Introduction
The future trajectory of the Arctic Boreal Zone (ABZ) as a carbon (C) sink or
source is of global importance due to vast quantities of C in permafrost and
frozen soils (Belshe, Schuur and Bolker, 2013). Cold and waterlogged
conditions in the ABZ have hindered soil organic material (SOM) from
microbial decomposition and led to long-term C accumulation at soil depths
below 1 m (Ping et al., 2015). Arctic warming, which stimulates plant
growth as well as respiration in tundra ecosystems (Mack et al., 2004;
Euskirchen et al., 2012; Natali et al., 2012; Barichivich et al., 2013;
Commane et al., 2017), has driven a period of C cycle intensification over
the last 50 years with greater C inputs and outputs across high-latitude ecosystems (Graven et al., 2013). Expert assessments of site-level
observations, inversion studies, and process models suggest that Arctic
C balance is near neutral, but large uncertainties allow for solutions
ranging from small sources to moderate sinks; however, most assessments favor
an overall strengthening of the regional C sink, with productivity gains
exceeding respiration losses on average (McGuire et al., 2012).
The effect of continued warming on future northern high-latitude (NHL)
ecosystem C balance is uncertain but appears to be increasingly dependent on
responses to changes in cold-season emissions, soil moisture, shifts in
vegetation community, and permafrost degradation (Abbott et al., 2016). These
vulnerabilities are likely driven by disproportionate warming during the cold
season (Fraser et al., 2014), which is projected to increase at twice the
rate of summer warming over the next century (Christensen et al., 2013). For
example, winter warming during the long cold season promotes increased soil
respiration, offsetting C uptake during the short Arctic growing season
(Oechel et al., 2014; Euskirchen et al., 2016; Commane et al., 2017), and
shifting tundra ecosystems from C sink to source (Webb et al., 2016). Winter
warming also promotes earlier and more rapid snow melt and landscape thawing
(Goulden, 1998; Schuur et al., 2015). This can impact seasonal C balance
through increased hydrological export of SOM by Arctic rivers (Olefeldt and
Roulet, 2014), which is projected to increase by 75 % by end of century
(Abbott et al., 2016). Early snow melt can also cause increased exposure of
the land surface to solar absorption (Lawrence et al., 2012), resulting in
increased evapotranspiration and summer drought risk (Zhang et al., 2011),
which decreases terrestrial biomass through reduced plant growth and
increased intensity and frequency of boreal fire emissions and fire
disturbance (Yi et al., 2014; Veravebeke et al., 2017). ABZ fire-driven C losses are
expected to increase 4-fold by 2100 (Abbott et al., 2016).
On longer timescales, permafrost degradation and resulting C losses from
deep, old C is expected to be the dominant factor affecting future Arctic
C balance (McGuire et al., 2012; Lawrence et al., 2015; Schuur et al., 2015).
In addition to these effects, warmer temperatures and longer non-frozen
seasons caused by earlier spring thaw and later autumn freezing can promote
accelerated deepening and increased duration of the active layer (layer of
soil near the surface which is unfrozen in summer and frozen in winter) and
thawing permafrost. More abrupt processes such as thermokarst lake
initialization can also lead to rapid thaw through pronounced sub-lake talik
formation (Jorgenson and Osterkamp, 2005). These processes can initiate
formation of a talik zone (perennially thawed subsurface soils) during
active layer adjustment to new thermal regimes (Jorgenson et al., 2010) in
lake and non-lake environments. Talik as well as longer, deeper active layer
thaw stimulate respiration of soil C (Romanovsky and Osterkamp, 2000;
Lawrence et al., 2008), making the ∼ 1035 Pg soil organic
carbon in near surface permafrost (0–3 m) and ∼ 350 Pg
soil organic carbon in deep permafrost (> 3 m) vulnerable to
decomposition (Hugelius et al., 2014; Jackson et al., 2017).
Climate models used in the Coupled Model Intercomparison Project Phase 5
(CMIP5) consistently project widespread loss of permafrost in the future due
to climate warming (Slater and Lawrence, 2013), though the
earth system models that
participated in the CMIP5 also project NHL terrestrial C uptake rather than
losses due to warming (Ciais et al., 2013). This projection conflicts with
expectations from field studies (Schuur et al., 2009; Natali et al.,
2014),
but newer approaches, such as explicitly representing the vertical structure
of soil respiration and its coupling to deep soil thermal changes, lead to
changes in the model-projected response from a net C gain with warming to
a net loss and hence a positive carbon–climate feedback (Koven
et al., 2011).
Permafrost C emissions are likely to occur gradually over decades to
centuries and therefore are unlikely to cause abrupt and easily detected
signals in the global C cycle or climate (Schuur et al., 2015). We use the
coupled permafrost and biogeochemistry Community Land Model version 4.5
(CLM4.5) to investigate in detail the subsurface thermal processes driving
C emissions from shallow (0–3 m) and deep (> 3 m)
permafrost C stocks and to project the rate of NHL permafrost C feedbacks
(> 55∘ N) over the 21st century. Using CLM4.5 in the framework of
an observing system simulation experiment (Parazoo et al., 2016), we ask how
we might be able to (1) identify potential thresholds in soil thaw,
(2) detect the specific changes in soil thermal regimes that lead to changes
in ecosystem C balance, and (3) project future C sources following talik
onset. We hypothesize that talik formation in permafrost triggers accelerated
respiration of deep soil C and, ultimately, NHL ecosystem transition to
long-term C sources.
Comparison to observed thaw at selected tundra and forested ecosystems along
north–south transects in Siberia and North America in the 20th and early
21st century provides a reference to evaluate historical thaw patterns and
projected thaw rates. The remainder of our paper is organized as follows:
Sect. describes our methods to simulate and analyze soil thaw and
C balance in CLM4.5; Sect. describes borehole datasets used to
analyze CLM4.5 soil thermal regime; Sect. presents results of
talik formation in CLM4.5 and comparison of simulated thaw profiles to
borehole data in North America; Sect. evaluates projected thaw
rates against long-term borehole data in Siberia; Sect. identifies
timing and location of C source onset and discusses formation mechanisms in
the presence and absence of talik; Sect. discusses the main findings.
Decade of projected talik formation and correlation to initial state
of simulated permafrost temperature and observed permafrost extent.
(a) Time series and (b) map of the simulated decade of
talik formation are estimated from CLM4.5 as the first decade when the mean
temperature of a soil layer exceeds a freeze–thaw threshold of
-0.5 ∘C in every month. Additional colors in panel (a)
represent progression of talik onset for different freeze–thaw thresholds.
(c) Initial permafrost temperature is defined as the annual mean
soil temperature at 3 m depth from 2006 to 2010.
(d) Permafrost extent is taken from
https://nsidc.org/data/docs/fgdc/ggd318_map_circumarctic/ (Brown
et al., 2001). Crosses in panel (a, c, d) represent locations of Siberian
borehole measurements along the East Siberian Transect from 1955 to 1900
(Table 1). Circles represent locations of borehole measurements in Alaska and
Canada from 2002 to 2013 (Table 2). Dashed black line in panel (a) shows
projected air temperature over the talik region. These results assume
a Representative Pathway 8.5 warming scenario through 2100 and an Extended
Concentration Pathway 8.5 through 2300. We note that peak talik formation
occurs around 2100.
Methods
Simulations
CLM4.5 provides an accurate characterization of the physical and hydrological
state of permafrost needed to evaluate permafrost vulnerability and identify
key processes (Lawrence et al., 2008; Swenson et al., 2012). CLM4.5 includes
a basic set of permafrost processes to allow projection of permafrost
carbon–climate feedbacks, including snow schemes, vertically resolved SOM
dynamics and soil hydrology, coupled hydraulic and thermal properties in
frozen and unfrozen soils allowing realistic seasonal evolution of the active
layer, and interaction with shallow (0–3 m) and deep
(> 3 m) permafrost C (Swenson et al., 2012; Oleson, 2013;
Koven et al., 2013, 2015; Lawrence et al., 2008). More abrupt thaw processes
affecting permafrost C dynamics and talik formation such as thermokarst or
other thaw-related landscape dynamics changes in wetland or lake distribution
are not accounted for in CLM4.5 (see Riley et al., 2011, for more
discussion).
