The tundra plays a pivotal role in the Arctic mercury
(Hg) cycle by storing atmospheric Hg deposition and shuttling it to the
Arctic Ocean. A recent study revealed that 70 % of the atmospheric Hg
deposition to the tundra occurs through gaseous elemental mercury (GEM or Hg(0))
uptake by vegetation and soils. Processes controlling land–atmosphere
exchange of Hg(0) in the Arctic tundra are central, but remain
understudied. Here, we combine Hg stable isotope analysis of Hg(0) in the
atmosphere, interstitial snow air, and soil pore air, with Hg(0) flux
measurements in a tundra ecosystem at Toolik Field Station in northern
Alaska (USA). In the dark winter months, planetary boundary layer (PBL)
conditions and Hg(0) concentrations were generally stable throughout the day
and small Hg(0) net deposition occurred. In spring, halogen-induced
atmospheric mercury depletion events (AMDEs) occurred, with the fast
re-emission of Hg(0) after AMDEs resulting in net emission fluxes of Hg(0).
During the short snow-free growing season in summer, vegetation uptake of
atmospheric Hg(0) enhanced atmospheric Hg(0) net deposition to the Arctic
tundra. At night, when PBL conditions were stable, ecosystem uptake of
atmospheric Hg(0) led to a depletion of atmospheric Hg(0). The night-time
decline of atmospheric Hg(0) was concomitant with a depletion of lighter
Hg(0) isotopes in the atmospheric Hg pool. The enrichment factor,
ε202Hgvegetationuptake=-4.2 ‰ (±1.0 ‰) was consistent
with the preferential uptake of light Hg(0) isotopes by vegetation. Hg(0)
flux measurements indicated a partial re-emission of Hg(0) during daytime,
when solar radiation was strongest. Hg(0) concentrations in soil pore air
were depleted relative to atmospheric Hg(0) concentrations, concomitant with
an enrichment of lighter Hg(0) isotopes in the soil pore air, ε202Hgsoilair-atmosphere=-1.00 ‰
(±0.25 ‰) and E199Hgsoilair-atmosphere=0.07 ‰ (±0.04 ‰). These
first Hg stable isotope measurements of Hg(0) in soil pore air are
consistent with the fractionation previously observed during Hg(0) oxidation
by natural humic acids, suggesting abiotic oxidation as a cause for observed
soil Hg(0) uptake. The combination of Hg stable isotope fingerprints with
Hg(0) flux measurements and PBL stability assessment confirmed a dominant
role of Hg(0) uptake by vegetation in the terrestrial–atmosphere exchange of
Hg(0) in the Arctic tundra.
Introduction
Mercury (Hg) is a high priority pollutant that causes neurodevelopmental
deficits in children and cardiovascular disease in adults
(Budtz-Jørgensen et al., 2000; Roman et al., 2011). Arctic populations
are particularly exposed to high Hg levels, despite few local anthropogenic
Hg emission sources, due to their traditional diet consisting of high
trophic level seafood (Sheehan et al., 2014). Anthropogenic
Hg emissions from midlatitudes reach remote ecosystems, such as the Arctic,
via the long-range transport of gaseous elemental mercury (GEM or Hg(0))
(Douglas et al., 2012). It has long been thought that springtime
oxidation of Hg(0) driven by photochemically produced bromine radicals
Br⚫ on surface snow, referred to as “atmospheric mercury depletion
events” (AMDEs), leads to enhanced deposition of divalent mercury (Hg(II))
in Arctic regions (Steffen et al., 2008). Such Hg(II) deposition to the
snowpack during AMDEs, however, can be subject to photoreduction and fast
re-emission back into the atmosphere, minimising the net load of Hg by
AMDEs to snow (Douglas et al., 2012; Johnson et al., 2008).
Tundra soils play a central role in the Arctic Hg cycle by storing
atmospheric Hg deposition from where it can then be mobilised and transported to
the Arctic Ocean (Obrist et al., 2017; Sonke and Heimburger, 2012; Sonke et
al., 2018). Over millennia, tundra vegetation and soils have drawn Hg(0) out
of the atmosphere, resulting in one of the largest pools of Hg (408–863 Gg Hg, top 1 m) stored at the Earth's surface (Obrist et al., 2017; Olson
et al., 2018; Schuster et al., 2018). Arctic rivers deliver 44–50 Mg a-1 of Hg to the Arctic Ocean (Dastoor and Durnford, 2014; Sonke et
al., 2018). In aquatic ecosystems, Hg can be transformed to methyl mercury that
bioaccumulates in the aquatic food chain, resulting in elevated Hg
concentrations in high trophic level fishes and mammals (Outridge et al.,
2009; Douglas et al., 2012).
On a global scale, vegetation uptake of atmospheric Hg(0) represents the
dominant pathway of atmospheric Hg deposition to terrestrial surfaces,
resulting in strong seasonal variation of atmospheric Hg(0) concentrations,
with minima in summer when vegetation activity is highest (Jiskra et al.,
2018). Direct Hg(0) flux measurements over selected surfaces (e.g. soil,
snow, or leaves), however, do not always provide conclusive answers about
the direction and magnitude of Hg exchange between terrestrial ecosystems
and the atmosphere. For example, a review of 132 terrestrial–atmosphere
Hg(0) flux studies conducted over various surfaces over the last 30 years
estimated a wide range of net fluxes in the range of -513 to 1650 Mg a-1 (37.5th to 62.5th percentile)
(Agnan et al., 2016). In recent years, a number
of studies using Hg(0) flux measurements based on micrometeorological methods
directly quantified net ecosystem exchange fluxes of Hg(0) over terrestrial
ecosystems (i.e. at the ecosystem level), including atmosphere–vegetation
exchange and underlying soil/litter contributions (Lindberg et al.,
1998; Fritsche et al., 2008; Bash and Miller, 2009; Castro and Moore,
2016; Osterwalder et al., 2017). Measurements of multi-level Hg(0) gradients
and interstitial snow air and soil pore air provided additional constraints on
the terrestrial surface exchange flux (Sigler and Lee, 2006; Moore and
Castro, 2012; Faïn et al., 2013; Obrist et al., 2014; Fu et al., 2016b; Agnan et al., 2018).
