Hydrology and pore water chemistry in a permafrost wetland,
Ilulissat, Greenland
Søren Jessen1, Hanne D. Holmslykke2,*, Kristine Rasmussen3, Niels Richardt4, Peter E. Holm5
1 Department of Geosciences and Natural Resource Management, University of Copenhagen, Copenhagen,
Denmark
2 Department of Reservoir Geology, Geological Survey of Denmark and Greenland, Copenhagen, Denmark
3 Danish Nature Agency, Department of Groundwater, Aalborg, Denmark
4 Rambøll Denmark A/S, Environment, Copenhagen, Denmark
5 Department of Plant and Environmental Sciences, University of Copenhagen, Frederiksberg, Denmark
*Corresponding author: E-mail: [email protected], phone: +45 3814 2000.
This article has been accepted for publication and undergone full peer review but has not beenthrough the copyediting, typesetting, pagination and proofreading process which may lead todifferences between this version and the Version of Record. Please cite this article as an‘Accepted Article’, doi: 10.1002/2013WR014376
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Abstract
Hydrological and geochemical processes controlling the pore water chemistry in a permafrost wetland, with
loam overlain by sphagnum peat, were investigated. The vertical distributions of dissolved Cl, and of pore
water δ18O, appeared unrelated to ion freeze-out and isotope ice-water fractionation processes, respectively,
dismissing solute freeze-out as a main control on the water chemistry. However, concentrations of major
ions, others than Cl, generally increased with depth into the active layer. A conceptual model for water and
solute movement in the active layer was derived. The model indicates upwards diffusive transport of
elements, released in the loam layer by mineral weathering, to the peat layer, in which lateral advective
transport dominates. Active layer pore water and water of melted core sections of permafrost were of Ca-
Mg-HCO3 type (1:1:4 stoichiometry) and were subsaturated for calcite and dolomite. The results are
consistent with an annual cycling of inorganic carbon species, Ca and Mg, via cryogenic carbonate
precipitation during fall freeze-up and their re-dissolution following spring thaw. Similarly, elevated Fe2+
concentrations appear to be related to cryogenic siderite formation. Pore water in the active layer showed
high partial pressures CO2, indicating the feasibility of bubble ebullition as a greenhouse gas emission
pathway from permafrost wetlands. Elevated concentrations of geogenic trace elements (Ni, Al and As) were
observed, and the controlling geochemical processes are discussed. The conceptual model for water and
solute movement was applied to quantify the contribution of released trace elements to a downstream lake in
the permafrost catchment.
1. Introduction
Carbon emissions from thawing permafrost wetlands are an expected global consequence of a warmer
climate. It is crucially important, and a key objective of this paper, to understand the pathways for such
carbon emissions. However, one should not discount the local consequences for the freshwater resources and
ecology in the Arctic, that global warming might also bring along. The Arctic foresees an increased number
of industrial facilities, mining and oil extraction operations, and increases in traffic and population, all of
which may cause pollution of the water resources of hosting permafrost catchments, their wetlands, streams
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and lakes. Current increases of ~1 cm/yr in the depth of the active layer, i.e. the surficial layer of permafrost
areas that thaws during summer, have been observed in Greenland (Elberling et al., 2010). This implies a
rapidly growing importance of the hydrogeological and geochemical processes taking place in the active
layer, on the water resources of permafrost catchments.
The seasonal freeze-thaw dynamics of a permafrost catchment obviously is a very important
control on its water resource. Kane et al. (1989) and Hinzman et al. (1991) provided excellent studies of the
effects of seasonal freeze-thaw dynamics on the lateral and vertical groundwater flow in an active layer.
Lateral groundwater movement is restricted in winter, when the active layer is frozen. In spring, snow melt
and surface runoff of the melt water is followed by a gradual thawing of the active layer, hence making water
in the active layer part of the hydrological cycle as opposed to its frozen winter-state. Thawing takes place
from the surface and downwards. The thawing draws latent heat mostly from above the thawing front, due to
the increasing solar radiation at this time of year, in combination with the absence of a reflective and
insulating snow cover. Near the onset of winter, initial snow fall melts soon after it contacts the surface,
drawing heat from the active layer. The active layer first cools from above and below to become isothermal
(~0 ºC) in its entire depth. This is followed by a phase change to ice, releasing latent heat. The phase
transition typically occurs mainly from the surface downwards (Fox, 1992), but can occur simultaneously
from the surface downwards and from the permafrost table upwards (Michel, 2011). Some researchers
suggest a more uniform freeze-up, where the phase transition occurs more or less simultaneously at all
depths in the active layer.
Despite the significant cooling of the active layer during the Arctic winter, a fraction of the
pore water remains unfrozen (Anderson et al., 1973; Hinzman et al., 1991; Freeze and Cherry, 1979). The
size of this fraction generally increases with decreasing grain size, and become quite substantial in silts and
clays, amounting to one tenth to one fifth of the pore water. For organic peat soil and an underlying silt
mineral soil, Hinzman et al. (1991) reported an unfrozen water content of ~7 % and ~16 % of the pore
water, for soil temperatures below -10 ºC. The saturated hydraulic conductivity of sediments decreases by
several orders of magnitude early after the onset of phase transformation to ice (Freeze and Cherry, 1979).
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Dissolved major ions and trace elements in the active layer will be affected by the freeze-thaw
dynamics. Ice formation alters mineral saturation indices, and dissolved CO2 together with other dissolved
gases may be excluded to bubble inclusions (Killawee et al., 1998; Papadimitriou et al., 2003). These
processes may cause mineral precipitation or dissolution which, due to kinetic constraints, may or may not
be fully reversible during the following spring thaw and summer. The freeze-thaw dynamics may also cause
freeze-out of ions and hence separate, to some extent at least, the processes of water flow from solute
transport. The extent of ion freeze-out depends on soil grain size and freezing front velocity. Ions may be
either concentrated in front of (sands and light sandy silts) or behind (clays and silty clays) the advancing
freezing front (Lundin and Johnsson, 1994; Chuvilin, 1999; Kokelj and Burn, 2005). The freeze-out of ions
may become absent in case of a sufficiently rapid fall freeze-up.
The aim of the present study is to increase our general understanding of water and solute
transport in permafrost wetlands, and in particular, to elucidate possible implications for inorganic carbon
cycling. As a starting point, we assess the hydrology of the active layer. The assessment is based on classical
hydrogeological methods in combination with the vertical distribution of both stable isotopes of water, and
of dissolved chloride, representing tracers of water and solute movement, respectively. This leads to a
conceptual model for water and solute transport in the wetlands active layer, which is then applied to the
major water chemistry, comprising inorganic carbon. Results for trace elements are finally included to serve
a baseline for the assessment of future impacts of increased anthropogenic activities in the Arctic.
2. Study area
The study site is located in a wetland (69º12’52”N, 51º05’45”W) about 300 m east of the city border of
Ilulissat, Greenland (Fig. 1). The wetland is a low relief part of the catchment area of a series of connected
lakes, used for drinking water by the local water works. The active layer of the wetland is saturated in most
of its depth, with the water table typically a few centimeters below ground. A network of natural drainage
channels are distributed over the wetland, visible on Fig. 1. The vegetation primarily consist of sphagnum
moss, with occurrence of dwarf birch (Betula nana), willow (Salix sp.), cotton-grass (Eriophorum sp.),
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northern Labrador tea (Ledum palustre), Crowberry (Empetrum nigrum) and in addition lichens, mosses and
grasses. The geology of the distal high relief part of the catchment area comprises granodioritic and quartz
dioritic gneiss, with basic agmatite, and garnet- and diopside-bearing amphibolite rocks. In the proximal low
relief part of catchment area, the bedrock is covered by a lower marine mud unit deposited under
glaciomarine conditions between deglaciation c. 9.5 ky BP and emergence of the area c. 7 ky BP and an
upper peat unit accumulated after the area emerged from the sea (Weidick, 1969; Weidick and Bennike,
2007).