CLM is spun up to C equilibrium for the year 1850 by repeatedly cycling
through 20 years of pre-industrial climate forcing with CO2
and N-deposition set at 1850 levels. C initialization is achieved via slow
mixing by cryoturbation between the seasonally thawed active layers and
deeper permafrost layers (Koven et al., 2009). Including vertically resolved
processes leads to a sign change in the projected high-latitude C response to
warming, from net C gains driven by increased vegetation productivity to net
C losses from enhanced SOM decomposition (Koven et al., 2011). The soil grid
includes 30 vertical levels that have a high-resolution exponential grid in
the interval 0–0.5 m and fixed 20 cm layer thickness in the
range of 0.5–3.5 m to maintain resolution through the base of the
active layer and upper permafrost, and reverts to exponentially increasing
layer thickness in the range 3.5–45 m to allow for large thermal
inertia at depth. Soil C turnover in CLM4.5 is based on a vertical
discretization of first-order multi-pool SOM dynamics (Koven et al., 2013;
Oleson, 2013), where decomposition rates as a function of soil depth
are controlled by a parameter Zτ (Koven et al., 2015; Lawrence et al.,
2015). This depth control of decomposition represents the net impacts of
unresolved depth-dependent processes. In this study, we
utilize Zτ = 10 m, which yields a weak additional depth
dependence of decomposition beyond the environmental controls and, as
discussed and evaluated relative to Zτ = 1 m
and Zτ = 0.5 m in Koven et al. (2015), results in CLM
permafrost-domain soil C stocks that are in closest agreement
(1582 Pg for Zτ = 0.5 m, 1331 Pg for
Zτ = 1 m, and 1032 Pg for
Zτ = 10 m) with observed estimates (1060 PgC to
3 m depth; Hugelius et al., 2013). This reduction in initial C is due to
higher decomposition rates at depth during the model initialization period.
There is no C below 3.5 m, so additional thaw below 3 m has
a small impact on the C cycle. We note that the relationship applied in
CLM4.5, which implies multiplicative impacts of limitations to decomposition,
is commonly applied in land biogeochemical models but is quite uncertain.
We use CLM4.5 configured as described in two recent permafrost studies
(Lawrence et al., 2015; Koven et al., 2015) using time-varying meteorology,
N deposition, CO2 concentration, and land use change to capture
physiological (i.e., CO2 fertilization) and climate effects of
increasing CO2 over the period 2006–2300. We use an anomaly forcing
method to repeatedly force CLM4.5 with observed meteorological from the
CRUNCEP dataset for the period 1996–2005 (data available at
https://www.earthsystemgrid.org/dataset/ucar.cgd.ccsm4.CRUNCEP.v4.html) and monthly anomalies added based on a single ensemble member
from a CCSM4 Representative Concentration Pathway 8.5 (RCP8.5) simulation for
the years 2006–2100 and Extended Concentration Pathway 8.5 (ECP8.5) for the
years 2100–2300. The period from 1996 to 2015 represents a base
climatological period used for calculating monthly anomalies, with a 20-year
record chosen to minimize large anomalies in the first few years. This
process is repeated for all variables and all times from 2006 to 2300
(constantly cycling through the same 1996–2005 observed data). Land air
temperature for the period 2006–2300, shown in Fig. 1a, is projected to
increase steadily over our simulation, with a slight decrease in the rate of
warming.
We caution that we are using only a single ensemble member from CCSM4, and
hence our results represent one realization from one model forced with one
climate scenario. This results in uncertainties from the historical
climate–weather forcing, the structure and parameterization of the model, and
climate scenarios (both across models and across emissions scenarios).
Simulations are carried out on a global domain at a grid resolution of
1.25∘ longitude × 0.9375∘ latitude and saved as
monthly averages. Simulation output is collected into decadal averages from
2011 to 2300 (e.g., 2011–2020 averages for the 2010s, 2021–2030 for the
2020s). Our method to link C balance changes to permafrost thermal state
relies on identifying the timing of two key processes: (1) talik formation
and (2) C source transition. Talik formation represents a critical threshold
of permafrost thaw. The C source transition represents a shift of ecosystem C
balance from a neutral or weak C sink to a long-term source as C balance
shifts to increasing dominance of C source processes including permafrost
thaw and fires (Koven, Lawrence and Riley, 2015). Using the hypothesis that
talik formation triggers a transition to long-term C sources, we quantify the
extent of talik formation and rate of transition to C source once talik has
formed in permafrost-affected NHL ecosystems.
Site information for long-term borehole temperature measurements
along the East Siberian Transect for the period 1957–1990. The nine sites
reported in this table, presented in a north-to-south order, meet the
criteria of at least 1 year of valid soil temperature data
(≥ 10 months per layer, ≥ 55 months across five layers). Talik is
observed in four of nine sites, two of which is observed in the first year of valid
reported data. Site-specific thaw trends are provided for sites with at least
6 years of valid data. Regional trends are calculated from all
available data for three regional locations.
Site
Location
Date range
Years with
First obs.
Site trend
Region
Regional trend
valid data
talik
(mmonthyr-1)
(mmonthyr-1)
Drughina
145.0∘ E, 68.3∘ N
1969–1990
8
N/A
-0.083
N Siberia
-0.057
Ustmoma
143.1∘ E, 66.3∘ N
1973–1975
3
N/A
N/A
Chumpuruck
114.9∘ E, 60.7∘ N
1981–1984
4
N/A
N/A
SW Siberia
0.019
Lensk
114.9∘ E, 60.7∘ N
1957–1990
11
1957
0.23
Macha
114.9∘ E, 60.7∘ N
1970–1990
13
1970
0.070
Oimyakon
114.9∘ E, 60.7∘ N
1966–1974
6
N/A
0.059
Tongulakh
114.9∘ E, 60.7∘ N
1966–1966
1
N/A
N/A
Uchur
114.9∘ E, 60.7∘ N
1966–1990
17
1974
0.24
Chaingda
130.6∘ E, 59.0∘ N
1967–1990
8
1989
0.51
SE Siberia
0.51
Following Koven et al. (2015), we define the timing of C source transition
from net annual sink to net source as the first decade when annual net biome
production (NBP) decreases below -25 gCm-2y-1 and
remains a source (NBP < 0 gCm-2y-1) through 2300.
Here, we use the sign convention of NBP < 0 to represent net C flux from
land to atmosphere (e.g., source). The timing of talik formation is defined
as the first decade when soil temperature (Ts) for any layer
between 0 and 40 m exceeds -0.5 ∘C for all months in
a calendar year (January–December), assuming that soils start off as
permafrost at the beginning of our simulations in 2006. We use a negative
freezing point threshold to account for availability of liquid water below
0 ∘C due to freezing point depression. We note that the real threshold
temperature at which liquid water remains available varies depending on the
soil salinity or mineral content, the latter effect of which is included in
the actual respiration calculations used by CLM. Here we use
-0.5 ∘C as the freeze–thaw cutoff and examine cutoffs at
0.5 ∘C increments from 0 to -2.0 ∘C.
We introduce the thawed volume–time integral, or “thaw volume”, as a metric
to better understand thaw dynamics and help identify thaw instability
thresholds. We integrate permafrost in both time (month of year) and depth
(soil layer from the surface to 40 m) into a logical function that is
one for thawed layers (Ts > -0.5 ∘C), zero for
frozen layers, and multiply each thawed layer by layer thickness to convert
to units of meter months. This conversion accounts for nonuniform layer
thicknesses, providing a consistent metric for comparing simulated and
observed thaw.
Our analysis focuses on NHL grid points within the ABZ north of
55∘ N. We analyze talik formation and C source transitions in the
context of the simulated initial state of SOM as well as published maps of
permafrost conditions from NSIDC
(https://nsidc.org/data/docs/fgdc/ggd318_map_circumarctic/) and
described in Brown et al. (2001). Permafrost extent is classified as
continuous (90–100 %), discontinuous (50–90 %), and sporadic
(10–50 %).
Observations
We compare simulated patterns of active layer dynamics and soil thaw to
patterns observed from contemporary and historical borehole measurements of
permafrost temperature profiles. We focus on sites in western North America
and eastern Siberia with daily continuous observations year-round
(January–December) over multiple consecutive years. The primary focus of
data in North America (2004–2013) is to evaluate seasonal progression of
soil thaw and talik formation near the surface (0–3 m). Siberian
data, which have a longer record on average (1950–1994), are used to
evaluate long-term trends in soil thaw at 0.0–3.6 m depth. Site
locations are shown in Fig. 1.
