Hydrology and pore water chemistry in a permafrost wetland, Ilulissat, Greenland

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

Revised manuscript for submission to Water Resources Research

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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


15

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)


16

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


17

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


18

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


19

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.


20

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.


21

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


22

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


23

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.


24

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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Zimov S.A., Schuur E.A.G. and Chapin F.S. (2006) Permafrost and the global carbon budget. Science 312, 1612–1613.


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).


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.


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.


-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.


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


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).


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


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.


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).


peat

loam

-Cl-Cl-Cl CationsCationsCations

Diffusive transportof reactive compounds

v = 8.3 m/yradv

v = 0.13 m/yradv


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.


Documents

Hydrology and pore water chemistry in a permafrost wetland, Ilulissat, Greenland