS). Two volcanic eruption
isochron used for correlation in the Greenland cores are highlighted.
The DV99.1 record also shows a pronounced decline in rBC concentration in
the late 20th century, but it occurs after the 1960s, which is later than in
most Greenland cores, except at Humboldt (Figs. 5 and 6). However, this difference in
timing could be due to uncertainties in the DV99.1 chronology
compared to that of annually dated Greenland cores. The DV99.1 mean rBC
concentrations over the period from 1960 to 1990 (μ=0.6–1.0 ng g-1;
μg=0.3–0.5 ng g-1) are lower than in the early modern
industrial period (early 19th century; μ=1.0–3.0 ng g-1;
μg=0.7–1.6 ng g-1). The only Greenland ice core in which
a similar situation occurs is from the ACT2 site (66∘ N, Fig. 5).
Neither the winter mass balance measurements, nor the reconstructed interannual
changes in A˙ on Devon ice cap (Colgan and Sharp, 2008) show any
sustained long-term trend since the early 1960s; therefore, the decrease in the rBC
concentration in the DV99.1 core during this period can not be
ascribed to changing precipitation rates on the ice cap. It seems more
likely that the decrease is at least in part due to a declining burden of
atmospheric BC in the Canadian High Arctic since the 1960s (Gong et al.,
2010). However, there are several methodological, site-specific and
regional-scale factors that must be taken into account when interpreting the
DV99.1 rBC record. These are discussed below.
Methodological and site-specific factors
Observations of atmospheric BC at Alert on Ellesmere Island (82∘ N, Fig. 1)
show a seasonal cycle with airborne concentrations peaking during
winter and spring months (December–March) and declining to their minimum in
summer and early autumn months (June–September) (Gong et al., 2010). Most BC
deposition in snow is thought to occur in spring and summer, when increased
cloudiness promotes in-cloud scavenging and wet deposition of particles containing BC
(Garrett et al., 2011; Browse et al., 2012; Shen et al., 2017). In
the interior of the Greenland ice sheet, the seasonal cycle of BC deposition
is well-preserved in snow and firn layers (e.g., McConnell et al., 2007).
This is not the case at the DV99.1 core site on Devon Island. Even in the
uppermost part of the core, where some seasonal δ18O variations
can be detected, there is no recognizable seasonal pattern of rBC
concentration peaks (Fig. S13). This is likely the result of the combined
effects of wind scouring/mixing of surface snow (as described earlier) and
of summer surface melt. Therefore, the question arises as to whether such processes
could also have obliterated or masked a 20th century anthropogenic
signal in the DV99.1 rBC record.
Simulation of the effects of wind scouring of snow
on the preservation of an anthropogenic signal of rBC deposition in a synthetic
ice-core times series of rBC spanning the period from 1800 to 1990. (a) The
synthetic series, with a pseudo-seasonal cycle superimposed on the
interdecadal baseline trend observed in the Greenland D4 record (McConnell
et al., 2007). (b) The synthetic series after randomly truncating the
amplitude of all winter deposition peaks (November–March) by 30 %–60 %. The
bold red line in both panels is a 5-year running geometric mean.
Maps of residence time probability for air arriving at
(a) Devon ice cap and (b) Summit, Greenland over the period from 1948 to 1999,
computed using HYSPLIT4. Air residence probability densities were normalized
to a scale of 0–1, and were spatially detrended by multiplying the original
residence time grids (in hours) by the distance between each grid point and
the coring site. This effectively removes the concentric increase in
probability density near the back-trajectory start point (Ashbaugh et al.,
1985). The spatial resolution of the grid is 200×200 km.