A climate record for 2004-2011 from the nearby Ilulissat Airport weather station, shows a
polar tundra climate, with a mean air temperature of -4 ºC and an annual precipitation of 260 mm. Average
daily temperature exceeds 0 ºC during a four and a half months summer period, from the middle of May to
the end of September. The period with an average daily minimum temperature exceeding 0 ºC is one month
shorter. The mean air temperature during the four and a half months summer period is +5 ºC. Precipitation
during the same period is 125 mm, of which half is rain and the other half is snow.
3. Methods
The field campaign was conducted during 17-26 August 2010. Observations were made mainly along a 400
m transect, extending from the hydraulically up-gradient margin of the wetland to a down-gradient lake (lake
39; Fig. 1). A cross section of the transect is shown in Fig. 2. In three locations along the transect (A, B and
C in Figs. 1 and 2), samples of pore water and sediment were collected in vertical profiles, as elaborated
further below. The profiles A, B and C are positioned hydraulically up-gradient, central, and down-gradient,
respectively, relative to each other. Active layer depth and lithology was investigated using hand drilling
equipment for every 10 m along the transect. The active layer temperature was measured as a function of
depth at location A (Fig. 2), by a pH meter connected with a cable to a digital thermometer, which was
carefully installed using hand drilling equipment.
3.1 Active layer pore water sampling and field analysis
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Samples of active layer pore water, all from the saturated zone, were collected using PRENART Super
Quartz (Teflon) suction cups (Supplementary Information). Each suction cup was connected to a 1 L
polypropylene (PP) bottle using nylon vacuum tubing. A 50 mL, acid washed, glass septum flask, equipped
with a 10 mm rubber stopper and two syringe needles, was placed in-line before the 1 L PP bottle, to obtain a
sample least possibly exposed to atmospheric air. The septum flask, PP bottle and connecting tubing were
flushed thoroughly with N2 gas prior to sampling. A pressure relative to atmospheric of -0.4 atm, maintained
by a vacuum pump, for 36-48 hours was required to extract samples of up to a few hundred mL, including
the 50 mL septum flask samples. At the end of sampling, the content of the PP bottle was passed through a
0.21 µm cellulose-acetate (CA) syringe filter (Sartorius Minisart) into several 20 mL polyethylene (PE) vials,
of which one received 1vol% of 7 M HNO3 suprapur, one was frozen, and the rest were kept refrigerated. EC
was measured by a HACH electrode; reported values correspond to EC at 25 ºC. Alkalinity was determined
by Gran-titration (Stumm and Morgan, 1981) on filtered sample. pH was determined on a few mL of sample
extracted from the septum bottles with a syringe and a needle and dripped onto pH indicator strips with a 0.2
pH unit resolution (MColorpHast™).
3.2 Sediment and permafrost pore water sampling
Active layer peat and sediment were excavated at pore water sampling sites (A, B and C, Fig. 1) at the end of
the pore water collection (Section 3.1) and transferred into 0.5 L blue cap glass bottles which were
immediately flushed with N2 and capped. Peat samples were stored frozen, and loam samples refrigerated.
Samples of the permafrost were obtained by motorized Stihl earth auger with a 40 cm long Hilti diamond
core bit (OD 57 mm). Retrieved cores were sealed in PVC tubes (OD 63 mm) and stored frozen. In the
laboratory, the sectioned core pieces, typically 10-20 cm in length, were split along the vertical axis, and the
one half transferred to a 500 mL centrifuge tube, melted at room temperature and centrifuged 10 min at 8000
g. The supernatant was doubly filtered through first a 0.45 µm cellulose filter and then a 0.2 µm Acrodisc
filter into 50 mL centrifuge tubes and stored refrigerated. The permafrost sediment samples, i.e., the residues
of the 500 mL centrifuge tubes, were stored frozen until analysis. Of the remaining core halfs, permafrost
pore water samples for 18O and D analysis were obtained by discrete subsampling of 1 cm-pieces, allowing a
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~1 cm depth resolution. The subsamples were transferred to 25 mL centrifuge tubes, melted at room
temperature and centrifuged. 1 mL of the supernatant was transferred to a glass vial for immediate analysis
(Section 3.3). Subsamples for determination of dissolved element concentrations were filtered and
transferred to 15 mL polypropylene centrifuge tubes (VWR, Denmark) and acidified to 1vol% HNO3 (J.T.
Baker, Baker Instra-analysed).
3.3 Laboratory water analysis
Active layer and permafrost pore water concentrations of major cations (calcium, Ca2+; magnesium, Mg2+;
sodium, Na+; potassium, K+) and of anions and ammonium (chloride, Cl-; sulphate, SO42-; nitrate, NO3
-;
ammonium, NH4+) were measured by Ion Chromatography (IC) (Metrohm, Interface 830) on refrigerated,
non-acidified samples and frozen non-acidified samples, respectively. For major cations, and anions incl.
NH4, respectively, a Metrosep C4250 and a Metrosep A supp5 column was used. Silicate (Si) and phosphate
(PO43-) concentrations were measured by spectrophotometry (Shimadzu UV-1800) on non-acidified samples
following the procedures in Danish Standard Association (2004) and Limnololgisk Metodik (1977),
respectively. Aluminum (Al) concentrations were measured on acidified samples by flame Atomic
Absorption Spectrometry (AAS) on a Perkin Elmer AAnalyst 400. Iron (Fe), nickel (Ni2+), arsenic (As) and
manganese (Mn2+) concentrations were determined using Inductively Coupled Plasma Mass Spectrometry
(ICP-MS) on an Agilent 7500C. Dissolved organic carbon (DOC) was measured on the refrigerated septum
flask samples by a Shimazdu TOC-VCN Analyzer. The pH and alkalinity of the permafrost pore water was
determined, respectively, by pH strip immediately after thawing and by Gran-titration on filtered,
refrigerated, non-acidified samples. δ18O and δD values, reported in the VSMOW/SLAP scale, were
determined on a Picarro Cavity Ring-Down Spectrometer (CRDS) L2120-i. Standard deviations were equal
to or lower than 0.21‰ for δ18O and 0.66‰ for δD.
3.4 Sediment analysis
The fractions of clay (<2 µm), silt (2-20 µm), fine sand (20-200 µm), and coarse sand (200-2000 µm) were
determined by a combination of dispersion, sieving and sedimentation according to Day (1965). For soil
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classification the USDA soil class system was used. Porosity and bulk density were calculated for the water
logged active layer samples of loam, by assuming a degree of saturation of 1 and using the gravimetric water
content, determined by 24 h drying at 105 °C, and a particle density of 2.65 g/cm3. Total soil organic matter
(TOC) was determined by dry combustion and subsequent quantification of evolved CO2 by infrared
spectrophotometry.
The cation exchange capacity (CEC) was determined by three displacements with 1 M NH4Cl
on 5-8 g sediment. CEC was determined as the sum of Na+, K+, Ca2+, Mg2+, Fe2+ and Al3+ equivalents in 0.21
µm CA filtered extracts. Four samples, collected at depths with highest ratio of dissolved NH4-to-total
dissolved ions were subjected to displacement by 1 M KCl, which indicated that NH4+ constituted <3% of
the CEC. The significance of possible Al(OH)3 precipitation was determined as described by Kjøller et al.
