BGBiogeosciencesBGBiogeosciences1726-4189Copernicus PublicationsGöttingen, Germany10.5194/bg-13-4315-2016Decadal and long-term boreal soil carbon and nitrogen sequestration rates
across a variety of ecosystemsManiesKristen L.HardenJennifer W.FullerChristopher C.TuretskyMerritt R.US Geological Survey, Menlo Park, CA, USAUniversity of Guelph, Department of Integrative Biology, Guelph, Ontario, CanadaK. L. Manies (kmanies@usgs.gov)1August201613154315432722January201627January201620May20167July2016This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://bg.copernicus.org/articles/13/4315/2016/bg-13-4315-2016.htmlThe full text article is available as a PDF file from https://bg.copernicus.org/articles/13/4315/2016/bg-13-4315-2016.pdf
Boreal soils play a critical role in the global carbon (C) cycle; therefore,
it is important to understand the mechanisms that control soil C
accumulation and loss for this region. Examining C & nitrogen (N)
accumulation rates over decades to centuries may provide additional
understanding of the dominant mechanisms for their storage, which can be
masked by seasonal and interannual variability when investigated over the
short term. We examined longer-term accumulation rates, using 210Pb and
14C to date soil layers, for a wide variety of boreal ecosystems: a
black spruce forest, a shrub ecosystem, a tussock grass ecosystem, a sedge-dominated ecosystem, and a rich fen. All ecosystems had similar decadal C
accumulation rates, averaging 84 ± 42 gC m-2 yr-1. Long-term
(century) C accumulation rates were slower than decadal rates, averaging 14 ± 5 gC m-2 yr-1 for all ecosystems except the rich fen, for
which the long-term C accumulation rates was more similar to decadal rates
(44 ± 5 and 76 ± 9 gC m-2 yr-1,
respectively). The rich fen also had the highest long-term N accumulation
rates (2.7 gN m-2 yr-1). The lowest N accumulation rate, on both a
decadal and long-term basis, was found in the black spruce forest (0.2 and
1.4 gN m-2 yr-1, respectively). Our results suggest that the
controls on long-term C and N cycling at the rich fen is fundamentally
different from the other ecosystems, likely due to differences in the
predominant drivers of nutrient cycling (oxygen availability, for C) and
reduced amounts of disturbance by fire (for C and N). This result implies
that most shifts in ecosystem vegetation across the boreal region, driven by
either climate or succession, will not significantly impact regional C or N
dynamics over years to decades. However, ecosystem transitions to or from a
rich fen will promote significant shifts in soil C and N storage.
Introduction
High-latitude soils store 50 % of the global soil carbon (C) pool
(Tarnocai et al., 2009; Davidson and Janssens, 2006)
due largely to physical factors such as low soil temperatures and wet soil
conditions. As a result, C losses are generally smaller than C inputs, even
over long timescales (Ovenden, 1990) when disturbances such as insects
and fire are included. The majority of net C storage is in the form of thick
(many centimeters to several meters deep) organic soils overlying the mineral soil
component. Climate change is expected to impact the boreal region in many
ways, including thawing of permafrost and reduced precipitation (Hinzman
et al., 2005). These and other changes can alter the dominant vegetation
types. If the factors that moderate C storage shift, it is likely that the
balance of C inputs and losses will also change, impacting the net C
balance. Because of the high amount of C stored in boreal soils, changes in
these C stocks can substantially affect the global C budget
(Chapin et al., 2000).
Many studies have examined boreal N availability, mineralization rates, and
their influence on C storage (for example, Keller et al., 2006; Bonan and
Van Cleve, 1991; Gundale et al., 2014; Allison et al., 2010), yet boreal
nitrogen (N) stocks are less well studied. It is known that boreal forests
have large stocks of soil organic N (Valentine et al., 2006),
with peatland stocks comprising approximately 10–15 % of the global N
pool (Loisel et al., 2014). The majority of N within boreal ecosystems
resides within the organic and mineral soil (Merila et al., 2014).
The size of these soil stocks changes with soil drainage and dominant
vegetation (Van Cleve et al., 1983), in part as a result of N
loss from fire (Harden et al., 2002). Understanding N stocks and
availability is important because N controls many aspects of plant
productivity and, therefore, cycling of C and N are closely linked
(Vile et al., 2014).
Accumulation rates of C and N in soils vary according to the timescale,
ecosystem type, and region studied. Short-term accumulation rates are higher
than long-term rates because there is little influence of disturbance
(Turunen et al., 2004). Accumulation rates also vary by
ecosystem. Peatlands accumulate about 20–30 gC m-2 yr-1 over the
long-term (see, for example, Yu et al., 2013; Jones and Yu, 2010; Turunen
et al., 2002; Roulet, 2000), with bogs typically having higher rates of C
accumulation than fens (Tolonen and Turunen, 1996). Long-term C
accumulation rates in peatlands are driven by growing season length and
photosynthetically active radiation (PAR: Charman et al., 2013). Black
spruce (Picea mariana) forests have C accumulation rates ranging between 10 and 40 gC m-2 yr- (Harden et al., 2000, 2012;
Trumbore and Harden, 1997;
Goulden et al., 2011; Rapalee et al., 1998), depending
on soil drainage class and timescale studied. C accumulation rates of these
ecosystems are related to fire and burial of C into deeper soil layers
(O'Donnell et al., 2011; Harden et al., 2012).