Hg stable isotopes are a powerful tool to study the deposition and
re-emission pathways of Hg to terrestrial ecosystems. The Hg stable isotope
fingerprint of soil samples reflects the source contribution of atmospheric
Hg(0) dry deposition, Hg(II) wet deposition, and Hg from geogenic origin, as
well as processes fractionating Hg isotopes during post-deposition
processes, e.g. re-emission (Demers et al., 2013; Jiskra et al.,
2015; Enrico et al., 2016). Mercury has seven stable isotopes, which can
undergo mass-dependent fractionation (MDF, described by δ202Hg), mass-independent fractionation of odd-mass-number isotopes
(odd-MIF, described by Δ199Hg and Δ201Hg), and mass-independent fractionation of even-mass-number isotopes (even-MIF, described by Δ200Hg and
Δ204Hg), thereby producing a multidimensional isotopic
fingerprint (Obrist et al., 2018). Atmospheric Hg(0) and Hg(II) in wet
deposition exhibit distinct Δ199Hg and Δ200Hg
signatures (Gratz et al., 2010; Chen et al., 2012; Sherman et al.,
2012b; Demers et al., 2013; Enrico et al., 2016). Foliar uptake of atmospheric
Hg0 discriminates heavier Hg isotopes, leading to consistently lower
δ202Hg values reported in foliage compared with atmospheric Hg(0)
(Demers et al., 2013; Enrico et al., 2016; Obrist et al., 2017; Yu et al.,
2018; Yuan et al., 2018). Using the triple isotopic fingerprint to
distinguish between Hg(0) and Hg(II) deposition, an increasing number of
studies around the globe have revealed that 60 %–90 % of Hg in soils is derived
from Hg(0) uptake by vegetation (Demers et al., 2013; Jiskra et al.,
2015; Enrico et al., 2016; Zheng et al., 2016; Obrist et al., 2017).
This paper is part of a larger study aiming to better understand the fate of
Hg in Arctic tundra ecosystems, centred around a 2-year field campaign on
the Arctic Coastal Plain of Alaska. In Obrist et al. (2017),
we performed a 2-year mass balance of terrestrial–atmosphere exchange over
the Arctic tundra. We also investigated the spatial distribution of Hg in
tundra soils (Olson et al., 2018), and spatial and temporal patterns of
Hg in snow (Agnan et al., 2018) and in vegetation
(Olson et al., 2019). In our previous work, we showed that the
uptake of atmospheric Hg(0) by vegetation and soil represents 70 % of
total atmospheric deposition and has led to high Hg levels in Arctic soils
(Obrist et al., 2017, 2018). In this study, we explore the
use of novel Hg stable isotope measurements of Hg(0) in interstitial snow
air and soil pore air to identify the processes driving tundra Hg(0)
deposition. We further discuss the effects of terrestrial–atmosphere
exchange processes and planetary boundary layer stability on the Hg(0)
concentration and Hg stable isotope signature measured in the atmosphere.
Materials and methodsStudy site
The study was conducted at Toolik Field Station (68∘38′ N, 149∘36′ W) on the Arctic Coastal Plain in northern
Alaska, USA, 180 km inland from the Arctic Ocean coast. All measurements
were conducted on an acidic tussock tundra, on Aquiturbels soils with an
active layer of 60–100 cm (Obrist et al., 2017). The
climate of Toolik Field Station is characterised by low mean annual
temperatures of -8.5∘C and a mean annual precipitation of 312 mm a-1 (Cherry et al., 2014). During the 1-year Hg
isotope campaign from October 2015 to September 2016, the tundra was
snow-covered for a total of 248 d (Agnan et al., 2018), leading
to a relatively short snow-free growing season.
Hg stable isotope sampling and measurement
Hg(0) was continuously sampled in the atmosphere (0.3 and 2 m above
ground), in interstitial air of surface snow (0 and 0.1 m above ground),
and in soil pore air (0.4 m below ground) at low flow rates of 0.2 L min-1. Interstitial snow air for Hg(0) stable isotope measurements was
sampled from a dedicated snow tower adapted from Seok et al. (2009),
consisting of an aluminium construction with three gas inlets on horizontal
bars at 0, 0.1, and 0.3 m above ground (Fig. S1a in the Supplement). Two soil wells as
described in Obrist et al. (2017) were dedicated to Hg(0)
stable isotope sampling of soil air. Each soil well had two gas inlets, positioned 1.5 m apart at 0.4 m depth, each inlet consisting of a 47 mm single stage filter assembly (Savillex, Eden Prairie, USA) with a Teflon® filter membrane (Fig. S1b).
Hg(0) was trapped on iodated activated carbon (IAC) traps (Brooks Rand, 0.1 g in custom made 12 cm long glass tubes with an inner diameter of 4 mm, Fu et al., 2014) and samples were changed manually during
site visits every 6 to 8 weeks. During site visits in March and June 2016,
higher temporal resolution (2–4 d) sampling was conducted with higher
flow rates of 1.5–2 L min-1. Overall, total volumes of sampled air per
sample ranged from 5.7 to 17.7 m3 (Table S1 in the Supplement). During the growing season
(June to September 2016), diel variation of atmospheric Hg(0) was assessed
using two parallel sampling lines operated with a time switch. Daytime
samples were collected from 06:00 to 22:00 AKST (Alaska standard time – used for all times referenced throughout the paper), and night-time samples were
collected from 22:00 to 06:00. Of the 14 soil pore air samples taken during
the 1-year Hg isotope campaign, only three samples contained sufficient Hg
(>2.5 ng) for isotopic analysis. The reason for this was that the Hg(0) in soil pore air was largely
depleted (below the detection limit of ≈0.1 ng m-3) (Obrist et
al., 2017; Agnan et al., 2018), in
particular during winter months, making isotopic characterisation of the soil
pore air Hg(0) pool impossible during this time period.
The protocol for Hg stable isotope measurements of Hg(0) was adapted from Fu
et al. (2014). We used lower amounts of IAC trap material
(0.1 g) to reduce possible matrix effects during cold-vapour generation.