(2004). For extracts supersaturated for amorphous Al(OH)3, decreases in the OH- concentration were
equivalent to Al losses of <0.2% of the dissolved Al concentration measured in the first displacement.
3.5 Speciation of water chemistry and exchanger composition
PHREEQC-2 (Parkhurst and Appelo, 1999), with the wateq4f.dat database, was used to calculate saturation
indices (SIs) and 𝑃CO2, as well as the equilibrium exchanger composition (Thomas-Gaines convention). Input
data were measured chemical composition and temperature (Profile A) of pore waters, and CEC of sediment
samples.
3.6 Hydrologic conductivity and Darcy velocities
Water table measurements and slug tests were conducted in eight screens along the transect
(Fig. 1) installed at dry locations (defined as locations with a water table below ground). The screens were
installed over the entire active layer depth. Additional water table measurements were obtained in wet
locations (i.e., locations with an above-ground water table) by reference to stakes pounded into the active
layer. The wet locations included three poorly defined natural channels intercepting the transect. Saturated
hydraulic conductivities of peat were interpreted from slug test data using the Bouwer-Rice method, and
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assuming flow only in the saturated peat. For each tested screen the data allowed for the calculation of a
minimum and maximum hydraulic conductivity. Saturated hydraulic conductivities of the active layer loam
were obtained by Rosetta Lite 1.1 as implemented in Hydrus-1D (Radcliffe and Šimůnek, 2010), using
measured texture and bulk density as input parameters (Stumpp et al., 2009). For comparability with slug test
values hydraulic conductivities returned by Rosetta Lite were corrected to 2 ºC. Lateral flow velocities and
fluxes were calculated from Darcy velocities divided by porosity, applying mean values of measured
hydraulic conductivity, gradient, saturated depth and porosity. However, a porosity of the peat of 0.8 was
assumed (not measured) as obtained from a weighted average of data from Table 4 in Hinzman et al. (1991),
which agree well with data of other studies (Hayashi et al., 2007).
4. Results
4.1. Geology and sediment characteristics
The distribution of peat along the transect is shown in Fig. 2. A peat layer thickness of 0.1 to 0.5 m was
observed, with a mean and median of 0.24 m. The peat layer both became more decomposed and stratified
with increasing depth. The peat is underlain by a non-stratified, grey, marine mud with a texture ranging
from silty clay, over clay loam and silty clay loam, to loam; henceforth collectively referred to as loam. The
boundary or transition from the peat layer to the loam was quite sharp and easy to identify. No variation in
the texture of the loam was observed across the permafrost table. The porosity of the loam in the active layer
ranges from 0.42 to 0.60 (n = 12) with a mean of 0.47 and median of 0.46. Beneath the permafrost table
alternating 0.5-2 cm thick layers of loam and ice (free of sediment) were observed typical for ice-rich
permafrost (Supplementary Information). Recorded thaw depths are also shown in Fig. 2. At the time of data
collection (mid-to-late august), the active layer will be nearly fully developed. The recorded active layer
thickness varied from 0.24 to more than 1 m, with a mean and median of ~0.50 m. Thaw depth was inversely
related to the depth to the water table, the latter expressing the unsaturated peat thickness which is an
important insulator inhibiting heat transfer to the permafrost table (Hayashi et al., 2007). The temperature-
depth relationship, determined at profile A only, was 4.5 ºC at 15 cm depth, decreasing almost linearly to 3.5,
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2.3, 1.0 and 0.4 ºC at, respectively, 22, 30, 43 and 50 cm depth, the latter being a few cm above the
permafrost table.
Total organic carbon contents of the surface peat ranged from 25 to 44% of dry weight (n =
29), with a mean of 37% and median of 38%. The observed peat TOC values therefore indicate that the peat
consist of a nearly 100% organic matter (e.g., CH2O with 40% carbon by weight) without significant mixing
with the underlying mineral loam. In the active layer loam, generally TOC contents ranged from 1.3 to 3.0%,
with mean and median values of 2.3% and 2.4%, respectively (n = 28). Loam samples underneath the
permafrost table had lower TOC values ranging from 0.2 to 2.1%, with mean and median values of 0.9% and
1.1%, respectively (n = 34). Locally higher TOC contents (7 to 17%) were observed in the loam, indicating
frost churning mixing pieces of peat downwards to near the permafrost table (Hinzman et al., 1991). Frost
boils were locally observed.
4.2. Hydrology
The results of water table measurements are included in Fig. 2. Overall, an average lateral hydraulic gradient
of 7.1‰ was observed between the water table measurements at 10 m and 360 m distance. A higher
hydraulic gradient is observed between 360 m distance and the lake water surface at 410 m distance. The
water table at the abovementioned dry and wet locations (Section 3.6) was typically within a few centimeters
from the surface. The down-slope hydraulic pressure gradient was observed independently on dry versus wet
location type, i.e., a down-slope hydraulic pressure gradient driving flow towards the lake was observed up-
slope, within, and down-slope the channels and other wet locations. Subsequent to a rain event on 23 August
2010 a down-slope hydraulic gradient was even sustained. The channels are believed to be important
conduits for water during snowmelt. Reddish iron stains up to a meter above ground level were observed on
rocks in depressions bounding the wetland area, and may indicate significant seasonal water level
fluctuations (Supplementary Information).
Hydraulic conductivity values for the saturated peat range from 3.1×10-6 to 2.4×10-4 m/s, with
a mean of 9.1×10-5 m/s and median of 5.8×10-5 m/s, as calculated from slug test data collected in eight
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screens located along the 400 m transect. At all locations the water table was positioned in the peat layer,
resulting in saturated thicknesses of the peat ranging from 4 to 19 cm, with mean and median values both
close to 12 cm. Hydraulic conductivities obtained by Rosetta Lite using observed loam textures and bulk
densities range from 4.1×10-7 to 1.9×10-6 m/s, with a mean of 8.1×10-7 m/s and median of 6.0×10-7 m/s.
4.3. Stable isotopes
The composition of δ18O of water from the active layer and from melted sections of the permafrost cores is
shown in Fig. 3. Active layer water in the up-gradient profile, A, was most depleted for the heavy 18O-
isotopes, with δ18O-values increasing with depth from -17.4‰ in the peat to -17.0‰ near the permafrost
table. The δ18O-values then increase abruptly to -15.7‰ in the shallowest sample from the permafrost.
Active layer water in the central and down-gradient profiles, B and C, are increasingly less depleted for 18O
relative to profile A. At the shallowest sample depths, profile B and C has δ18O-values of -16.2 and -13.7‰,
respectively. Beneath the permafrost table, δ18O in all three profiles converge to -15 to -16‰.