Accumulation rates of N in northern peatlands have been found to average
from 0.5 to 0.9 g N m-2 yr-1 (Loisel et al., 2014; Wang et al.,
2014), although this rate has changed over time. Fens have higher rates of N
accumulation than bogs (Wang et al., 2014; Trumbore et al., 1999),
reflecting the importance of plant and groundwater inputs to fen ecosystems.
Little is known about C or N accumulation rates for the ecosystems other
than peatlands or forests (i.e., shrubs, grass tussock, etc.) that
characterize boreal landscapes. Although these other ecosystem types cover
less area than black spruce forests and peatlands (DeWilde and
Chapin, 2006), they still comprise an important part of the Interior Alaskan
landscape.
Differences in the vegetation and environmental conditions among the varied
ecosystems of Alaska influence their C & N accumulation rates. Litter
production varies among vegetation (Camill et al., 2001),
thereby impacting rates of C and N input to the soil. The chemical content
and concentration of litter also varies among vegetation types
(Hobbie, 1996). Litter composed of more complex C compounds
and/or higher lignin : N ratios can have lower decomposition rates and,
therefore, lower rates of C loss and relative N retention. Vegetation is
also correlated with soil drainage (e.g., soil moisture;
Camill, 1999), the presence of permafrost, and thus the thickness of
insulating organic soil layers (Lawrence and Slater,
2008; Harden et al., 2000). All of these factors affect rates of
decomposition (Rapalee et al., 1998; Wickland et al., 2010; Dioumaeva et
al., 2002; Wickland and Neff, 2008; Treat et al., 2014), losses due to
combustion (Harden et al., 2002), and rates of mineralization
(Bonan and Van Cleve, 1991; Valentine, 2006), with
wetter sites having lower rates of C and N loss. Because litter inputs,
litter quality, the presence of permafrost, and soil moisture and
temperature all affect rates of C and N accumulation and vary among
ecosystem types, it follows that rates of C and N accumulation also vary
according to ecosystem type, with ecosystems with more labile litter and/or
with warmer soil temperatures storing less C and N over the long term.
Many accumulation studies focus on daily, seasonal, or annual timescales.
These studies use either chamber or eddy covariance techniques to measure
net ecosystem exchange (NEE). These short-term investigations have led to
insights regarding the importance of water table to the net C and N budget
(Chivers et al., 2009; Ise et al., 2008), the role of shallow soil layers
in trace gas emissions (Wickland et al., 2010), and the
importance of seasonal variations to the annual net C balance for various
boreal ecosystems (Euskirchen et al., 2014). Additional
insights into the drivers of C and N storage can be obtained by examining
accumulation rates over longer time frames, such as decades or centuries.
Through such investigations we have learned how C accumulation rates
increase as soil moisture increases (Rapalee et al., 1998), how
N deposition increases C accumulation rates (Turunen et al.,
2004), and how disturbances, such as fire (Pitkanen et al.,
1999), reduce C and N accumulation rates.
To help our understanding of longer term C and N accumulation rates in a
variety of boreal ecosystems, we compared soil-based C and N accumulation
rates in five different ecosystems within Interior Alaska, each varying in
soil moisture and dominant vegetation (black spruce, shrubs, tussock grass,
sedge, or moss). These ecosystems were located along a moisture gradient,
thereby controlling for factors such as parent material, climate, and
topography, which influence soil formation (Jenny, 1941). We examined C
and N accumulation rates on both decadal and century timescales to determine
how the interaction of soil and vegetation influences these rates, and thus,
C and N storage over time. Based on differences in soil temperature, soil
moisture, and litter quality, we predicted that the black spruce ecosystem
would have the lowest rate of C and N accumulation while the rich fen would
have the highest rate of C and N accumulation, with the values of the other
ecosystem's accumulation rates residing somewhere in between.
Site biological, physical, and chemical information. Depth
of organic soil, based on three soil cores, are averages with standard
deviations. July temperatures are averaged for 2005–2011. Water table depth
from measurements after July 15 for the years 2005–2008.
Black spruceShrubTussock grassSedgeRich fenDominant vegetationLedum groendlandicum,Salix spp., Betula spp.,Calamagrostis canadensis,CarexDrepanocladus spp.,Vaccinium caespitosum,Chamaedaphne calyculata,Drepanocladus spp.atherodesSphagnum spp.,FeathermossCalamagrostis canadensisCarexatherodesDepth of organic soil (cm)21 ± 230 ± 1529 ± 2216 ± 191 ± 12Shallow permafrost (< 1 m)yesyesnononoAvg. July temperature at 10 cm (∘C)a8.3 ± 2.35.7 ± 0.95.7 ± 2.39.1 ± 3.015.8 ± 5.2Avg. July temperature at 25 cm (∘C)a2.0 ± 0.53.6 ± 0.75.1 ± 2.47.9 ± 3.111.2 ± 1.7Avg. annual temperature at 25 cm (∘C)b-0.03 ± 1.5-1.5 ± 4.0-1.3 ± 5.4-0.03 ± 5.32.1 ± 5.0Water table depth (cm)34 ± 612 ± 715 ± 1311 ± 125 ± 11Soil moisture (% VMC at 5 cm)b15 ± 357 ± 866 ± 872 ± 784 ± 2
a 2005–2011, b McConnell et al. (2013).