The seasonally resolved ice-core record from the D4 site in Greenland
(71∘ N; Fig. 5) shows that during the historical period of
enhanced anthropogenic BC pollution in the Arctic, from the late 19th to mid-20th
centuries, rBC deposition increased in both summer and winter (McConnell
et al., 2007). If the Canadian High Arctic was impacted by airborne BC
pollution in a similar way, one would expect to find a marked increase in
rBC concentrations in the DV99.1 core during the early 20th century, even if
winter snow layers were scoured away by wind. To verify this, we performed a
simple simulation in which we generated synthetic time series of rBC
deposition spanning the period from 1800 to 1990, with a seasonal cycle superimposed
on baseline interdecadal variations similar to those observed in the
Greenland D4 ice-core record. Winter rBC deposition peaks in the series were
represented using a log-Gaussian function, and their amplitude was allowed
to vary from year to year to produce a range of temporal variations
comparable to, or lower than, that seen in the Greenland D4 core. Winter
deposition peaks were then randomly truncated by 30 %–60 % (mean 45 %)
to simulate the effects of wind scouring on the record, and 5-year running
means were computed from the resulting data. Smoothing was used to
simulate the effects of post-depositional snow layer mixing by wind. Results
of these experiments show that even if the wintertime rBC deposition peaks
between November and May were largely truncated by wind, the low-frequency
baseline variation would still persist, and should be recognizable above the
remaining interannual signal variance (Fig. 7). Therefore, it seems unlikely
that wind scouring would completely obliterate this rBC signal in the DV99.1
record, unless the amplitude of the seasonal cycle of atmospheric BC
deposition on Devon ice cap is much lower than observed at Alert or in
Greenland (Gong et al., 2010; Massling et al., 2015).
Unlike much of central Greenland, the summit of Devon ice cap is subject to
partial melting at the surface during summer months, and meltwater can
percolate and refreeze into the underlying snow and firn to form
infiltration ice features (“melt layers”). The volumetric percentage of melt
layers in the DV99.1 core was measured by Fisher et al. (2012) as a proxy for
past summer warmth. These data show that surface melt rates at the coring
site increased abruptly in the mid-19th century following the end of the
Little Ice Age cold interval, and have since averaged 22 % (median 19 %), occasionally exceeding 50 % in the 20th century (Fig. 4). The
DV99.1 coring site is above the present-day upper limit of the superimposed
zone (∼1400 m a.s.l.; Gascon et al., 2013) and the firn there
is >60 m thick. Therefore, it is very unlikely that there is any net
loss by runoff at this location: any meltwater produced in the summer must
refreeze in the firn. However, even without net losses, one must consider
whether meltwater percolation and refreezing could account for the limited
variability in the DV99.1 rBC record during the 19th and 20th centuries.
The post-depositional mobility of BC particles in melting snow is not well
known, and likely depends on the hydrophobicity of these particles, which is
influenced by the presence or absence of surface coatings, for example,
with SO42- (Liu et al., 2011, 2013). Doherty et al. (2013)
investigated the vertical redistribution of BC and other light-absorbing
particles in snow and firn near Dye 2 (66∘ N; ∼2100 m a.s.l.; Fig. 5) in a part of the Greenland ice sheet's percolation zone
where melt layers >10 cm thick are now commonly found (de la
Peña et al., 2015; Machguth et al., 2016). Only very limited vertical
redistribution of BC was observed in the snow and firn, and surface melt and
percolation did not obliterate seasonal variations of BC in the firn
stratigraphy. Doherty et al. (2013) attributed this result to the low
scavenging efficiency of these particles by meltwater (∼20 %–30 %).
At the DV99.1 site on Devon Island, ice layers >10 cm are
comparatively very rare, but A˙(0.14 m a-1) is only half of
that in the Dye 2 area (∼0.32 m a-1; Buchardt et al.,
2012). Therefore, surface melt could mask some seasonal variations of rBC in
the firn.