4.4 Water chemistry
Field measurements of the electrical conductivity (EC) of active layer pore water collected from the three
profiles A, B and C is shown in Fig. 4. EC-values of 98 to 121 µS/cm are observed at the shallowest depths
in the peat. EC increases with depth to up to 293 µS/cm in the up-gradient profile A, 436 µS/cm in profile B
and 718 µS/cm in the down-gradient profile C, at the deepest sampling depths of the active layer. In addition
to data collected in the catchment, we also obtained values of EC of the drinking water distributed from the
downstream water works. These values typically are around 100 µS/cm, i.e. corresponding to the values
observed in the peat. The EC of a water sample expresses the sum of dissolved ions, and accordingly
increasing concentrations with depths are observed also for the major cations Ca, Mg, Na and K and for
bicarbonate, also shown in Fig. 4. The concentrations of the dominant cations Ca and Mg, balanced by
bicarbonate, peak in the deepest part of the active layer, then show an abrupt decrease near the permafrost
table and a subsequent gradual increase with depth below the permafrost table. In the down-gradient
direction, from profile A through C, progressively higher concentrations in the active layer of Ca, Mg and
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bicarbonate (reaching 1.8 mM, 2 mM and 8 meq/L, respectively, in profile C) are observed, similar to the
trend described for EC. The concentrations of Na and K show a gradual increase across the permafrost table,
approaching 1 and 0.3 mM, respectively. For Cl, no increase with depth is observed (Fig. 4) in the active
layer. The concentration of Cl is around 0.3 mM in profile A and 0.4 mM in profiles B and C. Below the
permafrost table the concentrations of Cl are slightly higher, except for a peak at 1.4 mM at 75 cm depth in
profile C. Acidic pH in the range of 4.7 to 5 was observed in the peat, increasing with depth to neutral or
slightly alkaline values (pH 6-7.7). Lower pH values were observed in profile A compared profiles B and C.
4.4.1. Fe, SO4 and NH4
In peat of the active layer, Fe concentrations of 0.006 to 0.04 mM are measured (Fig. 4). Very high Fe
concentrations of up to 0.8 mM are measured in the upper and middle part of the loam, decreasing at the
bottom of the active layer. Beneath the permafrost table, Fe concentrations in profile A of up to 0.001 mM,
in profile B up to 0.016 mM and in profile C up to 0.48 mM are measured, decreasing with depth in profile B
and C (Fig. 4). Sulfate concentrations up to 80 µM are measured in the peat and upper part of the loam. In all
profiles the SO4 concentration shows a minimum just above the permafrost table. Below the permafrost table
SO4 concentrations first increase, and then decrease with depth, though in profile C, a second SO4 increase
below 110 cm depth is measured. NH4 concentrations (not shown) were below 0.03 mM in the peat, and
increased with depth in the active layer loam to up to 0.1 mM near the permafrost table. Higher NH4
concentrations of typically 0.1-0.2 mM were detected beneath the permafrost table.
4.4.2. Trace elements and DOC
Concentrations of Al, Ni, As and Mn (assumed to be Mn2+) are shown in Fig. 5. The Al concentrations in the
active layer range from below detection limit (0.004 mM) to 0.045 mM, with some tendency for relatively
higher values in the peat compared to the loam. However, much higher Al concentrations, up to 0.5 mM, are
measured immediately below the permafrost table. In the frozen zone, the Al concentrations decrease with
depth. The concentrations of Ni in the peat range from 0.15 to 1.0 µM. In all profiles, the concentrations of
Ni peak in the active layer loam, reaching a concentration of 1.8 µM at 35 cm depth in profile B, and
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generally decrease below the permafrost table. The variation in As concentrations with depth in the active
layer vary much between each profile and it is difficult to recognize a general trend for all three profiles. In
profile A and B, As is detected in the peat in concentrations up to 0.1 µM. In the active layer loam As
concentrations are below the detection limit (0.013 µM), except in profile B at 42 cm depth where a 0.28 µM
As concentration is measured. In contrast, in profile C, As concentrations are below detection limit in the
peat, and increase to 0.05 to 0.11 µM in the active layer loam. Below the permafrost table, As concentrations
decrease with depth in all three profiles. Mn concentrations up to 22 µM were observed. The distribution of
Mn is much similar to that of Ca, as confirmed by positive correlation of Mn with Ca (r2=0.71). Dissolved
organic carbon (DOC) is included in Fig. 5. Generally, DOC concentrations in the active layer range from 1
to 5 mM, and increases with depth in the active layer loam. Much higher DOC concentrations, up to 55 mM,
are measured below the permafrost table.
4.4.3. Mineral equilibria and 𝑃CO2
Saturation indices (SIs) for calcite (CaCO3), gibbsite (Al(OH)3), siderite (FeCO3) and chalcedony (SiO2) are
shown in Fig. 6. SI for calcite approaches 0 (saturation) with increasing depth in active layer. Across the
permafrost table, SIcalcite decreases quite abruptly, then continuously increase to near saturation at ca. 100 cm
depth. Gibbsite in the active layer is subsaturated (SIgibbsite<0) or close to saturation in profile A and C, and
increasingly supersaturated in profile B. Beneath the permafrost table all profiles show supersaturation for
gibbsite. Siderite is subsaturated in the peat, and approaches saturation with increasing depth in the active
layer loam. Below the permafrost table, pore waters in profiles B and C are close to saturation for siderite,
while profile A with a single point is subsaturated. SIchalcedony calculated for the only available Si
measurements (profile C) is slightly above 0 indicating supersaturation. Fig. 6 includes the calculated partial
pressure of dissolved CO2, 𝑃CO2, which are generally above 0.25 atm in the active layer, and diminutive in
samples of melted permafrost. As pH determined by indicator strips might be thought to be inaccurate, the
PHREEQC speciations of SIs and 𝑃CO2 (Fig. 6) were repeated using pH values both 0.3 unit lower and higher
than those measured by strip (Fig. 4). The grey areas in Fig. 6 indicate the span thus obtained in the
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calculated values; the effects are especially marked for 𝑃CO2, which is strongly pH dependent. However, the
overall results and derived conclusions remained in accordance with those described above.
4.5 Cation exchanger composition
The measured cation exchanger compositions of loam sediment samples in the profiles are shown in Fig. 7
collectively for all three profiles because of an indiscernible inter-profile variation. The exchanger is
dominated by Ca and Mg ions constituting, respectively, 4 to 6 and 2 to 4 meq/100 g in most depths. The
monovalent cations Na and K take up a much smaller amount of exchange sites, respectively, 0.1 to 0.2 and
0.25 to 0.1 meq/100 g, increasing with depth. Al constitutes 1 to 2 meq/100 g in the upper part of the active
layer loam. At depths greater than 30 cm, measured exchangeable Al is generally about 0.5 meq/100 g, with
possible exceptions at 33 and 50 cm depth, where Al diminishes on the exchanger. Exchangeable Fe2+
amounts to a substantial 2 to 3 meq/100 g in the active layer loam and decreases abruptly to <0.1 meq/100 g
across the permafrost table. An average CEC of 10.8±2.5 meq/100 g was calculated by summation of the
exchangeable cations shown in Fig. 7. Fig. 7 includes also two modeled cation exchanger compositions
which will be further described in the discussion, Section 5.3.
5. Discussion
5.1 Conceptual model for water and solute movement the active layer
5.1.1 Water movement in the active layer
As a starting point we will outline the seasonal groundwater flow dynamics in the transect. A preliminary
hydrological assessment can be made by generalizing the work on permafrost hydrology by Hinzman et al.
(1991) and others. Hence, at Ilulissat, snow melt takes place in mid-May, when the average daily air
temperature exceeds 0 ºC. This is followed by thawing of the active layer from the surface and downwards,
where progressively more of the active layer becomes hydraulically active. The groundwater flow system
remains shallow, even at maximum active layer depth, because the permafrost table acts as an impermeable
bottom. The freeze-up process is initiated in September, when the daily average air temperature falls to
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below 0 ºC, ultimately causing the hydraulic conductivity of the active layer to decrease by several orders of
magnitude. Because the active layer is saturated in most of its depth, movement of residual un-frozen pore
water during winter must be strongly restricted as pores are blocked by ice (Fox, 1992; Freeze and Cherry,
1979).