Methods
Study sites were located within the Bonanza Creek Long-Term Ecological
Research (LTER) site (64.70∘ N, 148.31∘ W),
approximately 30 km south-west of Fairbanks, Alaska, within the floodplain
of the Tanana River. This region of Interior Alaska is characterized by a
mean annual temperature of -7 ∘C and mean annual precipitation of
300 mm (Hinzman et al., 2006). We studied soils in five
ecosystems located along a ∼ 300 m transect, each of which
was dominated by a different type of vegetation. The ecosystems, presented
in order as they appear on the landscape, are (1) a closed-canopy black
spruce forest with a feathermoss and Ericaceous shrub understory (hereafter “black
spruce”), (2) a shrub system comprised of willow (Salix sp.) and birch (Betula sp.)
with an understory dominated by Chamaedaphnecalyculata and sparse moss cover (“shrub”), (3) a
tussock grass system dominated by Calamagrostiscanadensis with some brown mosses present
(“tussock grass”), (4) a peatland dominated by emergent vegetation such as
Equisetumfluviatile (“sedge”), and (5) a moss-dominated rich fen, comprised of both brown
mosses and Sphagnum sp. (“rich fen”). These ecosystems varied in moisture status
related to water table and presence of permafrost
(Table 1; Waldrop et al., 2012). This transect
extends from the toe slope of an adjacent upland forest into a
∼ 1.8 km2 fen complex. Although in the Tanana floodplain,
the sites are ∼ 1.5 km from the current location of the river
and appear to be relatively stable since site initiation in 2005. These
sites have also been a part of other studies, including examining controls
on ecosystem respiration (McConnell et al., 2013), examining
differences in the soil biotic community and their impact on soil C turnover
(Waldrop et al., 2012), understanding how
changing water table level impacts C cycling within the fen (Kane et al.,
2013; Chivers et al., 2009), and using eddy covariance methods to calculate
net ecosystem productivity (Euskirchen et al., 2014).
Three soil cores, encompassing all of the organic soil and extending into
the mineral soil below, were collected at each site at randomly selected
locations within an area less than ∼ 10 m2. Sampling for
the black spruce and low shrub site occurred during the summer and samples
were obtained using a combination of soil blocks cut to a known volume and
using a “Makita” coring device (4.8 cm diameter; Nalder and Wein,
1998). Soil cores from the other three sites were obtained in the spring,
when the ground was frozen, using a SIPRE corer (7.6 cm diameter;
Rand and Mellor, 1985). Each soil profile was then divided into subsamples
representing soil horizons. Soil horizon thicknesses ranged between 2 and 14 cm,
with 85 % of samples having a thickness ≤ 5 cm. This separation
occurred either in the field or, if frozen, in the lab, based on visual
factors such as level of decomposition and root abundance. Each horizon
sample was described using modified soil survey techniques
(Manies et al., 2016).
Soil horizon samples were processed in several steps: first they were air
dried (20–25 ∘C) and then homogenized. The samples were then
split into two parts: an archive split and an analytical split. The
analytical split was oven dried and then ground. Soils classified as
organics were oven dried at 65 ∘C and ground to < 0.25 mm
using a Cyclone mill (Udy Corporation., Ft. Collins, Colorado). Mineral
soils were oven dried at 105 ∘C and ground using a mortar and
pestle until the soil passed through a 60 mesh (0.25 mm) screen. Total C and
N content was analyzed using a Carlo Erba 1500 Series 2 elemental analyzer
(Fisons Instruments; Manies et al., 2016). C and N stock
inventories were calculated as the total amount of C or N within the profile
to the organic–mineral soil boundary. Recent ages were determined by
measuring 210Pb and 226Ra activities using gamma spectrometry by
means of a Princeton Gamma HPGe germanium well detector using previously
described methods (Van Metre and Fuller, 2009; Fuller et al., 1999).
Total 210Pb activity was measured and is the combination of supported
210Pb (produced in situ through the decay of 226Ra in the soil)
and unsupported 210Pb (produced in the atmosphere and added to the
ecosystem through atmospheric deposition). Unsupported 210Pb was
defined as the difference between measured total 210Pb and 226Ra.
Subsamples from each soil horizon within the profile, starting at the
surface, were measured until unsupported 210Pb was not detected.
