Sample collection
Cores were collected near Drew Point from April 10th-19th, 2018 using two coring systems. Near-surface cores (upper 4 to 6 m) were acquired using a SIPRE corer (7.5 cm diameter) and cores at depth were acquired using a JIPRO corer (7.5 cm diameter). To capture variations in near-surface permafrost characteristics, we sampled each of the three dominant geomorphic terrain units present in the Drew Point region: primary surface material that has not been reworked by thermokarst lake formation and drainage, an ancient drained thermokarst lake basin (DTLB) (Hinkel et al. 2003; Jones et al., 2012), and a young DTLB (Jones et al., 2012). Each permafrost core spanned from the tundra surface to below local mean sea level. Cores were collected in air temperatures between -10 and -20 degrees C. They were packed into coolers for transport back to Utqiagvik, Alaska and then flown frozen to the University of Alaska in Fairbanks where the cores were stored in a -20 degrees C freezer room prior to shipping them frozen to Sandia National Laboratory in Albuquerque, NM for processing.
At Sandia National Laboratory, sections from the frozen cores were cut using a band saw that was cleaned with Milli-Q water and ethanol after each use. Core material was then thawed in acid-washed glass beakers at room temperature for subsampling. Thawed soil/sediment was placed in Whirl-packs and frozen for bulk soil/sediment organic carbon and nitrogen content, stable carbon and radiocarbon analysis, and experimental work.
Core sections for this experimental work were chosen to represent similar depths within three distinct soil/sediment horizons within each core: the seasonally thawed "active layer", Holocene age terrestrial/lacustrine derived permafrost, and permafrost composed of relict, Pleistocene-aged marine sediment. This relict marine permafrost is located above and below modern day sea level in Drew Point bluffs. Additional information about core collection, sampling, and geochemical characteristics can be found in Bristol et al. 2021. Physical and geochemical data from these cores is also published on Environmental Data Initiative (https://doi.org/10.6073/pasta/cc4d53a91ed873765224fcb6d09f5eb7).
Seawater used for the incubation experiment was collected from Beaufort Sea surface waters near Barter Island during August 2018. Seawater was filtered using pre-combusted Whatman GF/F (0.7 um) filters and transported frozen to the University of Texas Marine Science Institute. The seawater had a salinity of 31 and a DOC concentration of 1.2 mg C L
-1.
Incubation experiment
We simulated the decomposition of eroding coastal soils by incubating active layer and permafrost samples from Drew Point in seawater in aerobic conditions at two temperatures (4 and 16 degrees C) in the dark for 40 days. As a comparison, we also incubated a subset of samples within and without seawater at 16 degrees C. To facilitate comparisons with other study areas, we generally followed experimental methods outlined by Tanski et al. (2019, 2021). Briefly, homogenized core sections were thawed at 4 degrees C and triplicate subsamples of 20 g wet weight were placed in precombusted ~120 mL glass serum bottles. Subsamples for wet:dry weight ratios were also taken at this step. For the seawater treatments, 30 mL of filtered seawater was added to each serum bottle. As a control, bottles containing only seawater were also incubated. The pH of the mixtures and seawater controls were measured before sealing the bottles with rubber septa. Then, the headspace of each bottle was flushed using a tank of compressed atmospheric air (480 ppm CO
2). This same tank of atmospheric air was used throughout all experimental work.
Headspace samples were collected at 5, 10, 20, 30, and 40 days for bottles incubated at 4 degrees C, and at 1, 3, 5, 10, 20, 30, and 40 days for bottles incubated at 16 degrees C. Before sampling, each bottle was over-pressurized by adding 12 mL of gas from the tank of compressed atmospheric air using a gas tight syringe. Bottles were gently shaken to allow equilibration. After a few minutes, bottle headspace was sampled by transferring 12 mL of gas to 5.9 mL pre-evacuated Exetainer vials. After headspace sampling, septa were removed to measure the pH of the slurries and seawater controls. Then, bottles were resealed and flushed with atmospheric air using the tank of compressed atmospheric air for 6 minutes with an approximate flushing rate of 900 mL min
-1. Experimental work was conducted at the University of Texas Marine Science Institute. Exetainers vials were promptly shipped via ground transport to Florida State University to measure CO
2 concentrations.
