Introduction
Peatlands occupy a small fraction of the global land area
(∼ 2–3 %) but store a large proportion of soil carbon
(∼ 50 %) (Gorham, 1991; Yu et al., 2010). Fluxes of carbon (C) in
peatlands are sensitive to variations in weather and climate, but large
uncertainties are associated with scaling process knowledge to landscapes or
regions (Baird et al., 2009), and with high variability within and among
sites (Bubier et al., 2003; Mastepanov et al., 2008; Treat et al., 2007).
Furthermore, there is a relative paucity of in situ measurements for many
globally important peatland types (Christensen, 2014). The majority of global
peatland area is located in boreal to Arctic regions of the Northern
Hemisphere, which has led most of the research on peatland C exchange to
focus on these regions (Lafleur, 2009). However, peatlands also make up an
important component of the landscape in other regions where much less is
known about the size of C stocks, rates and variability of C fluxes, and
sensitivity of those fluxes to environmental and climatic change (Frolking et
al., 2011; Lafleur, 2009; Limpens et al., 2008). Tropical and Southern
Hemisphere peatlands are particularly under-represented in the literature,
despite contributing 10 % of the global peatland area (Yu et al., 2010)
and southern wetlands as a whole potentially contributing > 50 % of
the global wetland methane budget in a given year (Bousquet et al., 2011;
Bridgham et al., 2013). Furthermore, many of the peatlands in these regions
are under pressure of changing land use and increased fire frequency (e.g.,
Page and Hooijer, 2016; Perry et al., 2014) while minimal baseline data are
available to assess the potential changes to regional C budgets. Given the
large amount of C stored in peatland ecosystems, anthropogenic impacts on
peatland function, and the potential for positive feedbacks between peatland
C fluxes and changing climate, a better understanding of C exchange processes
along the full spectrum of peatland types is needed.
In New Zealand and Australia, members of the exclusively Southern Hemisphere
Restionaceae family of rush-like vascular plants are the predominant peat
formers in low-altitude mires (Agnew et al., 1993), and relatively little is
known about their C exchange properties, either on an individual or ecosystem
scale (Meney and Pate, 1999; Goodrich et al., 2015a). In the Waikato region
of New Zealand, peatlands occupy about 5 % of the land area, the majority
of which has been drained for dairy pasture (McGlone, 2009; Pronger et al.,
2014). The draining of peatlands leads to subsidence associated with enhanced
mineralization and compaction (Pronger et al., 2014; Schipper and McLeod,
2002), both of which result in large C losses of the affected area. However,
intact peatland remnants adjacent to drained pastures seem to remain strong
annual sinks for CO2 despite artificially lowered water tables
(Campbell et al., 2014). The vascular plant-dominated vegetation in New
Zealand peatlands has very conservative evaporation rates and high canopy
resistance to water vapor exchange during dry, sunny periods (Campbell and
Williamson, 1997). This behavior also constrains gross primary production
(GPP) when vapor pressure deficit is high (Goodrich et al., 2015a). Lowered
water tables during drought conditions lead to reduced methane fluxes, which
remain low for months after water table recovery, substantially reducing
annual CH4 emissions during drought years (Goodrich et al., 2015b).
However, we do not yet know the full extent of the response of the net
ecosystem C balance (NECB) – and its components – in unaltered New Zealand
peatlands to dry versus wet conditions on seasonal to annual timescales.
Therefore we have little basis for predicting NECB changes in these systems
when environmental changes are imposed from neighboring land use
intensification (e.g., Fetzel et al., 2014) and with the potential for
increasing severity of summer droughts due to climate change (Perry et al.,
2014; Dai, 2013; Trenberth et al., 2014).
Northern Hemisphere peatlands can shift from annual sinks to sources of
CO2 in response to drought (Arneth et al., 2002), and drought-induced
lowering of the water table is generally the most important driver of
inter-annual variability in peatland C exchange (Gažovič et al.,
2013; Herbst et al., 2011; Olson et al., 2013). Drought response of peatlands
is also complicated by the potential for different effects on bogs compared
to fens. For example, the sensitivity of net C exchanges to water table depth
under relatively normal ranges has been shown to be higher in bogs than fens
(Lindroth et al., 2007), but water table and drought effects in
non-Sphagnum-dominated peatlands are not well studied (Fritz et al.,
2011; Cooper et al., 2015; Goodrich et al., 2015b). Furthermore, the relative
impact of drought on each NECB component is not uniform across peatland types
and may vary by dominant vegetation, litter quality, peatland hydrology,
growing season length, nutrient status, or timing of drought (Bubier et al.,
2003; Lund et al., 2012; Sulman et al., 2010). Expanding the coverage of
peatland C flux observations during drought to globally distinct vegetation
types may aid our ability to determine common features and responses that
lead to increased C losses or those that minimize drought effects.