Breakthrough was tested in the lab and under field conditions by connecting
a Tekran 2537 after the IAC trap, and Hg(0) measurements were always below
the detection limit (<0.1 ng m-3). IAC traps were combusted in
a two-stage oven system and Hg was recovered in a 4.2 N HNO3, 1.2 N HCl oxidising acid trap. Hg stable isotope ratios of trap solutions were
measured by cold-vapour separation multi-collector inductively coupled plasma
mass spectrometry (CV MC-ICP/MS) at the Observatory Midi-Pyrénées,
Toulouse, France (Jiskra et al., 2019; Sun et al., 2013). Six process blanks and one
field blank were measured during the sample processing and were
0.25 ng Hg/trap (max 0.85 ng Hg/trap) on average, representing 1 %–2 % of the typical Hg
amounts collected during sampling periods. The amounts of Hg collected on
IAC-traps sampling atmospheric Hg(0) were compared to Hg(0) concentration
measurements using a Tekran 2537 and revealed sample yields of 107±23 % (mean ± 1 SD, n=22).
Mass-dependent fractionation (MDF) of Hg stable isotopes is reported in
small delta notation (δ) in per mille (‰)
deviation from to the reference NIST 3133 Hg standard:
δxxxHg=(xxx/198Hgsample/xxx/198HgNIST3133-1)×103,
where “xxx” refers to measured isotope masses: 199, 200, 201, 202, and 204.
Mass-independent fractionation (MIF) is reported in capital delta notation
(Δ), which is defined as the difference between the measured δ199Hg, δ200Hg, δ201Hg, and δ204Hg values and those predicted for MDF relative to δ202Hg using the kinetic MDF law:
ΔxxxHg=δxxxHg-SFxxx×δ202Hg,
where SFxxx is the mass-dependent scaling factor of 0.252 for
199Hg, 0.502 for 200Hg, 0.752 for 201Hg, and 1.493 for
204Hg (Blum and Bergquist, 2007). Hg isotope enrichment
factors associated with two pools (εxxxHgpool1-pool2 and
ExxxHgpool1-pool2) were
calculated from the difference in the MDF and MIF signatures between two pools
(pool 1 and pool 2) as follows:
3εxxxHgpool1-pool2=δxxxHgpool1-δxxxHgpool24ExxxHgpool1-pool2=ΔxxxHgpool1-ΔxxxHgpool2.
The MDF enrichment factors of a reaction (εxxxHgreaction) were determined by fitting a linear regression
model (lm function of R) to the observational data following Mariotti et al. (1981):
δxxxHgresidual=δxxxHg0+εxxxHgreaction×lnf,
where δxxxHgresidual corresponds
to the Hg isotope signature of the residual Hg(0), δxxxHg0 corresponds to the initial Hg(0) isotope
signature, and f corresponds to the fraction of Hg(0) remaining in the gas phase. Note
that for high f>0.4 the systematic error of this simplified
approach is minimal. The MIF enrichment factor (ExxxHgreaction) was calculated as follows:
ΔxxxHgresidual=ExxxHgprocess×δxxxHgresidual.
The long-term precision was assessed through repeated analysis of the
ETH-Fluka Hg standard, which yielded values of -1.44±0.19 ‰, 0.08±0.1 ‰, 0.02±0.1 ‰, 0.02±0.09 ‰, and -0.03±0.2 ‰ (2σ, n=73) for
δ202Hg, Δ199Hg, Δ200Hg, Δ201Hg, and Δ204Hg, respectively, which is in agreement with
published values (Jiskra et al., 2015; Smith et al., 2015). The UM-Almaden standard, now available as NIST-8610, was measured less frequently and results were -0.58±0.15 ‰, -0.02±0.09 ‰,
0.00±0.1 ‰, -0.06±0.12 ‰, and -0.04±0.23 ‰ (2 SD, n=21) for δ202Hg, Δ199Hg, Δ200Hg, Δ201Hg, and Δ204Hg, which is in agreement with
previously reported values (Demers et al., 2013; Jiskra et al.,
2015; Enrico et al., 2016).
Hg(0) flux measurements
Micrometeorological flux measurements to quantify Hg(0) exchange at the
ecosystem level were conducted using the aerodynamic gradient flux method.
Surface–atmosphere flux was calculated by measurement of concentration
gradients in the atmosphere above the tundra in conjunction with atmospheric
turbulence parameters as follows:
FHg0=-k×u∗×zϕh×∂cHg(0)∂z,
where k denotes the von Kármán constant (0.4), u∗ is the friction
velocity, z is the measurement height, ϕh is the Monin–Obukhov scaling
coefficient (dimensionless) (Monson and Baldocchi, 2014), and
∂c(Hg(0))/∂z is the vertical Hg(0) gas concentration
gradient. Hg(0) concentrations at heights of 61 and 363 cm above the soil
surface were measured through 0.2 µm Teflon® inlet
filters connected to ∼35 m of perfluoroalkoxy polymer (PFA) lines. A
valve control system with three-way solenoid valves (NResearch, West
Caldwell, NJ, USA) allowed switching between the gradient inlets to take place every 10 min. A set of trace gas analysers with a total sampling flow of 1.5 L min-1 was connected to the gradient inlets by solenoid valves. The
trace analysers included an air mercury analyser (Model 2537A, Tekran Inc.
Toronto, Canada) and a cavity ring-down (CRD) greenhouse gas analyser to
measure CO2, H2O, and CH4 (Los Gatos Research, San Jose,
USA). Fluxes were only calculated during periods of appropriate turbulence
following Edwards et al. (2005) and as described in
Obrist et al. (2017). Gradient data recorded during very
stable (ζ>2) and very unstable (ζ>2) periods were removed from the dataset (4 %). The overall data coverage was 79 %.
Only flux data with wind directions from the tundra
(40–300∘ N) were considered (73 %). No gap filling
was performed. For quality control, sampling line blanks and line
intercomparisons, where the two gradient lines were put on the same height,
were performed approximately every 6 to 8 weeks (Obrist et
al., 2017). The planetary boundary layer (PBL) stability was assessed
via the stability index (ζ, dimensionless), defined as
ζ=zL.