According to the above, the conditions for significant lateral groundwater flow exist only
during a four-month summer period. The horizontal hydraulic pressure gradient throughout the thawed
period must be close to the observed value of 7.1‰ (Fig. 3), as it is restricted by the shallow depth of the
active layer and the general slope of the ground surface. Using the mean hydraulic conductivities of the peat
and loam reported above (Section 4.2), this gradient corresponds to average lateral pore water flow velocities
of 25 and 0.4 m/yr, in the peat and loam, respectively. Assuming a four-month summer, these velocities
correspond to annual lateral travel distances of 8.3 m in the peat and 0.13 m in the loam (Fig. 8).
Accordingly, flushing of the peat occurs at a much faster pace than flushing of the loam.
Fig. 3 shows that active layer water becomes increasingly enriched in 18O in the down-
gradient direction. Precipitation distributed over the small study area must have a uniform 18O isotopic
composition. The enrichment in Fig. 3 therefore is consistent with evaporation of water from the active layer
as the water travels down-gradient towards the lake, and hence suggests lateral hydraulic connectivity
(Supplementary Information). An evaporative loss of ~15% of the water would correspond to the observed
~4‰ δ18O enrichment from profile A to C, taking into account the prevailing relative air humidity of 50-70%
(Clark and Fritz, 1997).
The 18O isotope data (Fig. 3) may be used for assessing characteristics of the fall freeze-up
process. When ice forms, water molecules containing heavy isotopes 18O and D become preferentially
included in the newly formed ice, depleting the unfrozen residual water. In a closed and continuously mixed
water reservoir freezing from above, a possible proxy for a freezing-up active layer, the δ18O distribution in
the ice can be described by a Rayleigh distillation, according to the equation (Fritz et al., 2011):
δi = δ0 + 1000·ln αi-w·ln f + 1000·ln αi-w Eq. (1)
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Where δi is the δ18O value of the ice, δ0 is the initial δ18O value of the water reservoir, αi-w is
the ice-water fractionation coefficient, and f is the fraction of unfrozen residual water, which is equivalent to
the relative depth of the advancing freezing front. αi-w is close to 1.003 for equilibrium conditions, and
decreases with increasing freezing velocity, corresponding to an increasing degree of non-equilibrium
(Lehmann and Siegenthaler, 1991). In a porous medium, the medium impedes, but does not completely
block, the mixing of the residual water by diffusion and dispersion (Michel, 2011). On the other hand,
thawing of the ice takes place without isotopic fractionation, due to the impermeability of ice. In Fig. 3 the
calculated δ18O distribution in the ice according to Eq. (1) is included. To ensure readability of the graph, an
arbitrary initial δ18O value of the bulk water of -14.3‰ and a reduced αi-w of 1.0005 were used. Comparing
the modeled profile with the measured profiles, is it seen that none of the δ18O profiles in Fig. 3 indicate a
relation to the Rayleigh distillation curve form. This suggests a rapid fall freeze-up of the active layer (i.e.,
no fractionation as αi-w approaches unity) and/or a uniform, simultaneous pore water phase transition at all
depths in the active layer. These results have implications for the vertical water and solute transport since
also freezing-out of ions (Lundin and Johnsson, 1994; Kokelj and Burn, 2005) diminishes with increasing
freezing front velocity (Killawee et al., 1998), and would not effectively cause vertical redistribution of ions
in a uniformly freezing active layer. This conclusion, however, does not rule out that vertical water
movement and vertical mixing of pore water in the active layer may take place due to molecular diffusion,
cryoturbation, advective flow driven by vertical hydraulic pressure gradients, by transversal hydrodynamic
dispersion coupled to the above calculated lateral flow, and by temperature gradients (Loch and Kay, 1978;
Hanley and Rao, 1980; Chuvilin, 1999).
5.1.2 Solute transport in the active layer
The distribution of Cl can be used to test the above implication of the 18O isotope distribution, that freeze-out
of ions is not an important control on vertical solute transport in the studied active layer. It can be assumed,
that Cl moves conservatively and does not participate in reactions involving dissolution or precipitation. In
Fig. 4, except for one unexplained peak at 78 cm depth in profile C, the concentration of Cl is nearly constant
with depth. The opposite, a variation in Cl concentration with depth, would be expected if vertical solute
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transport was controlled by ion freeze-out (Lundin and Johnsson, 1994; Chuvilin, 1999; Kokelj and Burn,
2005). On the other hand, increases in the concentrations of other major ions with increasing depth are
observed. The discrepancy between the behavior of Cl and that of the other major ions is consistent with a
conceptual model where mineral dissolution reactions take place in the loam, followed by upwards diffusion
of the reaction products to a frequently flushed surface peat layer (Fig. 8). In support, the EC of treated lake
water distributed from local water works (~100 µS/cm) corresponds to that observed in the peat (Fig. 4),
suggesting an insignificant direct efflux of pore water and solutes from the active layer loam to the lake.
Rather, significant water input to the lake comes from peat layer efflux, snowmelt surface runoff, and direct
precipitation. Flushing of the loam is comparatively slow, although it may still affect observable solute
concentrations to some significant extent, especially in the upper part of the active layer loam which thaws
earlier than the lower part.
5.2 Major pore water chemistry
The pore water is a Ca-Mg-HCO3-type (Fig. 4) suggesting carbonate dissolution to be an
important control on the major pore water chemistry. The pore water generally shows a 1:1:4 stoichiometry
of Ca to Mg and alkalinity, which is in accordance with that of congruent dolomite dissolution by
consumption of carbonic acid:
CaMg(CO3)2 + 2H2CO3 → Ca2+ + Mg2+ + 4HCO3- Reaction 1
Most pore water samples showed subsaturation for calcite (Fig. 6) and dolomite (not shown), allowing for
the dissolution of these carbonates, if present. It should be noted, that the stoichiometric accordance with
dolomite dissolution is not evidence for the presence of dolomite in the active layer. In Reaction 1, the
CaMg(CO3)2-term can be substituted by any suite of CaCO3 and MgCO3 minerals reacting in equal amounts
with no effects on the right-hand side of the reaction. Alternatively, the major chemistry might instead be
explained by incongruent Ca- and Mg-rich silicate weathering, as supported by the host rock chemistry of
garnet and diopside-bearing amphibolite and gneiss with basic agmatite. Precipitation of the Si released by
silicate weathering is thermodynamically feasible as the pore waters appear to be supersaturated for
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chalcedony (Fig. 6). In the active layer, the strongly elevated partial pressure of CO2, 𝑃CO2, shown in Fig. 6
indicates ongoing carbonic acid production, as 𝑃CO2 in this shallow system would otherwise equilibrate with
the atmospheric 𝑃CO2 of ~0.0004 atm. Consequently, pore water pH may be controlled by a balance between
the rates of mineral dissolution, increasing alkalinity by proton consumption, and organic matter oxidation,
adding acidity as carbonic acid.