Site information for borehole temperature measurements at three sites
along a north-to-south transect in North America for the period 2004–2012.
Climatological soil thermal states presented on a site-to-site basis in Fig. 4
are based on all available valid monthly data for each site, with valid data
requiring at least 20 days of reported data for each layer. Layer of Deepest
thaw represents the deepest layer in which mean soil temperature exceeds
freezing (> -0.5 ∘C) in at least 1 month. Month of latest thaw
represents the latest month in which mean soil temperature exceeds freezing.
Here, we define May as the earliest possible month and April as the latest
possible month.
Site
Location
Date
Soil features:
Depth/
Layer of
Month of
range
surface organic
number of
deepest
latest
layer/soil type
layers
thaw
thaw
Mould Bay,
119.0∘ W,
2004–2012
Low organic layer
3 m/
0.69 m
Sep
Canada
76.0∘ N
(∼ 2 cm)/sandy silt
36
Barrow2,
156.0∘ W,
2006–2013
Low organic layer/
15 m/
0.58 m
Oct
Alaska
71.3∘ N
sandy silt
63
Gakona1,
145.0∘ W,
2009–2013
Thick organic layer
30 m/
2.5 m
Feb
Alaska
62.4∘ N
(50 cm)/silty clay
35
Siberian data are based on measurements along the East Siberian Transect
(EST)
(https://arcticdata.io/metacat/metacat/doi:10.5065/D6Z036BQ/default).
The EST consists of 13 sites that cover a southwest-to-northeast transect in
east Siberia (60.7∘ N, 114.9∘ E to 68.3∘ N,
145∘ E) during the period 1882–1994 (Romanovsky
et al., 2007). For this study, we focus on the nine sites which report
measurements as monthly averages at regular depths of 0.2, 0.4, 0.8, 1.6, and
3.2 m. Unfortunately, data gaps of years to decades exist on
a site-by-site basis, and many years do not report the full annual cycle over
multiple layers. To assess observed thaw trends from 1955 to 1990, we analyze
individual sites which report at least 10 monthsyr-1 of reported
monthly mean soil temperature at each layer, and 55 months across the
5 layers (out of 60 possible layer months per year). Based on these
requirements, we find that six of nine sites yield at least 6 years of data
over multiple decades and are well suited for examining historical thaw
trends. For comparison of observed trends to historical and projected trends
from 1950 to 2300, we analyze clusters of sites by combining the nine sites into
three groups based on approximate locations and calculate observed trends using
the inter-site average at each location. We use two sites in northern Siberia
(67∘ N, 144∘ E), six sites in southwest Siberia
(61∘ N, 115∘ E), and one site in southeast Siberia
(59∘ N, 131∘ E). Site information is shown in more detail
in Table 1.
North American transect data are taken from the global terrestrial network
for permafrost (GTNP) borehole database
(http://gtnpdatabase.org/boreholes): (1) Borehole 1108 at Mould Bay in
Canada (119∘ W, 76∘ N) from 2004 to 2012, (2) Borehole 33 in
Barrow along the northern coast of Alaska (156∘ W, 71.3∘ N)
from 2006 to 2013; and (3) Borehole 848 in Gakona in southeastern Alaska
(145∘ W, 62.39∘ N) from 2009 to 2013. Mould Bay is
a continuous permafrost tundra site with measurements at 63 depths from
0–3 m. Mould Bay has almost no organic layer (about 2 cm)
and then sandy silt with high thermal conductivity. Barrow is a continuous
permafrost tundra site with measurements at 35 depths from 0 to 15 m.
The soil at Barrow is represented by silt with a bit of mix with some
organics and almost no organic layer on top. Conductivity of the upper layer
is ∼ 1 WmK-1 for unfrozen and ≥ 2 WmK-1
for frozen soil. Gakona is a continuous permafrost forest tundra site with
measurements at 36 depths from 0 to 30 m. Gakona has a thick organic
layer of moss (0 to 5 cm), dead moss (from 5 to 13 cm), and
peat (from 13 to 50 cm), then silty clay at depth.
All North American transect datasets are reported as daily averages. For each
site, we aggregate from daily to monthly averages requiring at least
20 daysmonth-1 at each layer and for each year. Measurements are
reported at multiple depths and high vertical resolution (up to 0.1 m
in shallow layers) but are generally nonuniform in depth (multiple layers
missing, different layers reported for each site). Given these
inconsistencies and records ≤ 8 years, we use these data for
qualitative analysis of seasonal and vertical patterns in permafrost thaw.
Site information and soil characteristics are summarized in Table 2.
Results
Simulated talik onset in the 21st century
Our simulations show widespread talik formation throughout Siberia and
northern North America over the period 2010–2300 (Fig. 1b), impacting
∼ 14.5 millionkm2 of land in NHLs (55–80∘ N)
assuming a freeze–thaw threshold of -0.5 ∘C and RCP8.5 and ECP8.5
warming scenarios. In Europe, southwest
Asia, and North America (below 60∘ N), 10.6 millionkm2 of land either formed talik prior to the
start of our simulation in 2010 in regions already experiencing degraded
permafrost (e.g., Fig. 1d, permafrost extent < 90 % in southwest
Siberia and southern North America) or did not have permafrost to begin with.
A small amount of land along northern coastal regions
(∼ 1.6 millionkm2) shows no talik formation prior to 2300.
The long-term trend and decadal variability of talik formation are
quantitatively and qualitatively similar for freeze–thaw thresholds at or
below -0.5 ∘C (Fig. 1a). Peak formation generally occurs over the
period 2050–2150, accelerating rapidly early in the 21st century, and
leveling off in the late 22nd century. The timing and location of talik
formation correlates with the annual mean temperature of permafrost at
3 m (Tsoil-3 m) (Fig. 1c) and observed permafrost state
(Fig. 1d, from Brown et al., 2001) at the start of our simulation; we see
earlier talik formations in sub-Arctic regions (< 66∘ N) with
warm simulated permafrost (Tsoil-3 m > 0 ∘C) and
permafrost extent less than 90 % and later formation in northern regions
with cold permafrost (Tsoil-3 m < 0 ∘C) and
continuous permafrost. Talik formation progresses northward from the
sub-Arctic to the Arctic over time, starting in the warm, discontinuous
permafrost zone in the 21st century then to the cold, continuous permafrost
zone the 22nd century. This suggests a shift in permafrost state across the
pan-Arctic from continuous to discontinuous over the next 2 centuries.
Patterns showing the progression of soil thaw in the decades
surrounding talik onset. Individual lines represent averages across the
subset of talik-forming regions for each decade from the 2050s (darkest red)
to the 2250s (darkest blue). (a) Integrated soil thaw volume, where
the vertical solid line represents the mean timing of initial thaw at depth
and late into the cold season (January–April). Note that the upper limit to
the thaw volume metric in panel (a) is an artifact of the arbitrary
maximum soil depth of 45.1 m in CLM4.5. Other panels show
(b) date of spring surface thaw in the uppermost layer,
(c) annual maximum active layer thickness, and (d) annual
subsurface drainage (solid) and volumetric soil moisture averaged over the
soil column (dashed). Grey shaded areas show the standard deviation of results for
individual talik formation decades. Mean behavior exhibits a characteristic
pattern: gradual increase in thaw volume and active layer depth prior to
talik onset, abrupt shift in thaw volume, and active layer depth, followed by
stabilization to constant thaw volume as soil drying and subsurface drainage
increases.
Our simulations demonstrate consistent patterns of changing thaw volume
leading up to and following initial talik formation, independent of the
decade of talik onset. Time series of thaw volume as a function of decade
relative to talik onset (Fig. 2a) show a steady rise in thaw volume of
1–2 mmonthyr-1 in the decades prior to talik formation, with
thaw limited primarily to shallow soils (< 1.5 m) and summer–early
fall. Thaw volume accelerates to 10–20 mmonthyr-1 within
1–4 decades of talik onset, coinciding with thaw penetration at depth
(∼ 2 m on average, Fig. 2b) and deeper into the cold season
(∼ January–April). Thaw penetration into the January–April period
occurs for the first time at 2.6 ± 0.9 decades prior to talik onset
(vertical grey lines in Fig. 2a). At talik onset, thaw volume jumps from mean
values of 60 ± 10.7 to
377 ± 44 mmonthyr-1 at a mean depth of 4.1 m.