The depth at which meltwater could percolate in firn at the DV99.1 site is
not precisely known over the time period covered in the rBC record. The
thickness of the firn zone there (>60 m) is much greater than
other locations, such as Lomonosovfonna summit, Svalbard (∼25 m; Kekonen
et al., 2005). If we accept the estimated depth range of 0.5–2 m for
meltwater-induced relocation of water-soluble ions at Lomonosovfonna summit
for 2000–2007 reported by Vega et al. (2016), then it is highly unlikely
that the relocation of rBC particles could be deeper at the DV99.1 site. The summit
of Devon ice cap is ∼650 m higher than the Lomonosovfonna
summit (1250 m a.s.l.), has a much lower mean annual surface temperature
(-22 ∘C, compared with ∼-10 to -12 ∘C at
Lomonosovfonna; Ward van Pelt, personal communication, 2017), and the 10 m firn temperature on
Devon ice cap summit was <-15 ∘C (Bezeau et
al., 2013), in 2012, while at Lomonosovfonna it was -2 to -3 ∘C
(van de Waal et al., 2002), in 1997. Attempts were also made to quantify
post-depositional deposition of ions and/or particles by melt/percolation on
Penny ice cap on Baffin Island (66∘ N; Grumet et al., 1998;
Zdanowicz et al., 1998), where estimated summer melt rates over the last 150 years
are much higher (40 %–100 %) than at the DV99.1 site (Zdanowicz et al.,
2012). On Penny ice cap during the mid-1990s, ions and particles were
estimated to be redistributed over depths of 3–5 m. Therefore, a plausible,
conservative estimate of the maximum melt-induced relocation depth at the
DV99.1 site for the time period of interest might be 3 m (firn
depth). With a mean accumulation rate of 0.16 m H2O a-1 at the
site, soluble impurities could be offset by meltwater percolation in the
core by 5–8 years relative to their true depositional depth/age, and
probably less for BC particles given their hydrophobicity. In this paper, we
focus on interdecadal variations in rBC concentrations. At such a
time-averaging window length, the effect of impurity relocation by melt
should largely even out.
However, there is another consideration. Unlike in the Doherty et al. (2013) study,
rBC concentrations in the DV99.1 core were measured by SP2,
and the detection efficiency of this method for BC in liquid samples depends
on the type of nebulizer used for inflow. As previously mentioned, Schwarz
et al. (2012) and Wendl et al. (2014) showed that the relative
aerosolization efficiency of rBC by the U5000AT ultrasonic nebulizer used in
the analysis of the DV99.1 core drops rapidly for particles with a
volume-equivalent diameter >500 nm (∼10 %
efficiency at a volume-equivalent diameter of 600 nm). Coagulation and
agglomeration is known to increase the size of BC particles during thaw and
refreezing of snow (Schwarz et al., 2013), which raises the possibility
that the SP2 may underestimate the true mass concentration of BC particles
in those parts of the DV99.1 that contain icy layers (Fig. 3).
To verify this, we examined the probability distribution of rBC particle
mass in sections of the DV99.1 ice core with different percentages of melt
layers. We compared core sections from depths between 37 and 38 m (corresponding
to the time interval ∼1803–1814) which only had 1 % melt
layers, with sections from depths between 13 and 16 m (time interval ∼1943–1963),
which had up to 53 % melt features (min. 9 %); we found
no significant differences between these core sections (Fig. S14). If rBC
particles had coagulated to form larger clusters in sections of core where
a lot of percolating meltwater refroze, the probability distribution or rBC mass
in these sections should be positively skewed relative to that in core
sections not impacted by meltwater, but our data show no evidence of this.
While it remains possible that melt–refreezing may have contributed to
masking some of the historical variations in atmospheric BC deposition at the DV99.1 site,
it seems unlikely, based on available evidence, that this factor alone could
account for the low rBC concentrations in the DV99.1 core, when compared to
Greenland records analyzed using the same methods.
Some of the central and northern Greenland sites (e.g., Summit, NEEM) from
which ice-core rBC records were developed by the SP2 method (Fig. 5)
experience less surface melt than Devon ice cap, and BC particles in firn at
these sites are probably largely unaffected by post-depositional
coagulation. Other coring sites located in southern Greenland (ACT2, D4) or
at lower elevations (Humboldt) may experience some surface melt and
refreezing in summer, but statistics on ice layer frequency at these sites
are unpublished, so this cannot be verified.