We estimated the net ecosystem C balance (NECB) at a raised ombrotrophic bog
in New Zealand over four years that included one of the most extreme
meteorological droughts in the past 70 years (Porteous and Mullan, 2013). We
used continuous eddy covariance (EC) measurements of CO2 and
CH4 flux and a water balance approach to estimate DOC export in order
to calculate monthly and annual budgets of each C flux component and to
determine the main drivers of variability among years. Carbon flux
measurements and estimated NECB from this site extend the range of climatic
zones represented in peatland literature as well as add information on the
response of a distinctive plant functional type to a wide range of
environmental conditions. We also aim to highlight some useful parallels and
contrasts between this globally unique peatland system and the much better
represented Northern Hemisphere peatlands as well as the growing body of
tropical peatland literature, with respect to drought effects on C fluxes.
Methods
Site description
Kopuatai bog is located in the Hauraki Plains of North Island/Te Ika a
Māui, New Zealand/Aotearoa (37.387∘ S,
175.554∘ E). This
ombrotrophic raised bog is the largest remaining unaltered peatland
(∼ 90 km2) in the country since the majority of New Zealand
wetlands have been drained, primarily for agriculture (McGlone, 2009). The
vegetation at the site is dominated by the jointed wire rush,
Empodisma robustum (Wagstaff and Clarkson, 2012), which forms a
dense canopy (mean height ∼ 0.8 m) of interwoven live and dead stem
material. The total green plant area index (GAI) at the site is
1.32 ± 0.32 m2 m-2 and standing brown living plant material
and dead litter amount to 1.33 ± 0.54 kg m-2 (Goodrich et al.,
2015a), primarily contributed by E. robustum. Sphagnum and
other moss coverage is sparse throughout the peatland, occurring only where
the dominant vegetation is relatively open and light penetrates to the
surface, and therefore the primary peat forming material is E. robustum roots (Agnew et al., 1993). These roots form negatively geotropic clusters covering the surface,
and serve a similar nutrient capture and water holding capacity role to
Sphagnum in Northern Hemisphere peatlands (Agnew et al., 1993;
Clarkson et al., 2009). Peat depths at Kopuatai reach 14 m, with an average peat
accumulation rate of 0.9 mm yr-1 throughout the Holocene (Newnham et
al., 1995).
Eddy covariance CO2, H2O, and CH4 flux
measurements
We measured net ecosystem exchange of CO2 (FCO2) and
H2O (latent heat flux, LE) using the eddy covariance (EC) technique
from 19 November 2011 to 31 December 2015, while methane flux
(FCH4) measurements began on 4 February 2012. Our EC
instrumentation was mounted at 4.25 m above the peatland surface on a 4.5 m
tall triangular lattice tower and included a sonic anemometer (CSAT3,
Campbell Scientific Inc., Logan, Utah, USA), an open-path
H2O / CO2 analyzer (LI-7500, LI-COR Biosciences Inc.,
Lincoln, NE, USA), and an open-path CH4 analyzer (LI-7700, LI-COR
Inc.). Sensors were mounted on a horizontal boom approximately 1.5 m from
the face of the tower with uninterrupted fetch extending > 500 m in all
directions and relatively uniform canopy height and negligible slope over
that distance. Based on the analytical flux footprint model of Kormann and
Meixner (2001), the average distance, centered on the EC tower, within which
80 % of fluxes originated, was 330 m. Data were collected at 10 Hz
using a CR3000 datalogger (Campbell Scientific Inc.) and stored on a memory
card.