Thaw volume levels out within 1 decade following initial talik formation
and accelerated thaw of all soil layers; this leveling is an artifact of the
maximum depth of soils in CLM4.5 (equal to 45.1 m) and represents
the complete transition from permafrost to seasonally frozen ground in the
model. The transition to deep cold-season thaw and rapidly increasing thaw
volume represent key threshold signaling imminent talik onset.
Evolution of simulated decadal thermal and hydrological state as
functions of month and depth averaged across talik-forming regions in the
2090s. Each panel presents decadal average seasonal profiles in the decades
surrounding talik onset from the 2050s (a) to the
2130s (i). Contours are soil temperature in 0.5 ∘C
intervals, with solid (dashed) lines denoting temperature above (below)
a freeze–thaw threshold of -0.5 ∘C. Stars indicate “thaw” months
where soil temperature exceeds -0.5 ∘C. Color shading is
volumetric soil moisture anomalies relative to the 2040s, where red indicates
drying. Note that soil depth on the y axis is plotted on a nonlinear scale.
The soil thaw profile exhibits a shift from predominantly frozen and wet to
perpetually thawed and drying conditions at depth while remaining seasonally
frozen near the surface.
Onset of surface thaw in the uppermost soils during the spring freeze–thaw
transition provides another reliable predictor for talik onset. In
particular, we find consistent dates and trends of spring thaw in the surface
soil layer in the decades leading up to talik onset (Fig. 2c), shifting by
about 1 week over 4 decades from day of year (DOY) 134 ± 2.8
(∼ mid-May) to DOY 127 ± 3.5 during talik formation
(∼ early May).
Changes in total column soil water and subsurface drainage following talik
onset may provide clues a posteriori that talik is already present. Lawrence
et al. (2015) show that deepening of the active layer and thawing of
permafrost allows water to drain deeper into the soil column, which dries out
near surface soils. Our simulations show a similar, but very slight, drying
pattern in shallow layers in the 4 decades prior to talik onset (1.3 %
loss of soil moisture over 0–1 m depth; Fig. 2d), accounting for
about half of total water storage loss in the column. More significant
changes in water balance occur following talik onset, including more rapid
drying in shallow layers (∼ 10 % over 4 decades) and in the column
(∼ 16 %), and a substantial increase in subsurface drainage, as
discussed below.
Observed and simulated early 21st century soil thermal state as
a function of month and depth for the North American Transect boreholes
(black circles, Fig. 1). (a–c) Observed multi-year means for Mould
Bay, Canada (2004–2012), Barrow, Alaska (2006–2013), and Gakona, Alaska
(2009–2013). The color scale shows the mean temperature and the stars mark
the months when each layer is thawed (T > -0.5 ∘C).
Simulated soil thermal state from 2006 to 2010 for borehole
locations (d–f) and regions with 3 m permafrost temperature
within 0.5 ∘C of observed (g–i) show similar
north-to-south spatial gradient to observations, especially for similar
permafrost temperature. Note that the thaw state at Gakona, Alaska, persists
at depths of 1–3 m into the deep cold season (January–February),
perhaps signaling the threshold for rapid talik formation (see Fig. 3d).
The time evolution of soil vertical thermal and hydrological structure for
the subset of grid cells that form talik in the 2090s is shown in more detail
in Fig. 3. Here, we have subtracted the thermal and hydrological profiles in
the 2040s to show relative change. The 4 decades prior to talik onset are
shown in Fig. 3a–d (2050s–2080s), the decade of talik onset in Fig. 3e
(2090s), and the 4 decades following talik onset in Fig. 3f–i
(2100s–2130s). CLM4.5 represents the process of soil thawing as passage of
a “thaw front” in space and time through soil layers, penetrating and
warming colder, deeper layers, and bringing the frozen soil environment at
depth closer to thermodynamic equilibrium with the warming atmosphere. At
4 decades prior to talik onset (Fig. 3a), our simulated thawed layer exhibits
a tilted time–depth profile with earlier thaw and longer thaw duration
(∼ 4–5 months) in the near surface (< 1 m) compared to
later thaw and reduced thaw duration (1–2 months) at maximum thaw depth
(∼ 2 m). In the 3 decades leading up to talik onset, we find
gradual deepening of the thawed layer to 3–4 m and penetration of
thaw period into January–February.
Our simulations indicate an increased rate of heat transfer and thawing at
depth following talik onset, leading to rapid subsequent thawing, drying, and
decrease in the thickness of the seasonally frozen layer above talik
(Fig. 3e–i). This rapid thawing is depicted in Fig. 2a as the large jump in
thaw volume, and in Fig. 2d as enhanced drying and drainage, with drying
peaking at 3.5–4.5 m depth. In our simulations, talik onset
effectively pulls the “bath plug” that was the ice-filled pore space at
depth, with year-round ice-free conditions allowing soil water to percolate
and be diverted to subsurface drainage (Lawrence et al., 2015). We note that
bedrock soil is not hydrologically active in CLM4.5, and thus the rate of
thawing and drainage in response to permafrost thaw may be overestimated in
deeper CLM4.5 layers near bedrock due to reduced heat capacity.
Soil thaw observation time series from borehole measurements of soil
temperature at sites along the East Siberian Transect over various periods
from 1957 to 1990. Site coordinates are provided in the legend and plotted as
crosses on the map provided in Fig. 1. Thaw trends are derived from estimates
of thawed volume over a depth of 3.2 m for sites with > 55 months
of data over multiple decades: Drughina, Lensk, Macha, Uchur, and Chaingda.
Trend values are reported in Table 1. Vertical dashed lines mark the onset of
talik formation at Lensk (1957), Macha (1970), Uchur (1974), and Chaingda
(1989). Sites in southern Siberia show significant negative thaw volume
trends over the 20th century, representing net increases in soil thaw. The
trend at Drughina is not statistically significant, indicating that soil thaw
is unchanged in northern Siberia.
Our simulated pattern of phase lag for heat transfer to depth mimics observed
thaw profiles in North America (Fig. 4), which are sensitive to latitude and
ecosystem, but with more “vertical” time–depth tilt in CLM4.5 compared to
observations. Borehole data show shallow (∼ 0.5 m) and
seasonally short (∼ 3–4 months from June to September) thaw at the
northernmost tundra site in the Canadian Archipelago (Fig. 4a;
76∘ N, Mould Bay), shallow but longer thaw (5 months from
June to October) moving slightly south to Alaska North Slope (Fig. 4b;
71.3∘ N, Barrow), and deep (∼ 3 m) and seasonally long
(May–February) thaw at the low-latitude continental boreal site in
southeastern
Alaska (Fig. 4c; 62.4∘ N, Gakona). CLM4.5 shows reduced depth and
seasonal duration of thaw when sampled at these specific geographical points,
although the north–south gradient of increasing thaw moving south is
preserved (Fig. 4d–f). Given the challenging task of comparing point
locations with grid cell means, we also examine the mean behavior of CLM4.5
at locations where soil temperature at depth is similar to that observed.
Accounting for permafrost temperature at 3 m (by sampling all
locations with Tsoil-3 m within 0.5 ∘C of the observed
temperature) better reproduces thaw depth, but with reduced seasonal duration
throughout the soil column (Fig. 4g–i). These results suggest the current
ensemble CLM4.5 run overestimates the rate of soil refreeze in early fall.
Comparison of observed soil thaw to historical and future
simulations at sites along the East Siberian Transect (crosses in Fig. 1).
Observed thaw (filled circles) from 1955 to 1990 is based on soil thaw data in
Fig. 5 and on the inter-site average at three locations: northern Siberia (blue),
southwestern Siberia (yellow), and southeastern Siberia (brown). Simulated thaw
from 1950 to 2200 is derived from CLM4.5 and sampled at the nearest grid cell
of three above locations. Asterisks show simulated talik onset. Observed and
simulated thaw trends are derived from soil thaw volume and estimated over
the same period 1955–1990. We note a key discrepancy between observed and
simulated thaw volume: simulated thaw volume is integrated over depths from
0 to 40 m; observed thaw volume is integrated from 0 to 3.6 m.
The effect of this selection bias is a potential low bias in observed thaw
volume. In general, soil thaw is projected to remain stable in northern
Siberia but become increasingly unstable in southern Siberia.