Regional-scale factors
Other reasons for the differences between the DV99.1 and Greenland rBC
records (Figs. 6 and 7) may be found in the atmospheric transport paths that
deliver BC to the Canadian High Arctic, relative to Greenland. Shindell et al. (2008)
used multiple atmospheric transport models to investigate the
sensitivity of near-surface airborne BC concentrations in the Arctic to
regional anthropogenic emissions. They found that Europe and North America
likely contribute equally to BC deposition over Greenland, whereas the
central and Russian sectors of the Arctic are more impacted by European
emissions. Atmospheric BC in the Canadian High Arctic may be affected by
both European and North American emissions, but the region is expected to be
less sensitive to changes in these emissions compared to other parts of the
Arctic, partly because it is very remote from all BC source regions
(Shindell et al., 2008; their Figs. 9 and 10).
Sharma et al. (2006) and Huang et al. (2010) used air back-trajectory
analyses to investigate the probable source regions of BC detected at Alert
in winter and spring, and identified Russia and Europe as dominant, followed
by North America. The summit of Devon ice cap is 1000 km further south and
∼1.9 km higher; thus, it could be affected by a different mix
of BC source contributions than Alert. To verify this, and also to contrast
the situations of Devon ice cap and Greenland, we computed ensemble 10-day
air back-trajectories from both Devon ice cap summit and from Summit,
Greenland, using the Hybrid Single Particle Lagrangian Integrated Trajectory
model (HYSPLIT v.4) of the NOAA Air Resources Laboratory (Draxler and Hess,
2014; Stein et al., 2015). As input, we used meteorological fields of the
NCEP-NCAR 50-year reanalysis product, which are available on a global 2.5×2.5∘ grid at 6-hourly temporal resolution (Kistler et
al., 2001). Back trajectories starting daily at 12:00 UTC were computed
over the period from 1948 to 1999. However, unlike Sharma et al. (2006) and Huang et
al. (2010), we did not use trajectory clustering, as results are highly
sensitive to the quality and density of meteorological data coverage used in
trajectory computations, and to the arrival height of trajectories (i.e.,
starting point of back trajectories; Kassomenos et al., 2010; Su et al.,
2015). Instead, we computed probability density maps or air parcel residence
time from all combined trajectories over an equal area grid with 200×200 km resolution, following a methodology analogous to that of
Miller et al. (2002).
Results (Fig. 8) show that for 10-day transport periods, air parcels
arriving at Summit, Greenland, are more commonly advected from the
south-southwest than from other directions, and frequently reach central
Greenland after transiting over the North Atlantic, consistent with earlier
findings by McConnell et al. (2007; their Fig. S1). In contrast, air that
reaches the summit of Devon ice cap more comes frequently from the
west-northwest, and transits over the Arctic Ocean, which agrees with
findings from analyses of low-level air transport to Devon ice cap by Colgan
and Sharp (2008) for the period from 1979 to 2003. Therefore, it is likely that a
large part of BC transported to Devon ice cap is from regional emission
sources located in northwestern North America and/or in the central or
eastern parts of Eurasia.
(a) Historical variations in rBC concentration in the
DV99.1 core, 1760–1992, compared with reconstructed historical trends in
(b) fire frequency in the eastern boreal forest region of Canada
(Girardin et al., 2006), and (c) burned area across northern Canada
(Girardin, 2007) and in the boreal and grassland regions of Russia and
central Asia (Mouillot and Field, 2005). All data were log-transformed to
facilitate visual comparisons.