In the above-mentioned equation, z represents the height of the Metek USA-1 sonic anemometer (Metek
GmbH, Elmshorn, Germany), i.e. 236 cm above ground, and L represents the
Monin–Obukhov length. The PBL was considered stable when ζ>0.1, unstable when ζ<-0.1, and neutral
when -0.1<ζ<0.1 (Peichl et al., 2013).
During the sampling period, auxiliary variables showed the following daily
average values: air temperature of -7.4∘C (from -40.6 to
20.4 ∘C), relative humidity of 74 % (from 37 % to 98 %), and a wind speed of 2.36 m s-1 (from 0 to 7.82 m s-1).
Backward trajectory modelling
Backward trajectories of air masses were modelled using the HYSPLIT
transport and dispersion model (Stein et al., 2015) from the
NOAA Air Resources Laboratory, which is accessible via the READY website
(Rolph et al., 2017).
Results and discussion
We divide the presentation and discussion of results into three seasons of
the year. In winter (20 October 2015 to 17 March 2016), the tundra site was
continuously snow-covered and the climatic conditions were characterised by
low temperatures (mean =-17.8∘C, hourly max = 1.2 ∘C) and low solar radiation (mean = 0.02 kW m-2, hourly
max = 0.41 kW m-2). In spring (17 March–5 April 2016),
temperatures were low (mean =-19.2∘C, hourly max =-2.4∘C) and the tundra was still snow-covered; however, solar
radiation increased (mean = 0.13 kW m-2, max = 0.54 kW m-2)
and occasional atmospheric mercury depletion events (AMDEs) were detected at
the study site (Obrist et al., 2017; Agnan et al., 2018). During summer
(3 May–9 September 2016), the air temperature was above freezing (mean = 6.7 ∘C, hourly max = 25.1 ∘C), solar radiation was
high (mean = 0.19 kW m-2, hourly max = 0.80 kW m-2), and the
study site was predominantly free of snow.
Winter
Over the winter period, atmospheric Hg(0) concentrations and CO2 mixing
ratios were relatively constant and there was little diel variation (Figs. 1a, S2). Low solar radiation led to relatively stable PBL conditions
throughout the day (Fig. 1b). Hg(0) flux measurements revealed a small
deposition (-0.34±5.8 ng m-2 h-1, mean ±1 SD,
median =-0.2 ng m-2 h-1) (Obrist et al.,
2017). The Hg(0) net deposition flux is supported by observed depletions of
atmospheric Hg(0) in interstitial snow air (0.69±0.22 ng m-3,
mean ±1 SD, 0 and 0.1 m sampling height) relative to atmospheric
levels (1.07±0.04 ng m-3 – mean ±1 SD) (p=0.02,
two-sided t test), implying a net sink of atmospheric Hg(0) in the ecosystem
(Fig. 2c). A wintertime Hg(0) sink can either occur due to (i) Hg(0)
deposition to Arctic snow, (ii) uptake by soil or litter, or (iii) assimilation by vegetation still active under the snowpack. Depletion of
atmospheric Hg(0) in interstitial snow air was associated with an increase
in δ202Hg (1.08 ‰ ± 0.20 ‰ versus 0.77 ‰ ± 0.16 ‰ in ambient Hg(0), mean ± 1 SD, p=0.02,
two-sided t test) (Fig. 2a) and a decrease in Δ199Hg (-0.31 ‰ ± 0.05 ‰ versus -0.23 ‰ ± 0.06 ‰ in ambient Hg(0),
mean ± 1 SD, p=0.04, two-sided t test) (Fig. 2b). Both Hg(0) dry
deposition to surface snow (Douglas and Blum, 2019) and by
vegetation uptake (Demers et al., 2013; Enrico et al., 2016; Obrist et al.,
2017) have been reported to discriminate heavier Hg(0) isotopes, consistent
with these observations in the interstitial snow air. From a mass balance
perspective, however, Hg(0) dry deposition to snow is considered to only
play a minor role in the interior Arctic tundra. For example, using snow
data in Agnan et al. (2018), we calculated a total seasonal
snow Hg pool of only 50 ng m-2 at Toolik Field Station. Assuming that
all this Hg in the snow was originating from the dry deposition of Hg(0),
this would account for <10 % of the Hg(0) deposition during the
snow-covered period (total of 2.4 µg m-1 a-1; Obrist et al., 2017). In contrast, Douglas and Blum (2019) recently suggested that Hg(0) dry deposition to snow was
the major source of Hg in meltwater collected on the coast of the Arctic
Ocean close to Utqiagvik (former Barrow), ∼400 km north-west
of Toolik Field Station. In snow on the coast of the Arctic Ocean,
concentrations of halogens, which are considered to mediate reactive Hg(0)
uptake, are elevated compared with inland sites (Douglas and Sturm,
2004; Agnan et al., 2018; Douglas et al., 2017), leading to much higher snow Hg
pools (>2000 ng m-2) in coastal snowpacks (Douglas et al., 2017).
A major wintertime Hg(0) deposition
pathway to soils would be inconsistent with the observed enrichment in
heavier Hg(0) isotopes in the interstitial snow air (Fig. 2a), as Hg(0)
oxidation by humic acids in soils would lead to more negative δ202Hg values (Zheng et al., 2018) (see discussion below
in Sect. 3.3.3). Hg(0) re-emission from snow was insignificant in winter
due to the absence of sunlight, as indicated by lower Hg(0) concentrations
in interstitial snow air than in the atmosphere. By excluding snow and soil
uptake as dominant deposition pathway, we infer that the uptake of Hg(0) by
ground vegetation and/or litter on the soil surface may be driving Hg(0)
deposition during the winter months. MDF and MIF signatures observed in
interstitial snow air Hg(0) would be consistent with Hg(0) uptake by
lichen. Hg isotope signatures of lichen measured at the same site (δ202Hg =-0.80 ‰ ± 0.20 ‰, Δ199Hg = 0.20 ‰ ± 0.21 ‰, mean ± 1 SD, n=12) were in the opposite direction of interstitial snow air Hg(0) (Olson et
al., 2019). Lichen have been reported to actively exchange CO2 for
photosynthesis under snow cover (Kappen, 1993), and could possibly
also take up atmospheric Hg(0) during winter months.