Based on the pattern of January and February freeze–thaw dynamics observed at
Gakona in the 2010s and the time lag of 1–3 decades from this occurrence to
talik onset in our simulations, we project that Gakona will form talik as
early as the 2020s, assuming the atmosphere continues to warm as prescribed
in CLM4.5. Talik onset in CLM4.5 is variable in the region containing Gakona
(southeastern Alaska) with earliest onset by mid-century (∼ 2050s,
Fig. 1a); however, our comparison to borehole temperatures at Gakona suggests
that simulated thaw rates in southwestern Alaska and across pan-Arctic regions
with similar permafrost temperatures are underestimated and that earliest
onset may occur sooner than predicted. Overall, we find that simulated
patterns of permafrost thermal state change are consistent with available
observations but that the exact thaw rates are uncertain. Although there are
many possible explanations for differences in observed and simulated thaw
rates, we can attribute high observed thaw rates in part to a combination of
(1) relatively dry upper soil at Gakona and Mould Bay and (2) low surface
organic layer and high conductivity of the Barrow and Mould Bay soils. We
keep these uncertainties in mind as we examine patterns of change and talik
formation simulated into 2300.
Evaluation of simulated thaw rates and talik onset against Siberian borehole data
The Siberian borehole locations have similar permafrost extent
(> 50 %) to the North American locations according to the Circumpolar
Permafrost Map (Brown, 2001) and similar mean annual air temperature
(∼ -13.6 ∘C) in the 2000s according to CLM4.5. However, air
temperature is more seasonal in Siberia, including colder winters
(4 ∘C colder) and warmer summers (6 ∘C warmer). Spring thaw
for the Siberian sites occurs 2 weeks earlier on average than for the North
American sites in the 2000s, but follows the same pattern of later thaw date
moving north along the borehole transect.
Next we examine thaw trends observed from borehole soil temperature data in
Siberia in the 20th century and evaluate patterns of CLM4.5 projected trends
in the 21st century. We note several caveats in these comparisons: (1) model
simulations are based on only one realization (i.e., model ensemble member)
of historic and future warming and projected permafrost thaw,
(2) availability and access of long-term records in Siberia is limited, and
(3) there is significant variability in space and time in simulated and
observed thaw rates, making direct comparisons challenging. These comparisons
thus serve primarily as a first benchmark for future model analysis and
development.
We focus first on site-specific long-term historical trends by analyzing the
six Siberian borehole sites which recorded at least 55 months and
5 years of temperature data spanning multiple decades: Drughina,
Lensk, Macha, Oimyakon, Uchur, and Chaingda. Records at these locations show
an increase in thaw volume with an average positive trend of
0.19 mmonthyr-1 from 1955 to 1990 (Table 1, Fig. 5). All sites
except Drughina show positive trends, with larger trends in southern
locations, ranging from 0.51 mmonthyr-1 from 1957 to 1990 at
Chaingda in southern Siberia, to a statistically insignificant trend of
-0.083 monthsyr-1 from 1969 to 1990 at Drughina in northeastern
Siberia, suggesting a more or less constant thermal state at this site.
Further examination indicates that active layer thickness at Drughina
actually decreased to 0.8 m from 1989 to 1990 compared to 1.2 m
in the 1970s (data not shown). Drughina also shows smaller average thaw
volume magnitude compared to other sites, consistent with shallower thaw.
Together, these findings indicate that active layer thickness is decreasing
at Drughina.
There is considerable spatial variability in thaw volume and trends, but in
general thaw trends increase from west (0.18 mmonthyr-1) to
east (0.51 mmonthyr-1). Talik forms at several sites, at
different times between 1957 and 1990 (shown by vertical dashed lines on
Fig. 5), with earlier talik to the west consistent with higher mean initial
thaw volumes. We acknowledge the difficulty in identifying talik onset due to
discontinuities in the dataset and limited vertical information; however, we
note that the 15–30-year gap between talik formation in the western
site cluster vs. Chaingda 15∘ east is geographically consistent with
model simulations of later talik formation in eastern Siberia in the 21st
century (Fig. 1b) and thus may represent a gradual expansion of warming into
the east. In general, permafrost appears to be degrading more rapidly at the
southern locations compared to the northern location.
We recompute observed thaw trends at regional clusters using combined records
at the two sites in northern Siberia (blue), six sites in southwest Siberia
(yellow), and one site in southeast Siberia (brown, Table 1) and compare to
historical and projected thaw volume trends in CLM4.5 (Fig. 6). Northern
locations show a consistent pattern of low thaw volume
(< 10 mmonthyr-1) and negligible thaw trend
(∼ 0 mmonthyr-1) in the historical simulations and
observed record from 1950 to 2000. Thaw projections in northern Siberia
indicate an unchanged trend and continued stability of permafrost through the
early 22st century, followed by a shift to accelerated soil thaw in the early
2120, marked by onset of deep soil thaw late in the cold season.
Southern locations show a systematic underestimate of mean thaw volume
(< 20 mmonthyr-1) compared to observations
(∼ 40 mmonthyr-1) from 1950 to 2000. Simulated thaw trends
are negligible prior to 2000, but these likely represent an underestimate
given low simulated thaw volumes and significant positive observed trends in
both southeast and southwest Siberia beginning in the 1960s following talik
onset (Fig. 5). Thaw projections show more abrupt shifts in thaw volume in
the early 21st century in the southwest (∼ 2025) and in the mid-21st
century (∼ 2050) in the southeast. The strong discrepancy between
observed and simulated thaw and talik onset in southern Siberia warrants
close monitoring and continued investigation of this region through sustained
borehole measurements and additional model realizations of potential future
warming.
Carbon cycle responses to changing ground thermal regime
Projected decade when permafrost regions shift to long-term
C sources over the period 2010–2300, and relation to talik onset, soil C,
and fire emissions. (a) Map of the decade of transition to C source,
reflected in the color code, showing earlier transitions in cold northern
permafrost. (b) The area of land that transitions peaks in the late
21st century and is driven by regions where the C source leads talik onset
(dashed). (c) The decadal time lag from talik onset to C source
transition shows positive lags in warm southern permafrost (C source lags
talik) and negative lags in cold norther permafrost (C source leads talik).
(d) Histogram shows trimodal distribution of permafrost area as
a function of decadal time lag, with negative lags related to high soil
organic matter (green bars and map in e) and large positive lags
related to fires (red bars and map in f) but delayed by high
productivity. See text for details. These results assume a Representative
Pathway 8.5 warming scenario through 2100 and an Extended Concentration
Pathway 8.5 through 2300.
Figure 7a plots the decade in which NHL ecosystems are projected to
transition to long-term C sources over the next 3 centuries (2010–2300).
A total of 6.8 millionkm2 of land is projected to transition,
peaking in the late 21st century, with most regions transitioning prior to
2150 (4.8 millionkm2 or 70 %; Fig. 7b, solid black).
C source transitions which occur in the permafrost zone, accounting for
6.2 millionkm2 of land (91 % of all C source transitions),
also form talik at some time from 2006 to 2300 (Fig. 7c). The remaining
C source transitions (0.6 millionkm2, or 9 %) occur outside
the permafrost zone, primarily in eastern Europe.
Cumulative net biome production (NBP) over northern high-latitude
(NHL) regions (> 55∘ N) from 2010 to 2300. NBP < 0 represents
a net C source. NHL regions are divided into the following categories: all
NHL land (diamonds), NHL land regions which form talik from 2010 to 2300
(crosses), and regions which transition to long-term C sources from
2010 to 2300 (black solid). C source transition regions are further broken down
based on the lag relationship between talik onset and C source transition as
follows: regions where the C source transition lags talik onset (dotted),
leads talik onset (dashed), and occurs in the absence of talik (dashed
dotted). C source transition regions also divided by soil C content and fire
activity: regions where soil organic matter (SOM) exceeds
100 kgCm-2 (green), fire emissions exceed
25 gCm-2yr-1 (red), and SOM and fires do not exceed these
thresholds (blue). Regions which transition to C sources prior to talik
formation make up half of the total C source area but account for most of
the cumulative C source (∼ 80 %) in large part due to high soil C.