Unsupported 210Pb values were used to calculate dry mass accumulation
rates (MAR, g cm-2 yr-1) for each soil horizon, from which dates
of formation were calculated using both the constant flux–constant
sedimentation method (CF:CS; Robbins, 1978) and constant rate of
supply method (CRS; Appleby and Oldfield, 1978). To account for
compaction and loss of mass due to organic matter decomposition, both
methods modeled unsupported 210Pb as a function of cumulative dry mass
(g cm-2), not depth (Appleby and Oldfield, 1992). Cumulative
dry mass is the product of bulk density of the horizon (g cm-3) and the
horizon thickness (cm). The CF:CS method is based on fitting the decrease in
unsupported 210Pb vs. cumulative dry mass to a single exponential
function based on decay, and thus, estimating an overall MAR by assuming a
constant MAR through time. The CRS method assumes a constant rate of input
of unsupported 210Pb activity per unit area and determines a mass
accumulation rate for each soil horizon sampled by mass balance using the
integrated unsupported activity of the whole profile and, thus, accounts for
changes in MAR over time. The age of each sample interval is calculated from
the resulting MAR from the surface downward. Uncertainty of the CRS MAR and
resulting ages are derived from counting error, propagated from the top of
the core downward (Binford, 1990; Van Metre and Fuller, 2009). As the
soil profiles become deeper, and thus older, the total 210Pb activity
approaches the supported activity, with the difference (unsupported
activity) becoming similar to or less than the uncertainty in the
measurement (which is propagated from the top of the core downward; Binford, 1990; Van Metre and Fuller, 2009). At some point the magnitude
of these errors becomes larger than the age estimated for that horizon (for
example, the estimated age of the 19–22 cm horizon of BZBS 4 was 143 years old
with an estimated error of 144 years; Table S1 in the Supplement). This tends to occur for
horizons dated older than 1920. To minimize these errors we constrained our
decadal C accumulation rates to only include organic soil that had formed
within the 6 decades previous to our sampling. Decadal C accumulation
rates were calculated as the cumulative mass of C from the moss surface for
the base of the that soil horizon, divided by the age of this soil horizon
using the CRS age, which is more appropriate for ecosystems with variable
rates of accumulation (Appleby and Oldfield, 1978;
MacKenzie et al., 2011).
We also dated macrofossils, obtained from several processed, and therefore
homogenized, soil horizons, using AMS radiocarbon measurements for
comparison to 210Pb ages (Table S2). Additionally, bulk soil
samples, with roots removed, were submitted from the basal organic soil
horizon to determine the timing of basal organic soil horizon formation.
These samples were submitted to the USGS extraction laboratory (Reston, VA)
for complete combustion and trapping of CO2. Targets were prepared and
submitted for accelerator mass spectrometry at Lawrence Livermore National
Laboratory. Resulting 14C data were corrected for 13C and then
calibrated using CALIB v 7.0 (intercal13; Reimer et al., 2013), or, if
they dated post-1950, CALIBomb (intercal13, NHZ1 curve extension). Long-term
C accumulation rates were calculated as the amount of C within the organic
soil profile divided by the 14C age of that ecosystem. Ecosystem age
was calculated as the average of the minimum and maximum 14C calibrated
ages (Table S2).
Site C (g m-2), N (g m-2), and Unsupported
210Pb (dpm cm-2) storage data. Unsupported 210Pb inventories
represent the total atmospheric input of 210Pb to that site. Data are
averages of three cores with standard deviations. Different letters after
values indicate that the values among ecosystems are statistically different
based on the Tukey Honest Significant Difference test.
Black spruceShrubTussock grassSedgeRich fenC storage in organic soil (g m-2)6460a± 94014140a± 485013950a± 91307930a± 193061500b± 7290N storage in organic soil (g m-2)170a± 20700ab± 150940b± 540610a± 1203690c± 190Unsupported 210Pb (dpm cm-2)12.7a± 5.610.7a± 5.314.6a± 2.710.2a± 3.910.2a± 1.3ResultsCarbon, Nitrogen, and 210Pb Inventories
The rich fen site has significantly deeper organic soils than the other four
sites (p < 0.001), resulting in four or more times the amount of C
and N than the other ecosystems (Table 2). Average unsupported 210Pb
inventories (dpm cm-2) for each of the five ecosystem types were
statistically similar (p= 0.62, Table 2), which indicates that atmospheric
input is the same for all ecosystem types and there are no apparent losses
or transport of 210Pb among sites. Whereas all the unsupported
210Pb was found in organic soil in most systems, between 10 and 15 % of
the unsupported 210Pb activity was found in the mineral soil horizons
(2–4 cm thick horizons) for the tussock grass site. Because unsupported
210Pb is deposited on the organic soil surface while bound to
atmospheric aerosols and dust particles (Shotyk et al.,
2015), we did not expect to find it in mineral soil layers. Its presence in
mineral soil suggests that some of 210Pb bearing particles may be
transported downward in the grass ecosystem. The potential downward movement
of unsupported 210Pb would result in higher apparent CRS MAR and thus
younger ages. Therefore, the tussock grass site was not included in the
comparison of decadal accumulation rates.
Comparison of 210Pb and 14C ages for depth
increments where both analyses are available. The material dated for
14C ages was deciduous leaf fragments (shrub samples), seeds (sedge
sample), or moss leaves and seeds (rich fen samples). The 210Pb values
listed include estimated error.
14C dates and dating methodology comparison
14C dating of the basal organic soil layers provided information
regarding the initiation of soil development. This approach shows that the
rich fen is the oldest ecosystem, at approximately 1390 years old (Table S2). Age estimates for the shrub and sedge ecosystems ranged between 700 and
856 years cal BP. Unfortunately, we did not get ages for the black spruce or
tussock grass ecosystem (due to sample size limitations). Therefore, for all
ecosystems except the rich fen we used an initiation age of 780 years (the
median of the two ages listed above). We justify this approach using the
following logic. First, all of the ecosystems appear to be relatively stable
and lay within ∼ 300 m of each other, along an emergent
landform that grades from the rich fen up to the black spruce forest.