Fluxes were processed with an averaging interval of 30 min using the EddyPro
software (v5.1.1, LI-COR Inc.). Time lags between the wind and scalar
concentration time series were removed by covariance maximization. A fully
analytic approach was chosen for correction of low-pass (Moncrieff et al.,
1997) and high-pass filtering (Moncrieff et al., 2005) and the standard Webb
et al. (1980) method was applied to compensate for the effects of air density
fluctuations on LE, FCO2 and FCH4. A double-axis
rotation was applied for sonic tilt correction and the concentration time
series were de-trended by block averaging. Spikes in the high-frequency data
were removed according to Vickers and Mahrt (1997). We utilized the composite
EddyPro quality control flagging system (flags 1–5, with 1 being best
quality) based on tests for steady-state and well-developed turbulence
conditions (Foken et al., 2004; Mauder and Foken, 2006; Foken and
Wichura, 1996; and Göckede et al., 2006). We corrected FCO2
for storage changes in the layer below the EC instruments (with maximum
values on the order of 0.25 µmolm-2s-1 during summer
mornings) based on changes in 30 min CO2 concentrations measured by
the LI-7500, but did not adopt this procedure for FCH4 because
it made very little difference to daily and longer-term sums, instead
introducing more noise. Ancillary measurements included incoming total and
diffuse photosynthetic photon flux density (PPFD) (BF5 Sunshine Sensor,
Delta-T Devices Ltd., Cambridge, United Kingdom) above the canopy
(∼ 1.2 m above the peat surface); incoming and outgoing shortwave and
longwave radiation fluxes and canopy surface temperature (Tsurf)
(NR01, Hukseflux, Delft, the Netherlands) at 2 m height on a secondary mast
5 m northwards of the EC tower; and air temperature (Tair) and
vapor pressure (eair) (fully aspirated HMP 155, Vaisala,
Helsinki, Finland) at 4.25 m above the surface. Water table depth (WTD) was
measured using a submersible pressure sensor (WL1000W, Hydrological Services,
NSW, Australia) suspended within a 1.5 m long dipwell constructed from
50 mm diameter polyvinyl chloride (PVC)
slotted along its length, anchored to a wooden board laced to the peat
surface. Rainfall was measured with a tipping bucket rain gauge (TB3,
Hydrological Services, NSW, Australia).
Quality control, gap filling, and flux partitioning
Fluxes assigned QC flag values > 1 were discarded from the analysis. Data
were then filtered for insufficient atmospheric turbulence using a threshold
for friction velocity (u∗ < 0.15 m s-1), below which all
flux data were discarded. We chose this cut-off after calculating annual sums
of FCO2 and FCH4 using a range of u∗
thresholds and determining the value at which the annual sums stabilized,
following Loescher et al. (2006). In addition, fluxes were discarded when the
associated wind directions fell within a 55∘ sector that included the
tower and site infrastructure.
Gaps in all fluxes were filled using artificial neural networks (ANNs). The
ANN used to fill gaps in FCH4 was described in Goodrich et
al. (2015b). Given that the FCH4 measurements began on
4 February 2012, we used the ANN to estimate January 2012 fluxes and to
obtain a full 4-year dataset. Gaps in the FCO2 time series were
also filled using an ANN approach (Papale and Valentini, 2003), separately
for daytime (PPFD > 5 µmolm-2s-1) and nighttime
(PPFD ≤ 5µmolm-2s-1). The nighttime ANN consisted
of nine input nodes, including air temperature (Tair), peat
temperature at 50 mm below the surface (Tpeat), water table
depth, four fuzzy datasets representing the season, one fuzzy dataset
representing the year of the study period, and an offset node. These “fuzzy
datasets” are transformations of the decomposed time variables (year,
season, month), which provide a way to avoid arbitrary accumulation of time
information in the neural network (Papale and Valentini, 2003). The daytime
ANN had inputs of PPFD, Tair, canopy surface temperature
(Tsurf), atmospheric vapor pressure deficit (VPD), VPD within the
upper canopy estimated using measured Tsurf and eair
(VPDsurf), WTD, and the same fuzzy datasets described for the
nighttime ANN. Both nighttime and daytime ANNs included four hidden nodes,
and sigmoid transfer functions were applied to the weighted sums from the
hidden and output layers (Dengel et al., 2013; Papale and Valentini, 2003).
Since each neural network run gives a unique result, both daytime and
nighttime ANN models were trained and fitted 25 times and the median values
were used to fill gaps (Knox et al., 2014). Similarly, the ANN used to fill
gaps in daytime LE consisted of six input variables (horizontal wind speed,
Tair, VPD, Tsurf, net radiation (Rn), and
VPDsurf), and eight fuzzy variables describing season of year and
time of day. Nighttime gaps in LE were filled with ANN output driven by
Tair, VPD, VPDsurf, Rn, and horizontal
wind speed, and three fuzzy variables describing season of year.