Net C emissions from C source transition regions are a substantial fraction
of the total NHL C budget over the next 3 centuries (Fig. 8). The cumulative
pan-Arctic C source increases slowly over the 21st century, reaching
10 PgC by 2100 with RCP8.5 warming, then increases more rapidly to
70 PgC by 2200 and 120 Pg by 2300 with sustained ECP8.5
warming (Fig. 8, solid black). This pan-Arctic source represents 86 % of
cumulative emissions in 2300 from the larger NHL talik region (crosses),
despite the 2-fold smaller land area, and exceeds the talik region through
2200 due to mitigating widespread vegetation C gains (Koven et al., 2015).
Cumulative emissions over all NHL land regions (diamonds, > 55∘ N) increase
in similar fashion to the talik region, reaching 120 PgC by 2200
and 220 by 2300, with no sign of slowing.
The geographic pattern of C sink-to-source transition date is reversed
compared to that of talik formation, with earlier transitions at higher
latitudes (the processes driving these patterns are discussed in detail
below). Overall, the lag relationship between talik onset and C source
transition exhibits a trimodal distribution (Fig. 7d), with peaks at
negative time lag (C source leads talik onset, median lag = -5 to
-6 decades), neutral time lag (C source synchronized with talik onset;
median lag = -2 to 1 decade), and positive time lag (C source lags
talik; median lag = 12 decades; red shading in Fig. 7c), each of which is
associated with a distinct process based on soil C and fire emissions as
discussed below. Roughly half of these regions (3.2 millionkm2)
show neutral or positive time lag (lag ≥ 0). This pattern,
characteristic of the sub-Arctic (< 65∘ N), represents the vast
majority of C source transitions after 2150 (Fig. 7b, dotted), but only
accounts for 17 % of cumulative emissions (20 PgC by 2300;
Fig. 8, dotted). The remaining regions (3.0 millionkm2) in the
Arctic and high Arctic (> 65∘ N) show negative time lag and
account for most of late 21st century sources and cumulative emissions
(95 PgC by 2300, or 79 %; Fig. 8, dashed). C sources in regions
not identified as talik (0.63 millionkm2) either show talik
presence at the start of our simulation or are projected to transition in
the absence of permafrost or in regions of severely degraded permafrost
(Fig. 7c, dash dotted). This region contributes only 5 PgC
(4 %) of cumulative C emissions in 2300.
Net biome production (NBP) as a function of thaw volume. Symbols
represent NBP and thaw volume values averaged over regions which transition
to long-term C source from 2060 to 2140, binned into regions where the decade
of C source transition (a) leads talik onset, (b) lags
talik onset, and (c) lags talik onset AND where fires exceed
25 gCm-2yr-1. Colors indicate decade relative to C source
transition, denoted by the large green marker, which occurs when NBP exceeds
-25 gCm-2yr-1 (grey horizontal dashed line). The grey
square marker indicates the mean NBP and thaw volume values during talik
onset. Cases where C source leads talik (a) show small thaw volumes
during C source transition and amplified C sources during talik onset. Cases
where C source lags talik (b)–(c) show large thaw volumes during
C source transition and C sinks during talik onset.
Here, we investigate biological and soil thermal processes driving these
relationships, focusing first on regions where C source transition leads
talik onset (blue shading in Fig. 7c). In these regions, thaw volume is low
(< 50 mmonthyr-1) and shows a weak relationship to NBP (NBP
decreases much faster than thaw volume) prior to C source onset (indicated by
large green circle in Fig. 9a). By the time thaw volume reaches
300 mmonthyr-1 and talik formation occurs, these regions are
already very strong sources (NBP > 150 gCm-2yr-1).
This suggests that C sources in these regions are not driven by respiration
of old C from deep soil thaw, and thus alternative explanations are needed.
Evolution of simulated soil thermal and hydrological state, plotted
as a function of month and depth, for regions which transition to long-term
C sources in the 2060s but do not form talik for another 3 decades
(≥ 2090s). This represents cases where C source leads talik (e.g.,
Fig. 9b). Each panel presents decadal average seasonal profiles in the
decades leading up to C source transition. Shading and contour details are
explained in Fig. 3. These profiles exhibit shifts in thaw period (October),
depth (> 1.5 m), and soil moisture (drying) in the transition
decade.
Closer examination of thermal and moisture dynamics in shallow soils reveals
three potential indicators of C source transition: (1) seasonal duration of
thaw, (2) depth of thaw, and (3) soil drying. For example, vertical profiles
of soil temperature and moisture (Fig. 10) in regions which transition to
C sources in the 2090s show deeper seasonal penetration of soil thaw, a jump
in active layer growth, and enhanced year-round soil drying during the
C source transition decade (Fig. 10d). A broader analysis of soil thaw
statistics over all regions and periods indicates that most C source
transitions (∼ 2.3 millionkm2, or 77 % of land where
C source leads talik) occur at active layer depths below 3 m and thaw
season penetration into November.
Further examination of ecosystem biogeochemistry also shows high initial
C stocks in these regions (red shading in Fig. 7e). The median initial state
of SOM, 109 kgCm-2, is nearly a factor
of 2 larger than the median value in regions where C source lags talik onset
(SOM = 59 kgCm-2). These regions also show 40 % less
gross primary production (median GPP = 755 vs.
1296 gCm-2yr-1) and higher over saturation prior to
C source onset (water-filled pore space at 0.5 m depth at 10, 5, and
2 decades prior = 0.63, 0.59, and 0.57 mm3mm-3 for cold permafrost vs. a near constant value of
0.57 mm-3 in warm permafrost). The total area of land in which SOM
exceeds 100 kgCm-2 represents two-thirds of all land where
C sources lead talik onset (2.0 millionkm2) and peaks at
a negative time lag of -5 to -6 decades (Fig. 7d, green bars), which
perfectly aligns with the peak distribution of negative time lags. Cumulative
C emissions from regions of SOM > 100 kgCm-2 are also
two-thirds of total C emissions (80 PgC; Fig. 8, green). These results
indicate peat-like conditions characterized by saturated soils, high C
stocks, and low annual productivity, which allow low thaw volumes (active
layer depth < 2 m and peak thaw month of October, on average) and
rapid soil drying to produce early C losses in colder environments in the
absence of talik.
In regions where C source transitions lag talik onset (red shading in
Fig. 7c), NBP is strongly sensitive to changes in thaw volume until C source
onset occurs (Fig. 9b), and talik formation occurs when these regions are
weak sinks (NBP > 0 gCm-2yr-1). In general, C source
onset under high thaw volume indicates these regions are more sensitive to
C emissions from deep soil thaw. However, as noted above, neutral and
positive time lags show a bimodal distribution peaking near 0 and 15 decades,
and thus additional explanations are needed. Further examination shows high
fire activity in these regions at the time of C source onset (red shading in
Fig. 7f). The regions where fire C emissions exceed
25 gCm-2yr-1, representing our threshold for C source
transition, are exclusively boreal ecosystems, account for one-third of all land
with negative lags (∼ 1.1 millionkm2) and align perfectly
with the peak distribution of positive time lags (Fig. 7d, red bars) and
cumulative C emissions (20 PgC in 2300, Fig. 8, red). NBP is less
sensitive to thaw volume in regions where fire dominates the C balance, which
are strong C sinks at talik onset (Fig. 9c), where soil C respiration is
13 % less than non-fire regions (median SOMHR = 331 vs.
382 gCm-2yr-1), and productivity is 25 % more (median
GPP = 1548 vs. 1216 gCm-2yr-1). Fire regions are also
28 % drier on average in the surface layer than non-fire regions
(volumetric soil moisture = 0.28 vs. 0.39 mm-3 in summer
(May–September) in the upper 10 cm of soil). These results suggest
that soil thermal processes and talik formation are significant factors
driving C source transition in regions with reduced productivity, but fire
activity, spurred by soil drying, drives C source transition in higher
productivity regions.
The decadal time lag between talik onset and C source transition is more
normally distributed in the remaining region, represented by the residual
grey bars visible in Fig. 7d, which occurs predominantly in cold northern
permafrost in northwestern Siberia, where low SOM (< 100 kgm-2)
and fire emission (< 25 gCm-2yr-1) prevail. This
region has a mean lag of 1 decade from talik onset to C source, with high standard deviation
of lags (±8 decades) reflecting a skewed distribution of GPP; low
productivity in cold permafrost (GPP = 385 gCm-2yr-1)
increases the likelihood that soil thaw will lead to C source transition
prior to talik onset, and high productivity in warm permafrost
(GPP = 1111 gCm-2yr-1) increases the likelihood of
a transition after talik onset. Cumulative C emissions from this region are
on the low end (27 PgC by 2300; Fig. 8, blue) due to low soil C
(SOM = 59 kgCm-2).