Therefore, all ecosystems along this gradient likely formed within several
hundred years of each other. This assumption is supported by the fact that
the sedge ecosystem is only ∼ 100 years older than the shrub
ecosystem. The grass ecosystem also lies between the shrub and sedge
ecosystems along the gradient; therefore, its age of formation is likely
similar to the values measured for these two ecosystems. Although the black
spruce ecosystem lies at the end of the gradient, a sensitivity analysis
demonstrates that a dramatically different initiation age would be needed to
impact our results (Table S3). Therefore, even if 780 years is not accurate
for the black spruce ecosystem, realistic variations in this value (±400 years) would not change the outcome of our analyses.
For samples with both 14C and 210Pb, the ages defined by each
technique were in general agreement (Fig. 1). We expected the 14C dates
to lie somewhere within the 210Pb estimates due to the fact that the
macrofossils were obtained from a homogenized sample comprised of the
material from an entire soil horizon and so could have formed at any time
between when that soil horizon formed (the base) and the top of that
horizon. In two instances the range of dates predicted using 14C was
older than the 210Pb based age estimates (Fig. 1: Shrub, 8.5–12.5 cm;
Rich fen, 5–10 cm). Because the 5–10 cm rich fen 14C date is also older
than the two samples below it (10–15 and 15–20 cm), this 14C date is
likely not accurate. The younger 210Pb date for the 8.5–12.5 cm
Shrub-1 horizon could indicate that there has been some movement of
210Pb within the soil profile, which has been known to occur with this
dating technique (Turetsky et al., 2004). However, the
14C and 210Pb ages for the 4.5–8.5 cm horizon match well, which
we would not expect if downward transport was a significant issue. In
addition, adjusting our analyses to the 14C dates does not change our
results. Therefore, we feel comfortable moving forward using the 210Pb
age values.
Decadal (< 60 years) and long-term (780–1400 years) C
and N accumulation rates (g m-2 yr-1) with their standard
deviations. Accumulation rates were determined by averaging values
calculated for each individual soil profile by ecosystem type. Different
letters indicate significant differences among ecosystems for that
accumulation rate, based on Tukey Honest Significant Difference test.
Black spruceShrubTussock grassSedgeRich fenDecadal C accumulation rate (gC m-2 yr-1)59a± 13127a± 73–73a± 976a± 9Long-term C accumulation rate (gC m-2 yr-1)8a± 118a± 618a± 1210a± 244b± 5Short : long C ratio7.17–7.21.7Decadal N accumulation rate (gN m-2 yr-1)1.4a± 0.23.6ab± 1.7–5.6b± 0.64.6b± 0.5Long-term N accumulation rate (gN m-2 yr-1)0.22a± 0.030.90ab± 0.191.20b± 0.690.79a± 0.162.66c± 0.14Decadal accumulation rates
Decadal C accumulation rates (< 60 years) calculated from 210Pb
CRS MAR were not statistically different among sites (Table 3;
p value = 0.21), although the shrub ecosystem had the highest rate and the
black spruce had the lowest rate. Decadal rates ranged between 50 and 125 gC m-2 yr-1. Variability within each ecosystem type was high
(coefficient of variability: 12–60 %). This variability is likely due to
within-site heterogeneity, such as microtopography, changes in vegetation,
and differences in belowground biomass. Decadal accumulation rates of the
black spruce and rich fen ecosystems were similar to other literature values
(Fig. 2). N decadal accumulation rates ranged from 1.4 to 5.6 gN m-2 yr-1 (Table 3). The black spruce ecosystem had significantly lower
rates of N accumulation than the sedge and the rich fen ecosystems
(p= 0.004). The rich fen rate had higher decadal N accumulation rates
(4.6 g m-2 yr-1) than values found for a Norwegian bog
(0.6–2.1 g m-2 yr-1; Ohlson and Okland, 1998), but
similar to rates found for a variety of fens (3.7–7.1 gN m-2 yr-1; Trumbore et al., 1999).
A comparison of organic soil C accumulation rates for fen
and black spruce systems for this study (open symbols) and values found in
the literature (solid symbols). Rates were calculated over many different
time spans (annual to millennial). Errors (where available) are standard
deviations. Literature values from Aurela et al. (2004), Aurela et
al. (2009), Bauer et al. (2009), Camill et al. (2009), Dunn et al. (2007), Euskirchen et al. (2014), Harden et al. (2012),
Mathijssen et al. (2014), Oksanen (2006),
Turunen et al. (2002), Trumbore and Harden (1997), Yu et
al. (2003) and Yu et al. (2013).
Long-term accumulation rates
Long-term rates of C accumulation ranged from 8 to 44 g C m-2 yr-1
across sites (Table 3). Variability was highest in the grass tussock sites,
which had a coefficient of variability of 65 %, vs. 12–34% for the
other ecosystems. Long-term rates of N accumulation ranged from 0.22 to 2.66 gN m-2 yr-1 (Table 3) with the black spruce ecosystem having the
lowest rate of long-term N accumulation. The shrub, tussock grass, and sedge
ecosystems had similar rates of long-term N accumulation. The rich fen had
significantly higher rates of N accumulation than the other ecosystems. The
long-term N accumulation rate for the rich fen (2.66 gN m-2 yr-1)
is much higher than rates previously found for general peatlands
(∼ 0.5 gN m-2 yr-1; Loisel et al., 2014; Limpens
et al., 2006) and bogs (0.87 gN m-2 yr-1; Wang et
al., 2014).