To partition FCO2 into gross primary production (GPP) and
ecosystem respiration (ER), we estimated daytime ER by applying the nighttime
ANN to daytime driver data (Desai et al., 2008; Oikawa et al., 2017). Oikawa
et al. (2017) showed that results from flux partitioning based on neural
networks behaved similarly to those based on the Reichstein et al. (2005)
approach in an alfalfa field. However, both approaches may overestimate GPP
and ER (e.g., 10–13 %, Oikawa et al., 2017) because they rely on
extrapolating measured nighttime ER to daytime, whereas some studies have
demonstrated lower plant respiration during daytime (Kok, 1949; Wohlfahrt et
al., 2005; Wehr et al., 2016). The extent of this effect across all ecosystem
types is unknown (Oikawa et al., 2017; Wehr et al., 2016), so interpretations
based on partitioned ER and GPP should be stated cautiously. For this study,
we applied the standard partitioning approach whereby nighttime GPP was
assumed to be zero and daytime GPP was estimated by subtracting modeled
daytime ER from gap-filled FCO2. We use the term net ecosystem
production (NEP) to refer to monthly and annual summed FCO2,
representing the difference between GPP and ER, so that
NECB=NEP-FCH4-FDOC, and positive
NECB indicates C uptake by the ecosystem.
DOC export
A detailed description of the methods used for estimating C loss via
dissolved organic C (DOC) export in subsurface water (FDOC) was
given by Sturgeon (2013). Briefly, during 2012, DOC was sampled monthly at nine sites across the EC
footprint, at three peat depths, by extracting water from PVC wells sampling
depth ranges 0–0.3, 0.3–0.6, and 0.6–1.0 m. The concentration of DOC in
water samples was determined with a TOC-VCSH analyzer (Shimadzu, Kyoto,
Japan). Monthly water seepage from the EC footprint was estimated with a
water balance approach: Q=P-E-ΔS, where P is rainfall, E is
evaporation and ΔS is change in water storage (all with units mm).
Daily totals of E were calculated from gap-filled time series of 30 min
LE, and ΔS was calculated at monthly time steps from changes in water
table depth multiplied by peat specific yield. FDOC for 2012 was initially calculated as the
product of depth-weighted mean monthly DOC concentration and monthly Q.
There was a strong relationship (R2=0.92) between monthly (P-E) and
FDOC (Sturgeon, 2013), so monthly
FDOC for the whole study period was calculated from this
relationship (Fig. S1 in the Supplement).
Uncertainty estimates
Random uncertainty for each half-hourly value of FCH4 and
FCO2 was estimated based on whether the value was measured or
gap-filled (Dragoni et al., 2007). For measured values we applied the
“paired-days” approach of Hollinger and Richardson (2005) for which the
differences between matching half-hourly fluxes (either ΔFCO2 or ΔFCH4) on adjacent days were examined
if environmental data were similar (PPFD within 75 µmolm-2s-1, Tair within 3 ∘C, wind speed within
1 m s-1). To apply this approach to FCH4, additional
constraints were added for WTD (within 5 mm) and Tpeat (within
2 ∘C) given their influence on CH4 production and flux
(Goodrich et al., 2015b). Double exponential distributions (maximum
likelihood =1/(2β)e-x-μ/β, where β is
the mean of absolute deviations of the samples, and μ is the sample
mean) were fitted to ΔFCO2 and ΔFCH4
binned by flux magnitude and the uncertainty of each measured half-hourly
flux value (σm=(√2)β) was determined as a function
of the mean flux between the measurement pairs (Dragoni et al., 2007;
Hollinger and Richardson, 2005). We utilized the residuals from the 25 ANN
simulations for FCH4 and FCO2 to estimate
uncertainty owing to the gap-filling approach. These residuals were normally
distributed so the standard deviations (σgf) were determined
as functions of the gap-filled flux magnitudes (Dragoni et al., 2007).
Uncertainty in monthly FDOC was calculated as the 95 %
confidence intervals around the predicted value based on monthly P-E
(Fig. S1). Monthly uncertainty values were then combined in quadrature to
obtain annual uncertainty estimates.
Results
Meteorological and hydrological conditions
Mean annual air temperatures were 13.3, 14.0, 13.9, and 13.6 ∘C for
2012–2015, respectively, compared to the 30-year mean of 13.7 ∘C at
an official climate station 11 km to the east of the study site (New Zealand
National Institute for Water and Atmospheric Research, Taihoro Nukurangi).