5.2.1 Inorganic carbon cycling
The abrupt decreases of alkalinity, Ca, Mg and 𝑃CO2 (Figs. 4 and 6) observed from right above to right below
the permafrost table indicates that these parameters are mutually controlled via the carbonate system. The
size of the decreases of alkalinity, Ca and Mg is stoichiometrically in accordance with that of dolomite
precipitation, as could be described by reversing the arrow in Reaction 1. The occurrence of this data pattern
near the permafrost table indicates a relation to the water-ice phase-change. Our observations are partly
supported by studies by Killawee et al. (1998) and Papadimitriou et al. (2003), who observed increasing
saturation index for calcite and, ultimately, precipitation of cryogenic calcite (not dolomite or equal amounts
of CaCO3 and MgCO3) as larger proportions of their experimental solution froze. Concurrently with solution
freezing, the 𝑃CO2 increases (in accordance with Reaction 1) causing expel of CO2 to bubble inclusions
forming in the ice (Lipp et al. 1987). In this respect, the diminutive 𝑃CO2-values observed below the
permafrost table may be an artifact due to our method, which allowed loss of CO2 from inclusions in the
thawing permafrost core samples to the container head space. Cryogenic carbonates potentially formed might
re-dissolve if CO2 loss had been hindered and appropriate time allowed for their dissolution. Hence, the data
suggests an annual cycling of inorganic carbonate, Ca and Mg, i.e. carbonate mineral precipitation and CO2
bubble inclusion formation during fall freeze-up and their dissolution following spring thaw. In this scheme,
the primary source of Ca and Mg may still be silicate weathering. Shallow groundwater in permafrost of
North America (Williams and van Everdingen, 1973) and non-glacial drainage on Svalbard (Tye and Heaton,
2007) are reported to be generally of Ca-HCO3 or Ca-Mg-HCO3 type, indicating that the proposed cycling
can be a common process in permafrost regions. The concentrations of Ca, Mg and HCO3 reported by Tye
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and Heaton (2007) increased over the course of the summer; this would be consistent with a kinetically
constrained re-dissolution during summer, of cryogenic carbonates formed during fall freeze-up.
The very high concentrations of Fe (Fig. 4) combined with the subsaturation to near-saturation for siderite
(Fig. 6) indicates that siderite is also formed cryogenically, and hence may cause cycling of Fe2+ in a similar
scheme to that describe above for Ca and Mg. Lipson et al. (2010), in a drained thaw lake basin near Barrow,
Alaska, also observed highly elevated dissolved Fe concentrations.
5.2.2 Gas emissions by ebullition
The atmospheric carbon emission related to thawing permafrost has been the focus of several recent studies
(Schuur et al., 2009; Zimov et al., 2006). In relation to this, the observed high 𝑃CO2-values in the pore water
shown in Fig. 6 call for a further notion. In the pore water, formation of gas bubbles is feasible once the sum
of partial gas pressures exceeds a pressure of just above 1 atm, due to the shallow saturated thickness in the
active layer. Given that the partial pressure of N2 at equilibrium with the atmospheric N2 pressure is 0.78
atm, 𝑃CO2 exceeding 0.22 atm as shown in Fig. 6 is a strong indication of potential bubble formation.
Therefore, emission of CO2 from the active layer via ebullition is principally possible. While elevated CO2
pressures may be the primary cause for the initial bubble formation, once a bubble forms also other dissolved
gases such as methane (CH4), if present, will degas to the bubble. Hence ebullition may be a pathway for
CO2 and other greenhouse gases, from water saturated active layers. Ebullition as a pathway for gas
emissions from permafrost wetlands has been suggested by Sachs et al. (2008), based on observed short
duration emission pulses. CH4 ebullition has been modelled by Walter and Heimann (2000). However, their
model did not take into account the contribution of the partial pressure of CO2 to the total sum of partial gas
pressures.
An increase in DOC across the permafrost table as observed in this study (Fig. 5) was reported
also by Elberling et al. (2010). The observations suggest that the permafrost stores a pool of reactive DOC.
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In case of an increased thaw depth, the DOC liberated to the active layer from the newly thawed permafrost
may ultimately transform to gaseous CO2 and CH4.
5.2.3 Redox chemistry
A significant part, in not all, of the dissolved iron can be assumed to be in the reduced form as Fe2+ (Fig. 4)
(Lipson et al., 2010). Therefore, the presence of Fe indicates anoxic conditions already few centimeters
below the water table. In relation to the above described conceptual model for solute transport, the variation
of Fe with depth indicates an iron release in the active layer loam, followed by diffusive transport upwards to
the peat layer and downwards towards the permafrost table. However, the very high concentrations of Fe
(Fig. 4) combined with the subsaturation to near-saturation for siderite (Fig. 6) indicates that siderite is also
formed cryogenically, and hence may cause cycling of Fe in a similar scheme to that describe above for Ca
and Mg. The decreasing concentrations of SO4 with depth suggest downwards diffusion from the peat layer
to the loam. Consistently, proximity to the sea suggests that sea water aerosols is a probable source of SO4
(and Cl) due to prevailing winds from N and W during summer. The depletion of sulfate and decreasing
concentrations of Fe in the lower part of the active layer loam could indicate sulfate reduction followed by
Fe sulfide precipitation. However, the decrease in the concentration of Fe exceed that of SO4 by a factor of
four in profiles A and B, and more in profile C. Therefore, if the iron-sulfide precipitate formed is FeS
(mackinawite), an additional Fe sink must exist.
5.3 Trace elements
An important implication of the conceptual solute transport model presented above (Fig. 8) is that the
transport of dissolved trace elements from the loam to the peat will be by diffusion also. Consequently, the
down-gradient efflux of the trace elements released in the loam to the lake (lake 39; Fig. 1) will be controlled
by the upwards diffusive flux to the peat and subsequently by the lateral flow velocity in the peat. The study
area is a protected drinking water catchment, and the primary source of Al, Ni, As and Mn2+ therefore must
be geogenic.
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Arsenic strongly sorbs to iron oxides (Dixit and Hering, 2003), and Ni to manganese oxides
(Larsen and Postma, 1997), the reduction of which might release As and Ni to the pore water, concurrently
with respectively Fe2+ and Mn2+. With respect to As, the presence of significant Fe2+ concentrations indicates
As release due to iron oxide reduction. Results of sequential extractions, and oxalate and citrate-dithionite
iron extractions (Supplementary Information) were consistent with As being related to the ‘iron oxides’-
fraction in the loam. With respect to Ni, the presence in the pore water of Mn2+ may suggest Ni to be due to
reduction of Mn oxides. However, the correlation of dissolved Mn2+ and Ca (Section 4.4.2) indicates that
Mn2+ concentrations are controlled by carbonate dissolution and/or precipitation, rather than by reduction of
Mn oxides. This conclusion is supported by the sequential extraction of loam samples (Supplementary
Information), indicating Mn2+ associated to the ‘carbonates’-fraction below the permafrost table. Both As
and Ni are commonly associated with Fe sulfides (Bostick et al., 2004; Larsen and Postma, 1997), but the
reducing conditions throughout the active layer does not support a release of Ni and As by Fe sulfide
oxidation. Rather, Ni and As sequestration into Fe sulfides (Section 5.2.3) is possible. However, in profile B,
a high Ni concentration is observed in the shallow part of the peat layer, concurrently with a relatively high
SO4 concentration. This could be due to frost boils mixing loam into the peat layer, and hence allowing the
oxidation of Ni-bearing Fe sulfides contained in the loam.