As expected, long-term C accumulation rates were lower than decadal rates
for all ecosystems (Table 3; Fig. 2). This decline in C accumulation rates
is consistent with trends found in chronosequence studies using gas flux
(Baldocchi, 2008) and C stocks (Harden et al., 2012).
However, the difference between long- and decadal rates in the rich fen was
much smaller, indicating consistently high rates of C accumulation in this
ecosystem (Table 3) and suggesting some mechanism exists for preserving this
C over longer timescales. Long-term C accumulation rates for the rich fen
are especially high compared to the other ecosystems (p < 0.001),
which were statistically similar (Table 3). Our long-term C accumulation
rates for the rich fen are similar to other rates based on changes in C
stock (Fig. 2; Camill et al., 2009; Trumbore and Harden, 1997; Turunen
et al., 2002).
Discussion
The ecosystems studied here have varied historically in their dominant
vegetation, the presence or absence of permafrost, and hydrology. Despite
these differences in ecosystem structure we found no significant differences
in decadal rates of soil C accumulation (Table 3). Therefore, while inputs
and losses of C into and from the soil system may vary across these
ecosystems, the balance between inputs and losses for surface soil layers
has been relatively similar over the past 60 years. McConnell et al. (2013) measured ecosystem respiration (ER) at the same five
ecosystems and found higher ER in the grass and sedge ecosystems
(see also Waldrop et al., 2012), with the other
three ecosystems having similar, lower ER; thus the grass and sedge also
have higher rates of net primary production (NPP) and generally cycle C more
rapidly than the other systems. Across all ecosystem types, the shallow
organic soil layers, which have been created in the past 6 decades,
sequestered an average of 84 ± 42 gC m-2 yr1.
Carbon inputs and losses also balance out similarly over the long-term
(∼ 1000 years) for all of the ecosystems we studied except the
rich fen, which had greater long-term C accumulation rates than the other
ecosystems (44 ± 5 gC m-2 yr-1; Table 3). The similarity in
long-term C accumulation rates of the black spruce, shrub, grass, and sedge
ecosystems (14 ± 5 gC m-2 yr-1) was initially surprising, as
we expected the small, although not statistically significant, differences
in the decadal C accumulation rates to add up over time, resulting in some
significant differences in long-term accumulation. In hindsight, however,
this result makes sense, as the total C stored in the organic soils of these
four ecosystems are similar (Table 2). These results again demonstrate that
even if the magnitude of C fluxes into or from the soils systems vary across
these four sites, the overall balance between C inputs and losses is
similar. We note that these four ecosystems fall along the same ER – soil
temperature relationship (McConnell et al., 2013), suggesting that
soil temperature may be one of the main drivers of C cycling for these
sites.
Nitrogen accumulation rates have been studied much less frequently than
rates of C accumulation. The long-term N accumulation rate for the rich fen
in this study (2.66 g N m-2 yr-1) is five times higher than the
0.5 g N m-2 yr- estimated by Loisel et al. (2014). There are
several potential reasons for this discrepancy. First, Loisel et al. (2014)
synthesized data from a wide range of peatland sites, including bogs, fens,
and permafrost peatlands and thus included ecosystems with a broad spectrum
of peat properties. In addition, Loisel et al. (2014) used time-dependent
C : N ratios of 65 and 40 to assign % N values for their soil horizons,
resulting in average % N values that never exceed 1.7 %. In contrast,
the average % N value for our rich fen organic soil horizons was 2.4
%, resulting in an average C : N ratio of 17 (Fig. S1). In general, our
results support Treat et al. (2015), who showed that fen C : N ratios can be
much lower than estimates used by Loisel et al. (2014), despite high
variability (fen C : N averaging 29 ± 15). Regardless, the amount of N
within the rich fen ecosystem is relatively high. Reasons for this high N
storage could include high rates of N inputs, either through high rates of
biological N2 fixation or through high N concentrations in source
water. The majority of studies on N fixation in peatlands have focused on
Sphagnum species (Larmola et al., 2014; Vile et al., 2014). However, over 70 %
of the ground cover in our rich fen site is composed of brown mosses
(Churchill, 2011), some of which have been shown to fix N when
exposed to enough light (Basilier, 1979). Therefore, moss-based N2
fixation may play a role in the N dynamics of the rich fen. High N inputs
could also result from inflows of N-rich surface or ground water. Wetlands
in the Tanana River floodplain can be influenced by both surface runoff and
river-based groundwater, as evidenced by Ca++ values
(Racine and Walters, 1994). All ecosystems along the
gradient, with the exception of the black spruce forest, have been known to
experience flooding during years of very high precipitation, with these
flooding events dependent on the behavior of the Tanana River. During one
of such events, Wyatt et al. (2011) found that dissolved
inorganic N (DIN) at our rich fen site peaked post-flood at ∼ 0.50 mg L-1. Dissolved organic N (DON) at this site has been measured
from ∼ 0.86 to 1.42 mg L-1 (Kane et
al., 2010). While these DIN and DON concentrations are not uncommon for a
northern peatland (Limpens et al., 2006), the hydrologic
connection between the fen and river is undoubtedly important to the total N
budget of the wetland. In addition, long-term influences such as disturbance
likely play an important role in N cycling (see below for more discussion
regarding the influence of disturbance).