Annual totals of precipitation (100 % rain) were 1153, 1105, 1086, and
1167 mm compared to the 30-year mean of 1232 mm. Despite annual rainfall
being within 4 % of the mean for all 4 years, the summer (January–March)
rainfall sums in 2013 and 2014 were particularly low, with 65 and 103 mm,
respectively, compared to the much wetter 2012 summer (289 mm) and somewhat
less extreme 2015 summer (176 mm) (Fig. 1). These precipitation patterns
also manifest in late summer minima in water table depth, with 2013
exhibiting the lowest WTD of the measurement period (∼ 300 mm below
the surface) (Fig. 1). Water table depths recharged to within ∼ 50 mm
of the peat surface each winter, responding sharply to rainfall events.
Meteorological and hydrological variables at Kopuatai bog from
December 2011 to December 2015. (a) Daily total incoming
photosynthetic photon flux density (PPFD), (b) daily minimum (gray
dots), maximum (black dots), and 15-day running mean air temperature
(Tair) (line), (c) monthly total rainfall (black bars)
and monthly climatologies (1980–2010) taken from a nearby climate station
(gray bars), and (d) daily mean water table depth (zero line is the
peat surface).
Monthly CO2–C flux components at Kopuatai bog over the 4
measurement years. (a) Gross primary production (GPP),
(b) ecosystem respiration (ER), (c) net ecosystem
production (NEP). Note the different y-axis scale in (c).
Variations in NECB components
Annual NECB totals at Kopuatai bog were 210.2, 134.7, 143.3, and
216.9 gC m-2 yr-1 in the years 2012–2015, respectively
(Table 1). GPP and ER were the largest terms in the budget for all years.
Annual GPP totals were similar for 2012–2014 (ranging from 791.3 to
815.3 gC m-2 yr-1) but larger in 2015
(880.5 gC m-2 yr-1) (Table 1). Monthly GPP was
> 20 gC m-2 for every month of the study period (Fig. 2a),
indicating year-round growing conditions. Ecosystem respiration was roughly
10 % lower in 2012 (570.5 gC m-2 yr-1) than in all other
years (ranging from 629.2 to 636.5 gC m-2 yr-1) (Table 1),
primarily as a result of reduced respiration during the wet summer in 2012
with a generally higher water table (Figs. 1d and 2b). The resulting annual
totals of NEP were 244.9, 161.8, 169.9, and 243.7 gC m-2 yr-1 in
2012–2015 (Table 1).
Annual carbon balance and component fluxes (±estimated
uncertainties, see Methods) at Kopuatai bog from 2012 to 2015 (all units are
gC m-2 yr-1).
Year
CO2–C
Non-CO2–C
NECB
GPP
ER
NEP
FCH4
FDOC
2012
815.3 (±8.6)
570.5 (±9.7)
244.9 (±7.2)
21.9 (±0.4)
12.8 (±0.7)
210.2 (±14.9)
2013*
791.3 (±8.1)
629.2 (±13.3)
161.8 (±12.4)
14.7 (±0.4)
12.4 (±0.7)
134.7 (±19.9)
2014
799.4 (±9.6)
629.2 (±14.2)
169.9 (±14.1)
14.9 (±0.4)
11.7 (±0.7)
143.3 (±22.2)
2015
880.5 (±10.1)
636.5 (±19.0)
243.7 (±14.1)
14.2 (±0.3)
12.6 (±0.7)
216.9 (±25.7)
* Extreme drought year.
Annual FCH4 was a much smaller component of NECB than NEP, with
emissions representing 21.9 (10 %), 14.7 (11 %), 14.9 (10 %), and
14.2 (7 %) gC m-2 yr-1 in 2012–2015 (Table 1) with much
lower monthly fluxes during the drought months and subsequent slow recovery
after water table recharge (Fig. 3a). Annual FDOC contributed a
similarly small proportion but consistent fluxes of 12.8 (6 %), 12.4
(9 %), 11.7 (8 %), and 12.6 (6 %) gC m-2 yr-1 in
2012–2015 (Table 1). FDOC was the most variable flux from month
to month (Fig. 3b), being driven primarily by the water balance (Fig. S1).