Diel variation during winter (a, b, c), spring (d, e, f), and summer (g, h, i): (a, d, g)
average Hg(0) concentration (in red) and the CO2 mixing ratio (in blue) in the atmosphere (2 m height). (b, e, h) The median planetary boundary layer
stability parameter (ζ, grey), where positive values represent stable
conditions, and solar radiation is shown in yellow. (c, f, i) The average Hg(0) flux (in red)
and the average CO2 flux (in blue). The shaded areas in (a) and (c) represent the mean ± 1 SD interval of concentration and flux measurements, respectively.
Hg(0) measurements in interstitial snow air profiles. (a) Mass-dependent fractionation stable isotope signature of Hg(0) (δ202Hg), (b) mass-independent fractionation stable isotope signature of
Hg(0) (Δ199Hg), and (c) mean Hg(0) concentration. Note that the
Hg(0) concentration marked with an asterisk (*) was calculated from Hg recovered on the
IAC traps, whereas other Hg(0) concentration profiles were measured by an
automated trace gas system deployed in the snowpack (Agnan et al., 2018).
The dashed horizontal lines represent the average snow height during the
respective period.
Spring
During spring 2016, two major AMDEs occurred (events 1 and 3; Fig. 3).
During both events air masses were transported from the Arctic coast or the
Arctic Ocean to Toolik Field Station, as inferred from HYSPLIT backward
trajectory modelling. During the 19 March–20 March 2016 period (event 1 in Fig. 3), atmospheric Hg(0) concentrations dropped below the detection limit
(<0.1 ng m-3), while atmospheric Hg(II) concentrations
remained low (<0.1 ng m-3), with the exception of individual
spikes of up to 0.4 ng m-3. During the 1–3 April 2016 period (event 3 in Fig. 3), atmospheric Hg(0) concentrations decreased below the
detection limit at times, while Hg(II) concentrations remained high (around 0.4 ng m-3) for 2 days. During the similar AMDEs 1 and 3, O3
mixing ratios dropped below 10 ppb. Van Dam et al. (2013)
previously reported that during AMDEs ozone depletion events (ODEs)
also occurred at Toolik Field Station, concurrent AMDEs and ODEs were also
observed on the coast at Utqiagvik ∼400 km to the north-west.
Van Dam et al. (2013) suggested that AMDEs and ODEs were driven
by bromine emissions from the Arctic Ocean and transported to different
extents to the interior tundra. During event 1, no elevated Hg(II)
concentrations were observed; thus, we assume that air depleted in Hg(0)
was transported to Toolik Field Station, while the deposition
of Hg(II) likely occurred closer to the coast. During the 26–29 March 2016 period (event 2 in Fig. 3), when air masses were coming from
the south,
atmospheric Hg(0) concentrations temporally decreased to 0.75 ng m-3,
O3 remained high (>30 ppb), and Hg(II) concentrations
remained enhanced around 0.2 ng m-3 for 2 days while turbulent PBL
conditions prevailed. The source of these enhanced Hg(II) levels transported
from the Brooks Range mountains remains unknown.
Time series during the springtime AMDE period (18 March–4 April 2016) for
(a) the atmospheric Hg(0) concentration (red), the atmospheric Hg(II) concentration (green), and the
O3 mixing ratio (blue); (b) planetary boundary layer stability (ζ), where the shaded areas in green represent stable conditions (ζ>0.1) and shaded areas in red represent turbulent conditions
(ζ<-0.1); (c) solar radiation and air temperature; and (d) Hg(0) flux, where Hg(0) deposition is shown using green and Hg(0) re-emission is shown using
red. Midnight is indicated by dashed lines. Strong AMDEs when Hg(0)
concentrations dropped (1, 3) or Hg(II) concentrations increased (2) are
marked using numbers.
During the AMDE period (17 March–5 April 2016) an overall net Hg(0)
re-emission (1.6±11 ng m-2 h-1, mean ± 1 SD, median = 0.28 ng m-2 h-1) was measured which peaked after
the Hg(II) deposition, making the AMDE period the only time of the year when
net Hg(0) re-emission occurred (Obrist et al., 2017).
Strong Hg(0) re-emission from the snowpack has been reported during and
after AMDEs due to the fast reduction of Hg(II) deposition (Johnson et al.,
2008; Douglas et al., 2012).
Snowmelt occurred in May in 2016, during which the snow height declined rapidly between
7 May (24 cm) and 13 May (0 cm) and a small Hg(0) deposition possibly
occurred (-0.24±3.9 ng m-2 h-1, mean ± 1 SD, median:
0.08 ng m-2 h-1). In contrast to wintertime patterns, the
interstitial snow air Hg(0) during snowmelt showed low Δ199Hg
values of -0.62 ‰ and -0.44 ‰ (n=2)
versus -0.23 ‰ ± 0.06 ‰ in
ambient Hg(0) (Fig. S3b). The negative Δ199Hg values in Hg(0)
suggest a substantial contribution of Hg(0) re-emission after photoreduction
of Hg(II) in snow, which exhibited negative Δ199Hg values with
a minimum of -1.37 ‰ (Obrist et al.,
2017). This observation is consistent with previous chamber experiments,
where a negative Δ199Hg value of -2.08 ‰ was
reported for Hg(0) re-emission from snow (Sherman et al., 2010).
Even-MIF (Δ200Hg) is considered to be unaffected by
post-deposition processes such as re-emission (Sherman et al.,
2010; Enrico et al., 2016), providing a conservative tracer for the pathway
of atmospheric Hg deposition. Δ200Hg values measured in snow
impacted by AMDEs at Toolik Field Station and other sites in Alaska
(Obrist et al., 2017; Sherman et al., 2010, 2012a) are
similar to the Δ200Hg values of atmospheric Hg(0) (Fig. 4)
(-0.06 ‰ ± 0.06 ‰ versus
-0.05 ‰ ± 0.04 ‰, mean ± 1 SD). This similarity can be explained by a quantitative oxidation
of atmospheric Hg(0) to Hg(II) (e.g. event 1 in Fig. 3) that is deposited
to snow. Thus, Hg(II) in snow inherits the isotopic composition of the
source Hg(0) due to conservation of mass, irrespective of the isotopic
fractionation factor associated with Hg(0) oxidation. Several samples
exhibited Δ200Hg values between that of atmospheric Hg(0) and
Hg(II) in precipitation measured in temperate regions (Fig. 4). This
intermediate Δ200Hg signature can be explained by either AMDEs
with non-quantitative oxidation due to limited Br oxidant availability
(e.g. event 2 in Fig. 3), or a mixing of AMDE-derived Hg(II) with Hg(II)
present in the overlying Arctic free troposphere. Even-MIF (Δ200Hg) has been suggested as a promising tracer to distinguish between
atmospheric deposition of Hg(II) in precipitation, which exhibits positive
Δ200Hg anomalies, and direct Hg(0) deposition (e.g. uptake by
vegetation), which exhibits small negative Δ200Hg (Enrico et
al., 2016; Sun et al., 2019). We caution that the presence of AMDEs
complicates the use of Δ200Hg for mixing model-based Hg
deposition calculations in the Arctic (Obrist et al.,
2017).