The higher long-term C accumulation rate for the rich fen compared to the
other ecosystems suggests that long-term C cycling is fundamentally
different in the rich fen. The rich fen has significant deeper organic soil
(91 cm vs. 30 cm or less for the other ecosystems). Mechanisms for C
sequestration within this soil could be related to (1) higher inputs into
deep soil, from processes such as rooting, (2) less decomposable substrates,
which in turn reduces C losses, and/or (3) environmental conditions (i.e.,
soil temperature, oxygen availability) that reduce decomposition losses.
First, we examined rooting depth for each of the ecosystems. Descriptions of
the rich fen soil cores (Manies et al., 2016) show that live
roots are found throughout the 90 cm organic soil profile, which is
significantly deeper than the other four ecosystems (Table 1). Therefore,
input of C into the deep soil from roots is one possible mechanism for the
larger amount of long-term C found at the rich fen. Next, we examined the C
chemistry, or “quality”, based on the organic soil C : N
(Schädel et al., 2014). Lower C : N indicates
substrate that has undergone more decomposition and, therefore, would likely
be comprised of more recalcitrant material. A comparison of surface C : N
(< 20 cm) shows that the fen system has lower C : N than the black
spruce or shrub ecosystems, but similar values to the grass and sedge
ecosystems (Fig. S1 in the Supplement). This same pattern holds true for deeper soil layers
(>20 cm; Fig. S1, note that the sedge site does not have
organic soil deeper than 20 cm). More decomposable material could also be
reflected in higher ER rates. However, McConnell et al. (2013)
found that ER at the black spruce, shrub, and rich fen sites were
statistically similar. Therefore, differences in decomposable substrates
likely do not play an important role in supporting deep soil C storage at
the rich fen. Finally, we examined differences in environmental conditions,
such as temperature and oxygen availability between the fen and other sites.
Colder soil temperatures at depth at the rich fen could create slower rates
of C cycling due to thermal protection. However, the rich fen site has
warmer summer and annual soil temperatures at both 10 and 25 cm (Table 1),
and the soil temperature is above freezing even at depth (there is no
shallow permafrost at this site). Therefore, preservation of C by thermal
protection is not likely a contributor to the large amount of organic soil
in this ecosystem. Another mechanism for reducing rates of C cycling is
oxygen availability. McConnell et al. (2013) found lower Q10
values at the rich fen, indicating less temperature sensitivity. Instead,
with the shallowest water table (Table 1), it is thought that oxygen
availability plays a dominant role in the protection of deep C at the rich
fen (McConnell et al., 2013). Using average annual growth rates
for the last 60 years, we found that surface organics at rich fen become
submerged in 2 decades, while it takes the surface material of the other
ecosystems 40–90 years to reach the water table. Therefore, the rich fen
organics are exposed to oxygen-limiting conditions much more quickly than
the other ecosystems.
Long-term C and N accumulation rates are also impacted by long-term factors,
such as disturbance. The main disturbance in the boreal region is fire
(Zoltai et al., 1998; Turetsky et al.,
2011), which impacts the boreal C and N cycles directly through emissions
and indirectly via decreasing albedo (Ueyama et al., 2014),
removing insulating organic soil layers (Pastick et al.,
2014), and decreasing soil moisture (Carrasco et al., 2006),
all of which impact decomposition rates. Because the rich fen has a
shallower water table than the other ecosystems (Table 1), this ecosystem is
less likely to burn (Zoltai et al., 1998; Camill et al., 2009; Harden et
al., 2000), even in dry years, and less severely if it does burn (Camill
et al., 2009; Harden et al., 2000). Therefore, while the other ecosystems
likely experienced many fires over the last several millennia, these fire
events had a much smaller, if any, impact on C and N loss from the rich fen.
An examination of the soil horizon descriptions for these sites only found
one incidence of charcoal being observed (Manies et al., 2016;
shrub ecosystem, 12–17 cm). A macrofossil analysis of these or
future cores could further support this hypothesis.
Decadal and long-term C accumulation rates can be used to constrain C
accumulation rates as measured by eddy covariance flux towers.
Euskirchen et al. (2014) examined annual C accumulation rates
in 2012 and 2013 at this same rich fen location and found C accumulation
rates of 36 and 127 gC m-2 yr-1, respectively. By comparison, our
C accumulation rates ranged from 76 and 44 gC m-2 yr-1 (short- and
long-term rates, respectively). The tower based net ecosystem exchange (NEE)
in 2013 is 1.5 and 3 times higher than the decadal and long-term C
accumulation rates found in this study, respectively, while the 2012 NEE
rate is lower than both rates. As our decadal rates are averaged over the
last 6 decades, this discrepancy suggests that the large C loss values
Euskirchen et al. (2014) found in 2013 cannot be sustained over
decades. Interannual variations in NEE for boreal systems are influenced by
the length of the snow free season, soil temperature, light limitation
(i.e., cloudiness), and changes in water table (Baldocchi, 2008). If
tower measurements were continued over a longer time period we would expect
high variability in annual NEE values and those values to be based on that
year's weather conditions. Based on our decadal C accumulation rates years
of high net C accumulation, like 2013, should be balanced out with years of
net loss or low C accumulation to equal decadal rates from core profiles.