Geochemical analyses
Measurements of total organic carbon (OC) and total nitrogen (TN) content, stable carbon isotope ratios (δ
13C) and radiocarbon (
14C) analyses of bulk soils/sediments were conducted on 45 samples at the Woods Hole Oceanographic Institution (WHOI), National Ocean Sciences Accelerator Mass Spectrometer (NOSAMS) facility. Bulk samples were dried at 60 degrees C then finely ground using a mortar and pestle. Ground samples went through a vapor fumigation acid/base treatment step to remove inorganic carbon. This step involved placing samples in a vacuum-sealed desiccator in a drying oven (60 degrees C) with a beaker of concentrated HCl for 24 hours. Samples were then removed and placed in another vacuum-sealed desiccator with a dish of NaOH pellets, and again stored in a drying oven at 60 degrees C for another 24 hours. This latter step neutralized excess HCl. Samples were combusted using an Elementar el Vario Cube C/N analyzer. TOC and TN (% by weight) were quantified during this step. The resulting CO
2 was transferred to a vacuum line and cryogenically purified. The purified CO
2 gas samples were converted to graphite targets by reducing CO
2 with an iron catalyst under 1 atm H2 at 550 degrees C. Targets were subsequently analyzed for stable and radiocarbon isotopes (δ
13C as ‰ and
14C as fraction modern carbon). All Δ
14C data (in ‰) were corrected for isotopic fractionation using measured δ
13C values that were quantified during the
14C-AMS procedure. We measured δ
13C in these samples separately on a VG Prism Stable Mass Spectrometer at NOSAMS. Δ
14C and radiocarbon age were determined from percent modern carbon using the year of sample analysis according to Stuiver and Polach (1977).
CO
2 concentrations in headspace gas samples were measured at Florida State University. Pressurized exetainer vials were interfaced to a 1 mL gas sampling loop on a Shimadzu 8A gas chromatograph equipped with a methanizer and calibrated against Airgas standards. The GC had a carbosphere column and was operated at 140 degrees C.
CO2 production calculations
Headspace CO
2-C was calculated using measured headspace volumes and CO
2 concentrations, accounting for changes in pressure and concentration when bottles were over pressurized with atmospheric gas before sampling. Headspace pressure within the bottles was assumed to be 1 atm before over pressurizing. Dissolved CO
2 was calculated for seawater treatments using the solubility of CO
2 at the incubation temperature and seawater salinity (Weiss, 1974). Dissolved inorganic carbon (DIC; i.e., dissolved CO
2, bicarbonate, and carbonate) was calculated using the calculated dissolved CO
2 concentration, measured pH, and volume of seawater. When calculating CO
2-C production, we also accounted for the inorganic carbon that remained in the headspace and seawater after flushing, using atmospheric air CO
2 concentrations and pH at each timepoint. DIC calculations were performed using constants from Lueker et al. (2000), Dickson and Riley (1979) and Dickson (1990) using the 'seacarb' package in R v.3.3.0. (Gattuso et al., 2021).
CO
2-C production from decomposition was calculated considering the total increase in inorganic carbon in each vial, including both headspace CO
2-C and seawater DIC. CO
2-C production was normalized to soil/sediment dry weight and OC content. Results from triplicate subsamples were averaged before further analysis. In a few cases where gas samples were lost, an average of two replicates were used instead of three. CO
2-C production rates were calculated using the slope of a linear regression model of cumulative production over time. Additionally, OC loss (%) was calculated by comparing the CO
2-C produced over 40 days to the initial soil/sediment OC content. To describe how CO
2 production rates increase with increasing temperature, the Q
10 temperature coefficient was calculated using the following equation:
$$Q_1_0 = \left ( \frac{R_2}{R_1} \right )^\frac{10}{T_2-T_1}$$
where R2 and R1 represent CO
2 production rates at the respective incubation temperatures, T2 and T1, in degrees Celsius. All results reported here are averages of subsamples that were incubated in triplicate.