Seasonal variation in FCO2
Seasonal variation in diel ensemble CO2 fluxes was relatively
constrained (Fig. 4). Despite significant differences in mid-day (hours
10–14) CO2 uptake among seasons (ANOVA: F=579.4, p<0.001), the
winter mean (-3.2 µmolm-2s-1) was just 34 % lower
than summer mean uptake (-4.8 µmolm-2s-1)
(Fig. 4a, c). Mean nighttime (hours 20–5) CO2 fluxes were also
significantly different among seasons (ANOVA: F=429.7, p<0.001), with
mean winter nighttime losses 40 % lower than in summer (Fig. 4a, c). The
most substantial inter-annual deviations from mean FCO2 patterns
occurred in 2012 when summer mid-day CO2 uptake was 34 % greater
(-5.9 µmolm-2s-1) than the mean of the 3 other years
(-3.9 µmolm-2s-1) (Fig. 4a). Differences among years
were also prominent in spring and autumn mid-day CO2 uptake
(Fig. 4b, d), where 2015 exhibited the largest uptake in both cases.
Monthly total non-CO2–C flux components at Kopuatai bog
over the 4 measurement years. (a) Methane flux (FCH4),
(b) dissolved organic carbon export (FDOC).
Seasonal ensemble average FCO2 (measured data) for
(a) summer, (b) autumn, (c) winter, and
(d) spring over the 4 measurement years.
The bog switched from CO2–C sink to CO2–C neutral or source
1 month earlier in 2013 than any other year owing to the drier conditions and
elevated ER during the 2013 drought coupled with slightly lower GPP
(Fig. 2). However, the bog was a slight
CO2–C source for only 2 months during that year (2013) and neutral
for a third month. During 2015, which was neither abnormally wet nor dry, NEP
was positive for 11 months and neutral for 1 (Fig. 2c).
Controls on ecosystem C fluxes
Variation in monthly NECB was best described by a simple linear regression
with monthly total PPFD, whereby the ecosystem was a significantly stronger C
sink during summer months than during winter months (Figs. 5 and 6). As GPP
was the largest gross term in the budget, the seasonal progression of NECB
(Fig. 5) was generally similar to that of GPP and NEP (Fig. 2a, c),
effectively resulting in light limitation of overall NECB at monthly
timescales (Fig. 6). However, inter-annual differences in monthly NECB were
driven by changes to both ER and GPP (Fig. 7).
Monthly total net ecosystem carbon balance (NECB) at Kopuatai bog
over the 4 measurement years.
Monthly net ecosystem carbon balance (NECB) at Kopuatai bog as a
function of monthly total PPFD from January 2012 to December 2015.
Monthly means of summertime (December–February) and autumn
(March–May) (a) ecosystem respiration (ER) versus peat temperature
(Tpeat) with symbol fill color according to water table depth
(WTD), and (b) gross primary production (GPP) versus integrated photosynthetic photon flux
density (PPFD) with fill color according to daily maximum vapor pressure
deficit (VPD) over the 4 measurement years.
To assess the drivers of ER and GPP, we isolated summer (December–February)
and autumn (March–May) months since differences in mean fluxes between dry
and wet years were largest during these seasons (Figs. 2 and 4). Mean monthly
ER was strongly driven by WTD and Tpeat (Table 2). Higher
Tpeat led to higher respiration (Fig. 7a), while this enhancement
in ER was also exacerbated by lowered WTD (vertical stratification of colors
in Fig. 7a). Accounting for changes in both variables improved the regression
model, explaining about 86 % of the variance in ER, over the simple
models including only WTD or Tpeat (Table 2). The corresponding
variation in mean monthly GPP among years was largely driven by total PPFD
and VPD, whereby higher VPD led to reduced GPP at saturating PPFD (Fig. 7b). Since
changes in VPD were closely correlated with changes in PPFD (Goodrich et al.,
2015a), there were more subtle differences in regression results using one or
the other or both variables in explaining variance in GPP (Table 2) compared
to the equivalent for ER. In addition, a lowered water table per se did not
seem to impact GPP significantly (not shown). Although drier, warmer
conditions had a larger (up to 20 % increase) proportional impact on
summer ER (increasing with lowered water tables and higher Tpeat)
than the 5–18 % decrease in GPP (decreasing with higher VPD), the
contribution of lowered GPP to the overall NECB during those months was
similar to ER (Fig. 7) because of the larger relative magnitude of GPP
(Fig. 2).