Smoke plumes from forest or grassland fires, natural or provoked, can reach
the Arctic and contribute to BC pollution, particularly during summer (Stohl
et al., 2006; Paris et al., 2009; Warnecke et al., 2009; Quennehen et al.,
2012; Zennaro et al., 2014; Hall and Loboda, 2017). Back-trajectory analyses
of BB aerosols detected at Eureka on Ellesmere Island (80∘ N; Fig. 1)
indicate (as anticipated) that boreal forest/grassland regions of Russia
and Canada are the dominant source regions for these long-range plume
transport events, followed by north-central USA and Alaska (Viatte et al.,
2015). To investigate the impact of forest/grassland fire emissions on BC
deposition to Devon ice cap, we compared the DV99.1 rBC record with
reconstructed variations in fire frequency and/or burned area across Canada
and Russia during the 19th and/or 20th centuries (Fig. 9; data from
Girardin, 2007; Girardin and Sauchyn, 2008; Girardin et al., 2006; and
Mouillot and Field, 2005). On an interdecadal timescale, no statistically
meaningful correlations (p<0.05) could be identified between the
DV99.1 rBC record and the fire histories. If fire emissions contribute to BC
deposition on Devon ice cap, these contributions are either too small and/or
mixed in the DV99.1 record to be correlated with variations in fire
frequency or burned area in the source regions.
Aerosol species such as K+ or NH4+ are commonly associated
with BB emissions, and are often used as BB tracers in polar snow (Simoneit,
2002; Legrand et al., 2016). Cheng (2014) identified sectors of
south-central Russia and Kazakhstan as source regions for both BC and
K+ aerosols transported to Alert between 2000 and 2002. However, we did
not find any significant correlations (p<0.05) between
interdecadal variations of rBC in the DV99.1 core and either
(K+)BB or NH4+ in the DV98.3 record (Fig. S15). Whatever
contributions BB emissions make to (K+)BB or NH4+
deposition on Devon ice cap, these do not directly covary with BC
deposition, possibly due to the different post-depositional relocation of these
impurities in the DV98.3 and DV99.1 cores, and mixing from multiple
emission sources. For example, ammonia (NH3) emissions from seabird
colonies near Baffin Bay may be a larger regional source of NH4+
to Devon ice cap than distant wildfires (Wentworth et al., 2016).
Atmospheric BC deposition rates
In 90 % of the analyzed DV99.1 core, rBC concentrations are <3 ng g-1, and in the uppermost section of the core (depths 3–4 m), they
are mostly ≤1 ng g-1. These concentrations are very low compared
with the 8–14 ng g-1 reported by Doherty et al. (2010) for seasonal
snow sampled across the Canadian Arctic in 2009. Part of the apparent
discrepancy may be due to differences in analytical methods: the BC
concentrations in snow reported by Doherty et al. (2010) were measured using
a spectrophotometric technique which tends to yield larger mass
concentrations relative to the SP2 method (Schwarz et al., 2012). Also, as
previously stated, rBC levels measured in the DV99.1 core may underestimate
actual deposition due to wind scouring of winter snow. Atmospheric BC
deposition over the summit region of Devon ice cap could also be lower than
near sea level, where most of Doherty et al.'s (2010) samples were obtained,
because most of the ice cap's accumulation area (≥∼1150 m a.s.l.) is
above the typical altitude range of low-level Arctic
stratocumulus cloud decks which promote aerosol scavenging (Browse et al.,
2012).
Taking the aforementioned uncertainties into account, we estimated the
average late 20th century atmospheric flux of rBC (FrBC) over the summit
region of Devon ice cap using measurements of rBC concentrations in the
DV99.1 core for the period from 1963 to 1990, and data on spatial and temporal variations of
A˙ from Colgan and Sharp (2008) and from winter mass balance surveys
carried out over the ice cap since the early 1960s. The period from 1963 to 1990 was
selected because the 1963 radioactive layer in Devon ice cap firn provides a
reference level to constrain estimates of A˙ (Colgan and Sharp,
2008). Our calculations yield a mean FrBC of 0.2±0.1 mg m-2 a-1.
If μg, rather than μ, is used to estimate average
rBC concentrations, the estimated FrBC is slightly lower (0.1 mg m-2 a-1).
Furthermore, if the measured concentrations of rBC are assumed to
be underestimated by 60 %–80 % due to wind scouring of winter snow layers
and/or inadequate detection by the SP2 instrument, the adjusted figures for
FrBC are only slightly higher, ranging between 0.2 and 0.3 mg m-2 a-1.