Time series of ecosystem C fluxes showing seasonal and decadal
patterns during C source transition. This presents mean and standard deviations over the
period 2040–2270 for (a–b) gross primary production (GPP),
(c–d) sum of respiration from soils (SOMHR) and litter (LITHR),
(e–f) difference of respiration from soils and litter, and
(g–h) net biome production (NBP) where NBP < 0 indicates
source. The left columns show seasonal fluxes during the decade of C source
transition. The right column shows the evolution of decadal mean fluxes in
the 3 decades preceding and following C source transition. Regions where
C source transition leads talik (blue) show similar patterns to regions where
transition lags talik (red), most notably a jump in soil vs. litter
respiration during C source transition (f) corresponding in time and
magnitude to decreasing NBP (h). The primary difference between
regions is the seasonal distribution of SOMHR vs. LITHR (e), which
shows a large soil respiration source throughout the cold season in cases
where C sources lag talik. This indicates an annual source of deep old C.
Independent of the presence of talik, a key effect of an increasing number of
thaw months is an increasing rate of respiration from soil C pools. Warming
and CO2 fertilization increase the rate of photosynthetic C uptake,
increasing soil respiration mainly from younger near-surface C pools; whereas
deeper thawing affects both young and old C pools, so that the depth of thaw
dictates the timing and dominant C age of the net respiration flux. Figure 11
illustrates this with a comparison of decadal respiration trends for SOM
(SOMHR) and litter (LITHR) C pools for C source transitions in the mid-21st
century, for scenarios where C source leads talik onset (blue line, cold
permafrost) and lags talik (red lines, warm permafrost). Here, we examine
combined respiration (SOMHR + LITHR) and respiration difference
(SOMHR - LITHR) from soil and litter C pools.
GPP and combined respiration increase by ∼ 15 %decade-1
for each permafrost regime surrounding the decade of C source transition with
peak fluxes in the growing season (Fig. 11a–d). Combined respiration in cold
permafrost is systematically larger than in warm permafrost in the growing
season (May–September) and smaller in the cold season (October–April). In
particular, combined respiration is effectively zero for the late cold season
(January–April) in cold permafrost and significantly positive in warm
permafrost over the same period. The respiration difference also increases
surrounding the C source transition (Fig. 11e–f), but with two key differences
from combined respiration: (1) the decadal increase is exponential, starting
from a value near zero just 3 decades prior to C source transition, and
(2) peak respiration difference occurs in late summer and early fall. Because
litter respiration in the model is mainly drawing from C pools with short
turnover times, the litter respiration flux equilibrates rapidly to changes
in productivity and thus its change primarily reflects changes to inputs
rather than decomposition rates. Conversely, soil C pools, which have much
longer turnover times, equilibrate much more slowly to the productivity
changes and thus primarily reflect changes to the turnover times.
The trend in the respiration difference in warm and cold permafrost, which
increase by similar amounts (∼ 100 gCm-2yr-1), thus
reflects an increasing dominance of respiration from younger and older soil
C pools, respectively. These trends are identical to the corresponding NBP
trends, which decrease by 100 gCm-2yr-1 over the same
period from neutral to net source (Fig. 11g–h), such that the differences
between GPP and respiration driving the NBP trends are explained almost
entirely by the increasing fraction of soil vs. litter respiration.
Furthermore, warm permafrost shows sustained dominance of soil respiration
during the entire cold season. These results are consistent with an
increasing thaw effect on C budgets during C source transitions, but where
shallow thaw of young soil C dominates in cold permafrost and where talik
formation and deep thaw of old soil C dominate warm permafrost.
These results suggest that where talik forms, soil respiration increases
throughout the year as talik and perennial thaw mobilize deeper old soil C to
respiration. In the absence of talik in colder environments, soil respiration
increases primarily in the non-frozen season due to increased availability of
thawed shallow soil C. The lower GPP in colder regions suggests that
increased availability of substrate for respiration due to plant growth and
soil C accumulation has less impact on C source transition in our simulations
than soil thaw dynamics and the initial state of soil C. Thus, cold
permafrost locations become C sources due only to thaw-season dynamics while
warmer permafrost locations transition to C sources due largely to changes in
cold-season dynamics.
Discussion
Talik formation is widespread in our simulations, affecting half of all
NHL land (∼ 14.5 millionkm2) from
2010 through 2300. Simulations of the vertical thermal structure of soil thaw
leading to talik in CLM4.5 qualitatively reproduce deep soil temperature data
from borehole measurements in Siberia and western North America, although
rates of thaw at these and similar permafrost locations are underestimated.
Space-for-time comparisons along the north–south borehole transect in Alaska
and the Canadian Archipelago show a pattern of deepening and seasonal
expansion of thaw moving from the coldest location of the transect in
northern Canada (Mould Bay) to the warmest location in southeastern Alaska
(Gakona). Gakona shows the characteristic late cold-season thaw penetration
into February at 2–3 m depth which in our simulations signals
imminent talik onset (in the case of Gakona, as soon as the 2020s). Likewise,
projected soil thaw trends in east Siberia are in line with long-term
borehole measurements along the East Siberian Transect, but the rate of talik
formation here is also underestimated.
These comparisons indicate stable permafrost conditions in the colder sites
in Siberia and North America through the 21st century, where thaw is generally
slow, seasonally short, and stable. This suggests talik formation in the
northern Arctic is decades to centuries away, but potentially sooner than the
early 22nd century, as projected by the CLM4.5 simulation. Our analysis finds
more unstable permafrost conditions to the south, with observed talik in the
late 20th century although simulated talik is delayed until the early 21st
century.
Due to the potential for early 21st century talik and discrepancy between
observed and simulated trends in warm permafrost, continued model
investigation of factors controlling the rate of soil thaw is critically
needed. In particular, large-scale drying as projected in CLM4.5 near the
surface (Lawrence et al., 2015) may be restricting heat penetration and
active layer growth in the growing season, especially in organic-rich soils
which have very low thermal conductivity (O'Donnell et al.,
2009; Lawrence et al., 2011, 2012). Controlled experiments demonstrating the sensitivity
of talik to parameters that control soil drying such ice impedance or
baseflow scalars (e.g., Lawrence et al., 2015), and the effect of organic
content and mineral soil texture (Lawrence and Slater, 2008), could provide
key insight on soil thermal dynamics in frozen or partially frozen
conditions. Other factors affecting soil hydrology and carbon cycling not
considered in our CLM4.5 simulations include high spatial resolution in
discontinuous permafrost, shifts in vegetation community, lateral flow
representation, thermokarst activity and other thaw-related changes to the
ground surface, surface slope and aspect, soil heterogeneity, and potentially
several other factors (see Jorgenson and Osterkamp, 2005, for discussion of
some of the many complexities to be considered).
Our simulations show increasing C emissions over time across the talik region
(Fig. 1b), as cumulative NBP becomes increasingly negative (NBP < 0
equals a net C source), reaching a net source of 140 PgC by 2300
(Fig. 8, crosses), consistent with previous estimates of net C balance across
the larger pan-Arctic region from CLM4.5 (∼ 160 PgC; Koven
et al., 2015; Lawrence et al., 2015). Ecosystems which transition from net
C sinks to net C sources represent less than half the total talik area (6.8
of 14.5 millionkm2; Fig. 7a) but account for most
(∼ 85 %) of the cumulative emissions, reaching 10 PgC in
2100, 70 PgC in 2200, and 120 PgC by 2300 (Fig. 8, solid
black). Removing the effect of vegetation C gain (∼ 20 PgC in
2100 and 40 PgC in 2200 and 2300 according Koven et al., 2015), we
estimate a cumulative permafrost emission for C source transition regions of
30 PgC in 2100, 110 PgC in 2200, and 160 PgC in
2300. These numbers are on the low end but consistent with estimates of
permafrost C emissions summarized by Schuur et al. (2015), which range from
37 to 174 PgC by 2100 and 100 to 400 PgC by 2300.