Mass-dependent fractionation (δ202Hg) versus even
mass-independent fractionation (Δ200Hg) of atmospheric Hg(0)
(red circles), Hg(II) in wet deposition sampled at locations or during
seasons when no AMDEs occurred (blue triangles), and Hg(II) in Arctic show
sampled during the springtime AMDE season (green diamonds) for the Arctic tundra
at Toolik Field Station (filled symbols), and reported from elsewhere (empty
symbols). Atmospheric Hg(0) data are from Gratz et al. (2010), Demers et
al. (2013, 2015), Fu et al. (2016a), Enrico et al. (2016), and Obrist et
al. (2017); Hg(II) in wet deposition are from Demers et al. (2013), Sherman et al. (2015), Chen et al. (2012), and Gratz et al. (2010); and Hg(II)
in snow from AMDEs are from Obrist et al. (2017) and Sherman et al. (2010, 2012a). Error bars represent the analytical precision
determined by the 2 SD from multiple measurements of an in-house standard.
SummerDrivers of diel cycling in atmospheric Hg(0)
Figure 5 represents a time series of the atmospheric Hg(0) concentration, the
CO2 mixing ratio, the PBL stability, and Hg(0) fluxes during midsummer (15–28 July 2016). Atmospheric Hg(0) concentrations generally
declined during each night, and the strongest Hg(0) depletions (Hg(0)
<1 ng m-3) were observed when the PBL was stable (ζ>0.1, events 2 and 3 and green bars in Fig. 5b). These
Hg(0) night-time minima coincided with maxima of atmospheric CO2 mixing
ratios of 410 to 420 ppm. These patterns are consistent with measured Hg(0)
deposition fluxes during nights (daily minima at 00:00, -2.56±0.22 ng m-2 h-1 (mean ± 1 SD), Fig. 1i), under corresponding
CO2 accumulation in the PBL driven by night-time CO2 soil
respiration (Wofsy et al., 1993; Schlesinger and Andrews, 2000; Grant and
Omonode, 2018). On nights with unstable PBL conditions (e.g. event 1 in
Fig. 5), diel Hg(0) and CO2 variations were lower or absent due
to increased mixing with free tropospheric air containing background levels
of Hg(0) and CO2. During late summer (Fig. S4, 20–31 August 2016), the longer duration of stable nocturnal PBL conditions led to even
more pronounced night-time depletions in Hg(0). During daytime under strong
solar radiation, flux measurements showed a net Hg(0) emission around noon
(daily maxima at 11:00, 2.9±0.92 ng m-2 h-1, mean ± 1 SD, Fig. 1i). Daytime Hg(0) re-emission, however, did not lead to a
build-up of atmospheric Hg(0) above the surface due to prevailing turbulent
conditions, allowing for the efficient mixing with background free tropospheric air.
These patterns demonstrate how atmospheric Hg(0) and CO2 are both
controlled by the magnitude and direction of the net ecosystem exchange
fluxes in conjunction with the PBL stability. Overall, flux measurements showed
the tundra ecosystem to be a net sink of atmospheric Hg(0) over the duration
of the growing season (-0.76±6.3 ng m-2 h-1 – mean
±1 SD, median: -1.2 ng m-2 h-1).
Time series during a midsummer period (15–28 July 2016) of the (a) atmospheric Hg(0) concentration and CO2 mixing ratio; (b) median
planetary boundary layer stability (ζ), where the shaded areas in
green represents stable conditions (ζ>0.1) and shaded
areas in red represent turbulent conditions (ζ<-0.1); (c) solar radiation and air temperature; and (d) Hg(0) flux, with Hg(0)
deposition in green and Hg(0) re-emission in red. Midnight is indicated by
dashed lines.
Hg isotope fractionation during foliar uptake of atmospheric Hg(0)
Figure 6 shows a scatterplot of atmospheric Hg(0) concentrations and the
δ202Hg values sampled during different times of the day in
summer 2016. Hg(0) sampled during the night was characterised by lower Hg(0)
concentrations (1.06±0.13 ng m-3), higher δ202Hg
(1.31 ‰ ± 0.15 ‰), and similar
Δ199Hg values (-0.28 ‰ ± 0.08 ‰) (22:00–06:00, mean ± 1 SD, n=4),
compared with Hg(0) sampled during the day (Hg(0) =1.16±0.11 ng m-3, p=0.001; δ202Hg = 0.95 ‰ ± 0.22 ‰, p=0.01; Δ199Hg =-0.26 ‰ ± 0.05 ‰, p=0.70, mean ± 1 SD, 06:00–22:00, n=2; 24 h, n=4; p values based
on two-sided t tests). An enrichment of heavy atmospheric Hg(0) isotopes
during the night is consistent with the preferential uptake of light Hg(0)
isotopes by vegetation (Demers et al., 2013; Enrico et al., 2016; Obrist et
al., 2017; Olson et al., 2019; Yuan et al., 2018; Yu et al., 2016), which is the
dominant Hg(0) deposition pathway at the study site (Obrist
et al., 2017). Attributing the diurnal concentration and δ202Hg
variation of Hg(0) to vegetation uptake, we calculated an enrichment factor
of vegetation uptake, ε202Hgvegetationuptake, of
-4.22 ‰ ± 1.01 ‰ (mean ± 1 SE, R2=0.68, p=0.003) (Fig. S5) using a Rayleigh
model. Enrico et al. (2016) estimated an
ε202Hgplant-air of -2.6 ‰ for
foliar uptake by sphagnum moss, using a Rayleigh model to fit the
atmospheric Hg(0) concentration and δ202Hg measured at two
locations, a mountain site unaffected by local terrestrial–atmosphere
exchange and a peat bog where Hg(0) in air was depleted by foliar uptake.