These estimates are at the low end of the measured net rBC deposition rates in
Greenland ice cores between the early 1960s and late 1990s, which vary from
∼0.1 to ∼2.3 mg m-2 a-1 (Lee et al.,
2013; see also Fig. S16). Compared with most of central and southern
Greenland, the summit region of Devon ice cap experiences low snow
accumulation rates (0.17–0.25 m H2O a-1, or ≤0.31 m a-1
in ice equivalent; Colgan and Sharp, 2008), and this probably accounts, at
least in part, for the lower rBC accumulation rates there, given the
important role of precipitation scavenging in controlling atmospheric BC
deposition in the Arctic (Garrett et al., 2011; Browse et al., 2012). Other
reasons for the differences in rBC accumulation between Devon ice cap and
Greenland may be found in the predominant patterns of air transport trajectories
from source regions, as previously discussed.
Summary and conclusions
We developed a > 250-year time series of atmospheric rBC
deposition from Devon ice cap spanning the years ∼1735–1992.
The rBC ice-core record (core DV99.1) is the first from the Canadian Arctic,
and supplements existing ice-core records of rBC from Greenland developed
using the same analytical methods. The DV99.1 record differs from the Greenland
records in that it only shows a very modest and gradual rise in rBC
deposition through the 19th and early 20th century, unlike most Greenland
ice cores, in which there are large, well-defined rises from 1880 to the 1990s,
peaking in the 1910s. This rise was attributed to BC emissions from coal
combustion, which also emitted SO2 and trace metals such as Pb
(McConnell et al., 2007). Ice cores from Devon ice cap (DV98.3, DV2000) show
that the deposition of SO42- and Pb also increased there during
the 20th century, although the DV99.1 core shows no concomitant rise in rBC.
We suggest that differences between the DV99.1 and the Greenland rBC records are
due to a combination of methodological, site-specific, and regional-scale
factors. The DV99.1 coring site is subject to summer melt–freeze
cycles, and this may lead to some underestimation of true rBC concentrations
by the SP2 method. There is also evidence of wind scouring of snow at the
site, which may lessen the amplitude and resolution of historical variations
in BC deposition recorded in the core. Air back-trajectory analyses suggest
that, compared to Greenland, BC deposition on Devon ice cap is less
sensitive to BC emissions from the North Atlantic sector (eastern North
America and western Europe) than Greenland is. We hypothesize that BC
aerosols reaching Devon ice cap more frequently originate from
north-central/northwestern North America, and/or from Russia and central
Asia. The relatively long transport trajectories over the Arctic Ocean allow
for greater atmospheric mixing and deposition of aerosols to occur during
transit, which obscures source-receptor relationships. If correct, this
interpretation implies that historical trends in BC deposition over the
Arctic, and the resulting albedo-climate forcing, are likely subject to
large spatial variability, even over the relatively short distance between
Devon Island and Greenland. This variability, which is probably linked to
differences in BC aerosol transport patterns and atmospheric residence time
(Bauer et al., 2013), must be accounted for when attempting to model the
impact of past and future BC emission trends on the Arctic climate system.
This study also underscores the challenges of interpreting records of
aerosol deposition developed from firn or ice cores drilled on small ice
caps or glaciers, where local topographic and climatological effects can
impact the preservation of atmospheric signals, when compared with the
central regions of large ice sheets. A limitation of our study stems from
the fact that the DV99.1 record of rBC deposition is from a different site
than records of other aerosol species (SO42-, Pb) previously
obtained from Devon ice cap summit. To verify our interpretation of the
DV99.1 rBC record, a new core should be drilled from the ice cap summit, or
from another ice cap less affected by wind scouring and melt–freeze effects
(e.g., on northern Ellesmere Island), and on which co-registered
measurements of rBC and other aerosols could be made. This is particularly
important when one considers the large amount of spatial variability
inherent in ice-core records, even in areas of optimal preservation (e.g.,
Gfeller et al., 2014).