References
Bristol E M, Connolly C T, Lorenson T D, Richmond B M, Ilgen A G, Choens R C, Bull D L, Kanevskiy M, Iwahana G, Jones B M and McClelland J W 2021 Geochemistry of Coastal Permafrost and Erosion-Driven Organic Matter Fluxes to the Beaufort Sea Near Drew Point, Alaska Front. Earth Sci. 8 598933
Bristol, E., C. Connolly, T. Lorenson, B. Richmond, A. Ilgen, R. Choens, D. Bull, M. Kanevskiy, G. Iwahana, B. Jones, and J. McClelland. 2020. Geochemical characterization and material properties of coastal permafrost near Drew Point, Alaska ver 1. Environmental Data Initiative. https://doi.org/10.6073/pasta/cc4d53a91ed873765224fcb6d09f5eb7
Dickson, A G, and Riley, J. P. (1979). The estimation of acid dissociation constants in seawater media from potentiometric titrations with strong base. Marine Chemistry, 7, 89-99.
Dickson, Andrew G. (1990). Thermodynamics of the dissociation of boric acid in synthetic seawater from 273.15 to 318.15 K. Deep Sea Research Part A. Oceanographic Research Papers, 37(5), 755-766. https://doi.org/10.1016/0198-0149(90)90004-F
Gattuso, J.-P., Epitalon, J.-M., Lavigne, H., and Orr, J. (2021). seacarb: seawater carbonate chemistry. (Version R package version 3.3.0.). Retrieved from https://CRAN.R-project.org/package=seacarb
Jones, M.C., Grosse, G., Jones, B.M. and Walter Anthony, K. (2012). Peat accumulation in drained thermokarst lake basins in continuous, ice-rich permafrost, northern Seward Peninsula, Alaska. J Geophys Res: Biogeo 117. Doi: 10.1029/2011JG001766.
Hinkel, K.M., Eisner, W.R., Bockheim, J.G., Nelson, F.E., Peterson, K.M. and Dai, X. (2003). Spatial extent, age, and carbon stocks in drained thaw lake basins on the Barrow Peninsula, Alaska. Arctic, Antarctic, and Alpine Research, 35:3, 291-300.
Lueker, T. J., Dickson, A. G., and Keeling, C. D. (2000). Ocean pCO2 calculated from dissolved inorganic carbon, alkalinity, and equations for K1 and K2: validation based on laboratory measurements of CO2 in gas and seawater at equilibrium. Marine Chemistry, 70(1-3), 105-119. https://doi.org/10.1016/S0304-4203(00)00022-0
Stuiver, M. and Polach, H.A., 1977. Discussion: Reporting of 14C data. Radiocarbon, 19:355-363
Tanski G, Broder L, Wagner D, Knoblauch C, Lantuit H, Beer C, Sachs T, Fritz M, Tesi T, Koch B P, Haghipour N, Eglinton T I, Strauss J and Vonk J E 2021 Permafrost Carbon and CO2 Pathways Differ at Contrasting Coastal Erosion Sites in the Canadian Arctic Front. Earth Sci. 9 630493
Tanski G, Wagner D, Knoblauch C, Fritz M, Sachs T and Lantuit H 2019 Rapid CO2 Release From Eroding Permafrost in Seawater Geophys. Res. Lett. 46 11244-52
Weiss, R. F. (1974). Carbon dioxide in water and seawater: the solubility of a non-ideal gas. Marine Chemistry, 2(3), 203-215. https://doi.org/10.1016/0304-4203(74)90015-2
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