It is important to note that our sites are located close to the Tanana River
and thus our findings may be more indicative of locations where the
groundwater can be influenced by river water. We also found a high level of
within-ecosystem variability, with coefficients of variability of up to
60%. This variability is likely due to microsite variability in surface
vegetation, microtopography, and soil characteristics such as porosity, all
which influence C and N cycling, and thus, accumulation rates. This
variability limited our ability to make inferences about soil C and N
accumulation rates between the four non-fen ecosystems. We also acknowledge
that there is uncertainty associated with both dating techniques used in
this study. Downward transport of 210Pb could make the ages presented
here appear younger than the actual age of the soil horizon. There are also
potential uncertainties with 14C ages due to the movement of younger
atmospheric C into the soil through roots or fungi and the uptake of C from
non-atmospheric sources (Bauer et al., 2009). To minimize these
factors, future researchers could improve upon our methods by increasing the
number of soil cores, having higher resolution for soil horizons, and
studying the possibility of 210Pb downwash using 7Be
(Hansson et al., 2014). Regardless of the high
within-ecosystem variability and potential accuracy of ages, we found
significant differences in the long-term C and N accumulation rates of the
rich fen in comparison to the other four ecosystems studied.
Future changes to Interior Alaska's climate are likely to affect C and N
accumulation rates of the ecosystems studied here differently. Increases in
air temperatures (Hinzman et al., 2005) are likely to increase ER at the
black spruce, shrub, grass and sedge ecosystems, based on findings by
McConnell et al. (2013). This change will, in turn, reduce the
decadal C accumulation rates of these ecosystems. However, climate induced
shifts from vegetation from one ecosystem type to another among the four
similar ecosystems should not impact either short- or long-term C
accumulation rates, as we found similar rates among these four ecosystem
types. Therefore, shifts between these ecosystem types likely should not
impact the regional C budget. This statement assumes, however, that any
changes in climate influence the balance between C inputs and losses equally
among ecosystems. Projected increases in fire severity and frequency
(Turetsky et al., 2011) will also impact C
accumulation rates, especially in the long-term. In contrast, rich fens are
more likely to sustain their C and N accumulation rates as long as water
tables are maintained as this high water table appears to diminish
decomposition and reduce disturbance, thereby helping the rich fen maintain
its C and N stocks. However, the magnitude of the rate can be expected to be
quite variable from year-to-year (Euskirchen et al., 2014; Baldocchi,
2008). The C and N balance of rich fens is likely to be significantly
impacted only if there are dramatic drops in water table
(Waddington et al., 2014), which would require large
changes to both the precipitation regime and subsurface hydrology (i.e.,
input sources of water), thereby increasing the susceptibility of the rich
fen to wildfire and decreasing the zone of anoxic conditions, both of which
are important in maintaining the large C and N stocks of this site.
Conclusions
This study provides C and N accumulation rates for a variety of northern
ecosystems, many of which previously had little or no data available.
Knowing rates of C and N accumulation in these five ecosystems will aid in
the understanding of and ability to model their C & N cycles. For
example, the overall C balance for four of the five ecosystems was similar,
even though inputs and losses are different, despite differences in dominant
vegetation, presence or absence of near-surface permafrost, and depth to
water table. The significantly higher long-term C & N accumulation rates
at the rich fen support the idea that long-term biogeochemical cycling in
this ecosystem is different. We hypothesize that the black spruce, shrub,
tussock grass, and sedge ecosystems experience more wildfires than the rich
fen site, reducing their ability to preserve C and N over the long-term.
Additionally, C cycling in the rich fen ecosystem appears to be driven by
different biogeochemical processes (such as lower oxygen availability) which
results in the annual C balance of the rich fen more likely being a net C
sink, thereby increasing long-term C accumulation rates. Climate change may
increase rates of disturbance and soil temperatures for the non-rich fen
ecosystems, impacting C and N accumulation rates. However, shifts from one
ecosystem type to another among these four ecosystems would not impact
regional C budgets. Our data also suggest that climate change is less likely
to significantly impact C budgets at the rich fen, as large changes in rich
fen C accumulation rates would only occur if there is a dramatic drop in
water table, which would require large changes to both the precipitation
regime and subsurface hydrology.
Data availability
More information regarding these sites, the methods used to describe and process the samples, as well as data files containing the raw data
can be found at Manies et al., 2016.
The Supplement related to this article is available online at doi:10.5194/bg-13-4315-2016-supplement.
Acknowledgements
We thank the Bonanza Creek Long-term Ecological Research program for
granting us access to these research sites. Their personnel, especially
Jamie Hollingsworth, have been instrumental in providing support for this
research. Thank you to Lee Pruett, Renata Mendieta, and Pedro Rodriguez for
assisting with core collection, sample processing, or analyzing samples. We
also thank Claire Treat, M. Braakhekke, F. S. Chapin, and two anonymous reviewers for
providing helpful comments on an earlier version of this manuscript. Funding
for this work was provided by the US Geological Survey Climate Research
and Development program and the National Science Foundation (DEB-0425328).
The Bonanza Creek Long-term Ecological Research program is funded jointly by
NSF (DEB-0620579) and the USDA Forest Service Pacific Northwest Research
Program (PNW01-JV11261952-231).
Edited by: S. Zaehle
Reviewed by: M. Braakhekke, F. S. Chapin, and two anonymous referees
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