Similarly, observations of the difference in δ202Hg between
plants and atmospheric Hg(0) suggested a ε202Hgplant-air value between -1 and -3 ‰
(Demers et al., 2013; Enrico et al., 2016; Obrist et al., 2017; Olson et
al., 2019; Yuan et al., 2018; Yu et al., 2016).
Mass-dependent Hg isotope signature (δ202Hg) of
atmospheric Hg(0) versus Hg(0) concentration during the snow-free growing
period (11 June–10 September 2016). The dashed line represents a non-linear
Rayleigh fit and the shaded area is the 95 % confidence interval (see main
text). Error bars represent the analytical precision determined by the 2 SD
from multiple measurements of an in-house standard.
At Toolik Field Station, the difference between δ202Hg in
vegetation relative to atmospheric Hg(0) was also considerably lower (range
of -1.29 ‰ to -2.09 ‰, depending on vegetation
species; Olson et al., 2019) than the fractionation factor
derived from the atmospheric pattern. This discrepancy can be explained by
the fact that δ202Hg signatures measured in vegetation not only
reflect the isotopic fractionation during foliar uptake, but also contain a
re-emission component. Hg(II) reduction is expected to lead to more positive
δ202Hg values in the residual, foliar Hg(II) fraction,
irrespective of the reduction mechanism (Bergquist and Blum, 2007; Zheng
and Hintelmann, 2010; Kritee et al., 2007; Jiskra et al., 2015). Re-emission
of foliar Hg is supported by the observed negative shifts in odd-mass isotope
MIF (Δ199Hg) in vegetation relative to Δ199Hg of
atmospheric Hg(0), which have been observed at Toolik Field Station
(Olson et al., 2019) and elsewhere (Enrico et al.,
2016; Demers et al., 2013). Positive Δ199Hg in atmospheric Hg(0)
re-emitted from foliage has recently been constrained by flux bag
experiments (Yuan et al., 2018). Therefore, it is expected that
the fractionation factor of foliar uptake is larger than just the difference
between δ202Hg in foliage and Hg(0) in the atmosphere. Instead,
the difference reflects a net fractionation consisting of the isotopic
fractionation during foliar uptake, as well as during foliar reduction and
re-emission. Our observation that the fractionation factor derived from
δ202Hg of atmospheric Hg(0) (Fig. 6) is larger than the
difference of δ202Hg between vegetation and the atmosphere
could also be associated with the diel variation of Hg(0) fluxes and PBL
dynamics. During daytime, atmospheric turbulence is higher and, therefore,
local signals of terrestrial re-emission are expected to be diluted by
mixing with background Hg(0). At night, when the PBL is stable, foliar
uptake of lighter Hg(0) isotopes is imprinted on the residual atmospheric
Hg(0).
Sink of Hg(0) in soil inferred from Hg stable isotopes
At our study site, soil pore air Hg(0) concentrations were below ambient
levels measured in the atmosphere all year (Obrist et al.,
2017), with an average concentration of 0.54±0.14 ng m-3,
indicating a consistent sink of Hg(0) in soils. Hg(0) in soil pore air
showed a lower δ202Hg (-0.01 ‰ ± 0.39 ‰) and a higher Δ199Hg (-0.18 ‰ ± 0.07 ‰) (mean ± 1 SD, n=3) compared with ambient atmospheric Hg(0) (Hg(0) =1.1±0.09 ng m-3, p=0.001, δ202Hg = 0.81 ‰ ± 0.18 ‰, p=0.057,
Δ199Hg =-0.25 ‰ ± 0.04 ‰, p=0.27; 24 h, mean ± 1 SD, p value of
two-sided t test, n=6) during the summer and fall periods, when we were
able to quantify soil pore air isotope patterns. Fitting the Hg stable
isotope fractionation trajectory for MDF and MIF of three data points of
soil pore air samples and the atmospheric Hg(0) samples resulted in
enrichment factors of ε202Hgsoilair-atmosphere=-1.00 ‰ ± 0.25 ‰ (mean ± 1 SE, R2=0.69, p=0.005; Fig. 7a) and
E199Hgsoilair-atmosphere=0.07 ‰ ± 0.04 (mean ± 1 SE, R2=0.32, p=0.11; Fig. 7b). Recently,
(Zheng et al., 2018) investigated Hg stable isotope
fractionation during the oxidation of dissolved Hg(0) by low molecular weight
thiol compounds and natural humic acids (HAs). For oxidation by HAs, they
reported an enrichment of light Hg(0) isotopes (ε202HgHg(0)-Hg(II)=-1.54 ‰ ± 0.05 ‰, mean ± 1 SE) and a positive odd-mass Hg
MIF (E199HgHg(0)-Hg(II)=-0.18 ‰ ± 0.03 ‰, mean ± 1 SE) in the residual Hg(0)
fraction (Zheng et al., 2018). Our limited number of soil air
measurements (n=3) agree with the fractionation trajectory for HA
oxidation (red straight lines in Fig. 7), suggesting abiotic oxidation in
soils as a cause of the observed soil Hg(0) uptake. In spite of a consistent
soil Hg(0) sink in soils, Obrist et al. (2014)
estimated that the soil Hg(0) sink results in small Hg(0) fluxes (<-0.03 ng m-2 h-1) due to low diffusivity. Comparing the soil
Hg(0) uptake flux reported by Obrist et al. (2014) to the net ecosystem flux measured at
Toolik Field Station, we estimate that such a soil Hg(0) sink would only
account for <5 % of the total Hg(0) deposition. Soil air δ202Hg is enriched in light isotopes relative to wintertime atmospheric Hg(0), whereas interstitial snow air Hg(0) is enriched in heavier isotopes, in the opposite direction of soil air (Fig. 8), suggesting that soil uptake of Hg(0) has a minor effect on interstitial Hg(0) patterns in the snowpack above.