About half of this region (3.2 millionkm2) shows a pattern of
accelerated soil C respiration following talik onset, which shifts the
surface C balance of photosynthetic uptake and litter respiration from net
C sinks to long-term net sources totaling 20 PgC by 2300. The
pattern of C source transition following talik formation is most evident in
warm permafrost in the sub-Arctic, suggesting increased microbial
decomposition with warming soils. We also find evidence of talik-driven soil
drying near the surface associated with increased active layer thickness and
higher available water storage, which can lead to enhanced decomposition
rates by causing soils to be less frequently saturated or anoxic (Lawrence
et al., 2015). At the same time, these regions show high ecosystem
productivity which increases roughly in proportion to respiration and thus
may be driven by combination of warming and increased nitrogen availability
resulting from permafrost thaw (Mack et al., 2004; Natali et al., 2012; Koven
et al., 2015). As such, the transition time to sustained net ecosystem
C source is delayed by 1–2 centuries following talik onset as productivity
continues to outpace respiration as currently observed (Belshe et al., 2013;
Mack et al., 2004), with C balance transitions peaking in the mid- to late
22nd century. In nearly one-third of these regions, an estimated
2 millionkm2 of land, fires are a primary mechanism triggering
C source onset, rather than talik. Consequently, in regions of very high
productivity, talik appears to serve more as an indirect driver of long-term
C sources through accelerated soil drying rather than as a direct driver
through accelerated respiration of deep soil C.
Our estimate of C emissions following talik onset (∼ 20 PgC)
is low compared to the cumulative emissions from all long-term C source
transitions (120 PgC), but likely strongly underestimated. Soil
C is not permitted below 3.5 m in CLM4.5, or in most analogous
models, such that potential decomposition of the ∼ 350 Pg soil
organic C in deep permafrost (yedoma C, > 3 m) is not accounted
for (Hugelius et al., 2014; Jackson et al., 2017). This is significant for
our simulations, which show frequent talik formation and accelerating thaw
volumes below 3 m (e.g., Fig. 3). We therefore caution the reader in
the interpretation of the timing and magnitude of permafrost C emissions
following talik onset in our simulations, which represent a lower bound of
potential emissions based on the current formulation of CLM4.5.
We identify an equally large region of land in the high Arctic, representing
∼ 3.0 millionkm2, which is projected to transition to
a long-term C source much sooner than the sub-Arctic in the absence of talik
and emit 5 times as much C by 2300 (∼ 95 PgC). This region,
distributed across northern Siberia and North America, resembles peatlands
and is characterized by cold permafrost, high soil C stocks and soil
moisture, and low productivity. Thawing in this cold northern permafrost is
limited to young, shallow soils with significantly reduced contributions from
deeper, older C than warm permafrost, but with a factor of 2 higher C stocks.
These C rich soils become increasingly vulnerable to decomposition as they
are exposed to increased warming and drying as active layers deepen and
persist deeper into the cold season. The transition to long-term C sources in
this region peak is expected to peak between 2050 and 2100, nearly a century
prior to talik-driven sources in warm permafrost and decades to centuries
prior to talik onset, which eventually amplifies C sources in this region.
These results have important implications for designing an Arctic monitoring
system to simultaneously detect changes in the soil thermal state and
C state. In particular, C observations should not be limited to warm
permafrost regions of the sub-Arctic, since cold northern permafrost regions
are projected to become C sources much sooner and emit more C even without
forming talik. Our analysis of the seasonal dynamics and vertical structure
of permafrost thaw and soil C emissions provides a general strategy for
concurrent observing warm and cold permafrost based on time of year and depth
of thaw.
Observing warm permafrost will require year-round measurements of ground
thermal state to detect precursors to talik onset including thaw penetration
at depth (∼ 2–3 m) and late into the cold season
(∼ January–February), as well as sustained cold-season C flux
observations to detect changes in C balance associated decomposition and
respiration of deep, old soil C. Continued monitoring of these depths will
require sustained long-term measurements from deep boreholes and increasing
reliance on remote sensing technologies such as electromagnetic imaging
(EMI). In particular, EMI surveys along the continuous/discontinuous
permafrost transition zones during the cold season from November to March are
likely to provide key thermal state diagnostics. Systematic radiocarbon
(14C) measurements, which can be used to partition respiration into
autotrophic and heterotrophic young and old soil components (Hicks Pries
et al., 2015), would provide a valuable tool to help disentangle and track
future C emissions from deep permafrost, especially during the long cold
season when talik enables the microbial decomposition of deep old C and is
the primary source of C emissions.
Observing cold permafrost in the high Arctic is both more urgent, due to
earlier shifts in C balance and larger emissions, and more complicated, due
to challenging observing conditions (remote, cold, and dark) and less
detectable signals in thermal state (e.g., talik) and C age (e.g., depleted
in radiocarbon) change. Our results suggest that sustained observation of year-round soil thermal and hydrological profiles (soil drying; depth and duration
of thaw at 1–2 m depth) using boreholes and EMI surveys and cold-season net CO2 exchange (September–October) using atmospheric
CO2 sensors and eddy covariance towers can help detect changes in
soil thaw and soil vs. litter respiration driving annual C balance changes.
We also recommend an observing network focused on regions rich in soil
organic matter, where our simulations indicate increased sensitivity of soil
decomposition to warming.
Conclusions
Greening trends driven by high-latitude warming and CO2 fertilization
have led to amplification of the contemporary C cycle, characterized by
increasing photosynthetic C uptake during the short growing season and
increasing respiration of recent labile soil C during the cold season (Mack
et al., 2004; Piao et al., 2008; Randerson et al., 1999; Graven et al., 2013; Forkel
et al., 2016; Wenzel et al., 2016; Webb et al., 2016). Our simulations of
C–climate feedbacks with interactive terrestrial biogeochemistry and soil
thaw dynamics indicate this trend continues mostly unabated in NHL
ecosystems. However, sustained warming over the next 300 years drives
accelerated permafrost degradation and soil respiration, leading to
widespread shifts in the C balance of Arctic ecosystems toward long-term net
C source by the end of the 23rd century. Also, 6.8 millionkm2 of land
impacted in Siberia and North America will produce an integrated C source of
90 PgC by 2100 and 120 PgC by 2200. Our projected
permafrost C feedback is comparable to the contemporary land use and land use
change contribution to the annual C cycle.
Our main results emphasize an increasingly important impact of NHL
cold-season warming on earlier spring thaw, longer non-frozen seasons, and
increased depth and seasonal duration of soil thaw. Our simulations are
consistent with soil thaw patterns observed from borehole time series in
Siberian and North American transects during the late 20th and early 21st
centuries. Patterns of deeper and longer thaw drive widespread talik and
expose Arctic soils to increased warming and drying, which accelerates
decomposition and respiration of deep, old C and shifts ecosystem C balance
to a state increasingly dominated by soil respiration.
The timing with which Arctic ecosystems transition to long-term net C sources
depends on a number of factors including talik onset, vegetation
productivity, permafrost temperature, soil drying, and organic matter. The
timing is most sensitive to talik onset in warm permafrost regions in the
sub-Arctic, which account for a total of 3.2 millionkm2 of land,
representing ∼ 50 % of our simulated permafrost region. These
regions are also the most productive, which can delay the transition to net
C source by decades or even centuries. As such, warm permafrost regions
typically do not transition to net C sources until the mid-22nd century.
The cold permafrost region in the northern Arctic, which accounts for an
additional 3.0 millionkm2 of land, transitions to net C source
in the late 21st century, much earlier than warm permafrost and in the
absence of talik. High decomposition rates, driven by warming and drying of
shallow, young C in organic-rich soils, and low annual productivity make this
region perhaps the most vulnerable to C release and subject to further
amplification with future talik onset. This result is surprising given the
region is dominated by tundra and underlain by deep, cold permafrost that
might be thought impervious to such changes.
Rather than thinking of the permafrost feedback as being primarily driven by
a single coherent geographic front driven by talik formation along the
retreating boundary of the permafrost zone, this analysis suggests multiple
modes of permafrost thaw with a mosaic of processes acting in different
locations. C sink-to-source transitions are caused by active layer deepening
in some regions, talik-driven permafrost loss in others, fire-driven changes
in other places, and thaw-led hydrologic change in yet others. Our results
reveal a complex interplay of amplified contemporary and old C cycling that
will require detailed monitoring of soil thermal properties (cold-season thaw
depth, talik formation), soil organic matter content, soil C age profiles,
systematic CO2 flux, and atmospheric 14CO2 measurements to
detect and attribute future C sources. Further investigation of soil thermal
properties and thaw patterns is required to understand C balance shifts and
potential further amplification of emissions from high northern latitudes.