The primary source of dissolved Al is expected to be silicate weathering. Near-saturation for
gibbsite was observed close to the peat-to-loam transition (Fig. 6), possibly suggesting some control by
gibbsite equilibrium on Al concentrations. However, once SIgibbsite strongly deviates from 0, rather the
opposite conclusion must be reached. Instead, we presumed cation exchange to be an important control on
dissolved Al concentrations (cf. Kjøller et al., 2004). To elucidate this, the cation exchanger compositions in
equilibrium with the measured concentrations of dissolved ions (Figs. 4 and 5) were calculated by
PHREEQC, the result being shown in Fig. 7. Initial modeling (grey crosses in Fig. 7) underestimated Na+
and K+ exchange. In a second model run the half reaction exchange coefficients for Na+ and K+ were
therefore increased from 0 to 0.3 for Na+ and from 0.7 to 1.4 for K+. This improved the fit for Na+ and K+
while only slightly affecting the exchangeable di- and trivalent cations (black crosses in Fig. 7). The modeled
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exchanger composition is in overall agreement with measured composition, except for exchangeable Al and
Fe2+, which appears to be either significantly over- or underestimated. The discrepancies between measured
and modeled exchangeable Al and Fe2+ indicate a weak control on these solutes by cation exchange. A third
possible control on dissolved Al could be complexation with dissolved organic matter (Lövgren et al., 1987),
which may also control Fe concentrations (Pettersson and Bishop, 1996). This cannot be ruled out, yet no
correlation between concentrations of DOC and Al or Fe was observed (r2<0.04).
The concentrations of Al and As increases markedly across the permafrost table, and also Ni
concentrations are relatively high below the permafrost table. Thawing, in response to global warming, of the
upper part of the frozen zone therefore might release water with elevated concentrations of Al, Ni and As to
the active layer.
5.3.1 Diffusive trace element efflux
As an evaluation of the conceptual model in Fig. 8, we may calculate an upwards diffusive flux of a trace
element diffusing from the active layer loam to the peat layer. Afterwards we can compare the calculated
value with the observed concentration in the peat layer. The diffusive flux can be calculated using Fick’s
Law:
Fi = -Df ε2 dmi/dz Eq. (2)
where Fi is the flux (moles/s/m2) of solute i, mi is the concentration of i, Df is the free ion diffusion
coefficient, ε is the porosity and z is the elevation (m). In Eq. (2), the term ε2 corrects the free diffusion
coefficient for effects of tortuosity and porosity (Appelo and Postma, 2005). For the loam a general free ion
diffusion coefficient of 0.65·10-9 m2/s at 2 °C and a porosity of 0.47 is applied. For the example of Ni in
profile B, the concentration of Ni drops from 1.8 µM in 35 cm depth to 0.2 µM at 20 cm depth, shortly above
the loam-peat transition. This translates to a concentration gradient of ca. 11 µM/m. Assuming that diffusive
transport occurs only during a four-months summer, yields a maximum upwards diffusive flux of 15 µmoles
Ni/(m2·yr). Solutes arriving to the peat layer will be dispersed in all the pore water stored in the peat, which
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amounts to ca. 100 L/m2, applying the average saturated depth of the peat of 0.12 m and a porosity of 0.8.
For Ni in profile B, the diffusive flux therefore results in a concentration input of 0.15 µmoles/(L·yr). The
mean residence time of peat pore water is ca. 50 years, as calculated from the annual travel distance of 8.3 m
and the 400 meters distance to the lake. Multiplication with the diffusive flux yields a maximum Ni
concentration in the peat layer of 7.5 µM. This number is in reasonable agreement with the values observed
in the peat, up to 1 µM, especially when considering the simplifying negligence of retardation, and of
interaction between peat pore water and channels in channels intersecting the transect, and the use of the
highest measured vertical Ni gradient in the loam. If the constrictions imposed by these simplifications were
relaxed, this in all cases would reduce the concentration buildup in the peat layer.
6. Conclusions
Addressing inorganic carbon cycling, and carbon emission pathways, we investigated the hydrological and
geochemical controls on the pore water chemistry of a permafrost wetland, Ilulissat, Greenland. The wetland
comprises a geological sequence of loam overlain by sphagnum peat, and is saturated to near the surface.
The following conclusions were made:
i. A conceptual model for water and solute transport was derived. According to the model, mineral
weathering reactions take place in the active layer loam, which is only slowly flushed by advective
flow. The reaction products then become transported upwards by diffusion to the frequently flushed
peat layer. Neither the vertical distribution of Cl, nor that of pore water δ18O, appeared to be related
to ion freeze-out or isotope fractionation, respectively.
ii. Pore water in the active layer and water of melted core sections of permafrost were of Ca-Mg-HCO3
type, and were subsaturated for calcite and dolomite. The mutual behavior of 𝑃CO2, aqueous
carbonate species, Ca and Mg (and Mn2+) indicates cycling of inorganic carbon, Ca, Mg, (and Mn2+)
by cryogenic carbonate precipitation during fall freeze-up and re-dissolution after spring thaw.
iii. Elevated Fe concentrations, and subsaturation to near-saturation for siderite, suggests cycling for
Fe2+ via cryogenic siderite formation and re-dissolution.
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iv. High partial pressures of CO2 in the active layer pore water were observed, indicating that ebullition
from saturated active layers may be an important greenhouse gas emission pathway from permafrost
wetlands.
v. Cation exchange equilibria were observed for major cations, as modelled by PHREEQC.
vi. The distribution of trace elements were investigated, to serve a baseline for assessing future impacts
of increased anthropogenic activities in the Arctic. Elevated concentrations of geogenic Al, Ni, and
As were observed. Neither cation exchange nor complexation with DOC appears to be dominant
controls on the Al concentrations. Iron sulfides oxidation are not generally a likely source of As and
Ni, due to prevailing reducing conditions. As release appears linked to iron oxide reduction.
vii. A low efflux of released trace elements to a downstream lake in the permafrost catchment was
found, as quantified using the conceptual model.
7. Acknowledgements
We thank Bo Elberling for sharing experience on permafrost ice core collection, Dieke Postma for scientific
discussions, and two anonymous reviewers for their constructive criticism. The research was funded by the
Greenlandic Ministry of Domestic Affairs, Nature and Environment and the Danish Environmental
Protection Agency.
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Revised manuscript for submission to Water Resources Research
28
Figure captions
Fig. 1: Situation sketch showing the position of the three profiles A, B and C along the investigated NNW-
SSE stretched transect over the wetland. Lake 39 is part of an interconnected lake system used for drinking
water supply. The active layer lithology and depth was investigated with hand drilling equipment for every
10 m along the transect (white dots). Amphibolite and gneiss outcrops border the wetland to the south
(indicated) and north (not within the map).
Fig. 2: Cross section along the transect (Fig. 1) showing i) ground surface and water table position, ii) active
layer depth and its peat and loam distribution, and iii) the position and depth of the three profiles A, B and C.
Fig. 3: Pore water δ18O vs. depth in profile A, B and C. The model line (dotted) is calculated for ice-water
fractionation assuming a downwards movement of the freezing front (ice-water transition) during fall freeze-
up and a reduced fractionation factor αi-w of 1.0005.
Fig. 4: Major chemistry of the pore water in the up-gradient profile A (squares), central profile B (diamonds)
and down-gradient profile C (triangles).
Fig. 5: Pore water concentrations of trace elements Al, As, Ni, Mn2+, and of DOC in the three profiles A
(squares), B (diamonds) and C (triangles).
Fig. 6: Saturation indices (SIs) and partial pressure of CO2, 𝑃CO2, in the three profiles A (squares), B
(diamonds) and C (triangles). Grey area indicates the calculated pH uncertainty field for pH values plus and
minus 0.3 unit away from the measured pH of Fig. 4.
Fig. 7: Composition of the cation exchanger (meq/100 g). Blue dots indicate measured values. Grey and
black crosses indicate values modeled by PHREEQC for exchange equilibrium with the measured pore water
composition of Figs. 4 and 5, using, respectively, the exchange coefficients of the default wateq4f.dat
database (grey) and a modified set of coefficients (black).