Hg stable isotope systematics of Hg(0) in the atmosphere during
summer/fall and in soil pore air measured during the same time period. (a) Mass-dependent Hg isotope signature (δ202Hg) against Hg(0)
concentration. (b) Mass-independent Hg isotope signature (Δ199Hg)
against mass-dependent Hg isotope signature (δ202Hg). The
dashed green lines represent the regression of a Rayleigh model (a) and a
linear model (b). The straight red lines represent the trajectories for
abiotic dark oxidation of Hg(0) by natural humic acids (HAs) from Zheng et
al. (2018). For comparison with observations, the intercept of
the linear regressions was adjusted to fit through the average of
atmospheric Hg(0). The shaded areas represent the 95 % confidence
interval. Error bars represent the analytical precision determined by the 2 SD from multiple measurements of an in-house standard.
Overview of the mass-independent Hg isotope signature (Δ199Hg) against the mass-dependent Hg isotope signature (δ202Hg) of atmospheric Hg(0) (red) and Hg(0) in the pore air of snow (blue)
and soil (green) during winter (circles), the AMDE season (squares), and
summer/fall (triangles). Bulk vegetation measurements (green diamonds) are
reproduced from Obrist et al. (2017). The straight arrow
represents the fractionation during Hg(0) uptake by vegetation, and the dashed
arrow represents the expected development of the corresponding residual
Hg(0) pool. Error bars represent the analytical precision determined by the
2 SD from multiple measurements of an in-house standard.
Conclusions
We document that the concentrations and the isotopic composition of
atmospheric Hg(0) are strongly affected by terrestrial–atmosphere exchange,
in particular by vegetation uptake, of Hg(0). While directions and
magnitudes of the terrestrial–atmosphere exchange vary with season and time
of the day, atmospheric stability and the dynamics of the PBL strongly
affect atmospheric Hg(0) concentrations. During the vegetation period,
atmospheric Hg(0) isotope systematics were dominated by vegetation uptake of
Hg(0) discriminating heavy Hg(0) isotopes in the residual atmospheric pool,
which manifested itself most strongly during stable PBL conditions at night.
We found a larger discrimination of heavier Hg(0) isotopes during foliar
uptake when deriving a fractionation factor from atmospheric Hg(0)
observations, compared with deriving this factor based on the difference
measured between bulk Hg in vegetation and atmospheric Hg(0). While this
discrepancy is not fully understood, it may be attributed to photoreduction
and re-emission of lighter Hg(0) isotopes during the day. Hg(0)
concentrations were continuously depleted in interstitial snow and soil pore
air. Hg stable isotope measurements of Hg(0) in soil air indicated that the soil Hg(0) sink was driven by Hg(0) oxidation
by natural organic matter. However, we observed no isotopic traces of this
Hg(0) soil sink in interstitial snow air and the atmosphere above and
concluded that soil uptake of Hg(0) made only a minor contribution to the
net ecosystem exchange of Hg(0). Based on the stable isotope signatures of
Hg(0) in soil and snow pore air and mass balance considerations, we infer
that the uptake of Hg(0) by ground vegetation and/or litter on the soil
surface was likely responsible for Hg(0) deposition during the winter
months. The combination of stable isotope fingerprints with Hg(0) flux
measurements and PBL stability assessment confirmed a dominant role of Hg(0)
uptake by vegetation in Arctic tundra Hg cycling.
Atmospheric temperatures have increased twice as fast in the Arctic as in
temperate regions over the last 2 decades (Cohen et al.,
2014). The increase in Arctic temperatures has been accompanied by an earlier
snowmelt in spring and, thus, a longer vegetation period and higher maximal
greenness of Arctic vegetation (Box et al., 2019).
Vegetation uptake is estimated to currently drive the deposition of 210 Mg Hg a-1 of atmospheric Hg(0) to the Arctic tundra (Obrist et al.,
2017; Sonke et al., 2018). With an ongoing warming and greening trend
associated with climate change, an amplification of the Arctic tundra Hg(0)
vegetation pump can be expected. Contrarily, higher soil temperatures,
leading to permafrost thaw and degradation and development of thermokarsts, increase the risk of the re-mobilisation of large amounts of Hg that is currently stored in
Arctic tundra soils (St. Pierre et al., 2018; Olson et al., 2018; Schuster
et al., 2018). The net effect of these complex alterations in Arctic Hg
cycling driven by climate change is currently poorly understood and the risk
needs to be assessed.
Data availability
Hg stable isotope data are available in the Supplement. Metadata and atmospheric concentration data are available from the NSF Arctic Data Center (urn:uuid:e58b2369-0701-4dac-b3c6-b8e15a0cd187, https://arcticdata.io/catalog/view/urn:uuid:e58b2369-0701-4dac-b3c6-b8e15a0cd187, last access: 16 October 2019; Helmig et al., 2018). Flux data are available from Daniel Obrist (daniel_obrist@uml.edu) upon request.
The supplement related to this article is available online at: https://doi.org/10.5194/bg-16-4051-2019-supplement.
Author contributions
MJ conceived the study and led the Hg stable isotope sampling and analysis with support from JES and DO. YA led the data analysis of flux data. MJ wrote the paper with input from all authors.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
We thank Toolik Field Station and Polar Field Services staff for their support in setting up the field site and maintaining its operation for 2 years, with special thanks to Jeb Timm. We thank Christine L. Olson, Dominique P. Colgrove, Jacques Hueber, and Christopher W. Moore for assistance in the field and Jerome Chmeleff for support with Hg stable isotope measurements. The authors gratefully acknowledge the NOAA Air Resources Laboratory (ARL) for the provision of the HYSPLIT transport and dispersion model and READY website (http://www.ready.noaa.gov, last access: 5 August 2019) used in this publication.
Financial support
This research has been supported by the Swiss National Science Foundation (grant no. PZ00P2_174101), the H2020 European Research Council (MEROXRE; grant no. 657195), the US National Science Foundation (grant nos. 1304305, 1739567, and 1848212), and the European Research Council (MERCURY ISOTOPES; grant no. 258537).
Review statement
This paper was edited by Michael Weintraub and reviewed by five anonymous referees.
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