Revised manuscript for submission to Water Resources Research
29
Fig. 8: Conceptual model for water and solute movement in the active layer, indicating i) vertical diffusive
solute fluxes within the loam and across the peat-loam interface and ii) dominant lateral advective solute
transport in the peat. Indicated advective velocities are particle velocities. Conservative solutes (e.g. Cl) are
at equilibrium (no net vertical diffusive flux). Non-conservative solutes (e.g. Ca and SO4) released or
consumed within the loam show positive or negative net diffusive flux.
Revised manuscript for submission to Water Resources Research
30
Supplementary Information
1. Field photos showing i) iron staining on rocks along the edge of the wetland and ii) ice lenses in
permafrost cores: ice-rich permafrost.
2. Isotope δ18O vs. δD plot including GNIP data. Support for evaporative enrichment.
3. Sequential extractions showing As related to Fe oxides and Mn related to carbonates. Extractions
indicating Feo/Fed, to characterize the Fe oxides. Methods description for these extractions.
200 m
Ilulissatcity
Lake39
A
B
C
Amphiboliteoutcrop
Gneissoutcrop
Channel
69°1
2'6
0"N
51° 4'15"W
Ilulissat
N
Fig. 1: Situation sketch showing the position of the three profiles A, B and C along the investigated NNW-SSE stretched transect over the wetland. Lake 39 is part of an interconnected lake system. The activelayer lithology and depth was investigated with hand drilling equipment for every 10 m along the transect (white dots). Amphibolite and gneiss outcrops border the wetland to the south (indicated) and north (not within the map).
Intended figure width: 80 mm (1 column).
Lake 39
Gneiss
100 m
1 m
Peat
Loam
Permafrost
A
CB
SSE NNW
Fig. 2: Cross section along the transect (Fig. 1) showing i) ground surface and water table position, ii) active layer depth and its peat and loam distribution, and iii) the position and depth of the three profiles A, B and C.
Intended figure width: 80 mm (1 column).
-18 -16.5 -15
18δ O ( )‰
-13.5
Depth
(cm
)
0
50
100
150
Legend:
Profile A
Profile B
Profile C
Freeze-out model
peat
loam
permafrost tablepermafrost tablepermafrost table
Fig. 3: Pore water δ18O vs. depth in the three profiles. The model line (dotted) is calculated for ice-water fractionation assuming a downwards movement of the freezing front (ice-water transition) during fall freeze-up and a reduced fractionation factor αi-w of 1.0005.
Intended figure width: 70 mm (1 column).
0 0.1 0.2 0.3 0.4
K (mM)
0
50
100
150
0 250 500 750
EC (µS/cm)
0 0.5 1.5
Ca (mM)
0
Mg (mM)
0 0.3 0.6 0.9 1.2
Na (mM)
4 5 6 7 8
pH
0
50
100
150
0 2.5 5 10
Alkalinity (meq/L)
0 0.4 0.8 1.6
Cl (mM)
0 0.25 0.5 0.75 1
Fe (mM)
0 25 50 75 100
SO (µM)4
Depth
(cm
)
1 2 2.5 0.5 1.51 2 2.5
Depth
(cm
)
7.5 1.2
Legend:
Profile A
Profile B
Profile C
permafrost tablepermafrost tablepermafrost table
peat
loam
Fig. 4: Major chemistry of the pore water in the up-gradient profile A (squares), central profile B (diamonds) and down-gradient profile C (triangles).
Intended figure width: 170 mm (2 columns).
0 15 30 45 60
DOC (mM)
0
50
100
150
0 0.2 0.4 0.6
Al (mM)
0 5 10 15
Mn (µM)
0 0.5 1 1.5 2
Ni (µM)
0 0.1 0.2 0.3 0.4
As (µM)
permafrost table
peat
loam
Fig. 5: Pore water concentrations of trace elements Al, As, Ni, Mn2+, and of DOC in the three profiles A (squares), B (diamonds) and C (triangles).
Intended figure width: 170 mm (2 columns).
Depth
(cm
)
0
25
50
75
100
125
150
0 2 4
P_CO2 (atm)
0
25
50
75
100
125
150
0 1 2
P_CO2 (atm)
0
25
50
75
100
125
150
-5 -4 -3 -2 -1 0 1
SI Calcite
0
25
50
75
100
125
150
-5 -4 -3 -2 -1 0 1
SI Calcite
0
25
50
75
100
125
150
-6 -4 -2 0 2 4 6
SI Gibbsite
0
25
50
75
100
125
150
-6 -4 -2 0 2 4 6
SI Gibbsite
0
25
50
75
100
125
150
-5 -4 -3 -2 -1 0 1
SI Siderite
0
25
50
75
100
125
150
-5 -4 -3 -2 -1 0 1
SI Siderite
0
25
50
75
100
125
150
0 0.2 0.4 0.6
SI Chalcedony
0
25
50
75
100
125
150
0 0.2 0.4 0.6
SI Chalcedony
0
25
50
75
100
125
150
-5 -4 -3 -2 -1 0 1
SI Calcite
0
25
50
75
100
125
150
-6 -4 -2 0 2 4 6
SI Gibbsite
0
25
50
75
100
125
150
-5 -4 -3 -2 -1 0 1
SI Siderite
0
25
50
75
100
125
150
0 0.2 0.4 0.6
SI Chalcedony
0
25
50
75
100
125
150
0 0.5 1 1.5 2
P_CO2 (atm)
0 1 2 3 4
P (atm)CO2
0
50
100
150
-5 -3 -1 1
SI Calcite
-6 -2 2
-5
0 0.2 0.4 0.6
SI Chalcedony
0 4 6 -4 -2 0 1
Depth
(cm
)
-4 -2 0
SI SideriteSI Gibbsite
-4 -3 -1
permafrost tablepermafrost tablepermafrost table
peat
loam
Fig. 6: Saturation indices (SIs) and partial pressure of CO2, , in the three profiles A (squares), B (diamonds) and
C (triangles). Grey area indicates the calculated pH uncertainty field for pH values plus and minus 0.3 pH unit away from the measured pH of Fig. 4.
Intended figure width: 170 mm (2 columns).
0
50
100
150
0 10
CaX2
Depth
(cm
)
5 15 0 52.5 7.5 0 0.20.1 0.3 0 10.5 1.5 0 31 52 4 0 31 2
MgX2 NaX KX AlX3 FeX2
permafrostpermafrost table tablepermafrost table
peat
loam
Fig. 7: Composition of the cation exchanger (meq/100 g). Blue dots indicate measured values. Grey and black crosses indicate values modeled by PHREEQC for exchange equilibrium with the measured pore water composition of Figs. 4 and 5, using, respectively, the exchange coefficients of the default wateq4f.dat database (grey) and a modified set of coefficients (black).
Intended figure width: 170 mm (2 columns).
peat
loam
-Cl-Cl-Cl CationsCationsCations
Diffusive transportof reactive compounds
v = 8.3 m/yradv
v = 0.13 m/yradv
permafrost tablepermafrost tablepermafrost table
Fig. 8: Conceptual model for water and solute transport in the active layer, indicating i) vertical diffusive solute fluxes with in the loam and across the peat-loam interface and ii) dominant lateral advective solute transport in the peat. Indicated advectivevelocities are particle velocities.Conservative solutes (e.g. Cl) are at equilibrium (no net vertical diffusive flux). Non-conservative solutes (e.g. Ca and SO4) released or consumed within the loam show up- or downwards net diffusive flux.
Intended figure width: 80 mm (1 column).