Regression statistics for comparison of simple linear single- and
dual-driver models explaining summertime and autumn monthly ER and GPP over
the 4 measurement years. Root mean square error (RMSE) and Akaike's
information criterion (AIC) are given as measures of model error and relative
quality, where a lower AIC value is favorable.
Model
R2
RMSE
AIC
ER ∼ WTD
0.20
50.6
166.2
ER ∼ Tpeat
0.67
21.2
145.3
ER ∼ WTD +Tpeat
0.86
9.3
126.5
GPP ∼ PPFD
0.80
78.0
115.3
GPP ∼ VPD
0.62
147.1
115.2
GPP ∼ PPFD + VPD
0.83
69.9
116.5
Discussion
Peatland net ecosystem C balance
Kopuatai bog was a strong C sink during 4 years with contrasting
environmental conditions that included late summer droughts in 2013 and 2014.
For all 4 measurement years, Kopuatai NECB was much larger
(135–217 gC m-2 yr-1) than published Northern Hemisphere bog
NECB estimates, which range from losses of 14 gC m-2 yr-1 to
gains of 101 gC m-2 yr-1 (Dinsmore et al., 2010; Gažovič
et al., 2013; Koehler et al., 2011; Nilsson et al., 2008; Olefeldt et al.,
2012; Roulet et al., 2007).
The relative contributions of non-CO2–C components to Kopuatai's
NECB (∼ 10 % each) were comparable to those estimated in other
peatland NECB studies (Koehler et al., 2011; Nilsson et al., 2008; Roulet et
al., 2007). However, the relatively short season of C loss at Kopuatai was
largely due to the mild climate that resulted in year-round growing
conditions and relatively large annual NEP (Table 1). This result is in
agreement with Campbell et al. (2014), who found large annual NEP for a
drainage-impacted New Zealand bog despite having similar peak summertime
CO2 uptake to Northern Hemisphere peatlands. The Campbell et
al. (2014) study was conducted at Moanatuatua, a remnant bog with prevalence
of the taller, late successional restiad species, Sporadanthus ferrugineus (giant cane rush, Clarkson et al., 2004) in addition to
E. robustum, resulting in greater mid-day and annual GPP than we
observed at Kopuatai. However, mean nighttime FCO2 (Fig. 4b) and
total ER during summer drought months at Kopuatai (Fig. 2b) were similar to
those observed during summer at Moanatuatua bog, despite water tables
reaching 800 mm below the surface there (Campbell
et al., 2014) compared to < 300 mm below the surface at Kopuatai
(Fig. 1). Similarly, Lafleur et al. (2005) showed that ER at the relatively
dry Mer Bleue bog, in Ontario, Canada, was only weakly correlated with water
table depth. The lack of increase in ER at Mer Bleue with dropping water
tables may have been related to the compensating factor of decreased respiration
from desiccated surface Sphagnum offsetting increased respiration of
deeper heterotrophic microbial communities (Dimitrov et al., 2010). Our
results suggest that lowered water tables increase ER at Kopuatai, but there
may be a limit to this increase. Further significant drops in water tables
during severe drought may only be possible if the vegetation structure were
to undergo a long-term shift away from E. robustum, with its
conservative evaporation regime (Campbell and Williamson, 1997), to
vegetation with higher water use (Thompson et al., 1999).
Growing seasons at Northern Hemisphere peatlands are generally bounded by
frozen or snow-covered winters, but year-round GPP > 0 has been reported
at an Atlantic blanket bog, Glencar, subject to a relatively mild, maritime
climate (McVeigh et al., 2014; Sottocornola and Kiely, 2010). However, mean
summertime peak GPP and ER at Kopuatai (114.4 and
73.0 gC m-2 month-1, respectively, Fig. 2) were substantially
higher than reported for Glencar (63.7 and 38 gC m-2 mo-1,
respectively) (McVeigh et al., 2014), which may be partly due to the lower
peak LAI there (∼ 0.6 m2 m-2) compared to Kopuatai
(1.3 m2 m-2) (Goodrich et al., 2015a), as well as less solar
radiation at the higher-latitude Irish site. In contrast, a moderately rich
treed fen in western Canada with higher LAI (2.61 m2 m-2) had
larger peak GPP and ER than we found at Kopuatai, leading to similar annual
totals (713 and 596 gC m-2 yr-1, respectively) despite a
shorter, 6-month growing season (Syed et al., 2006). Lund et al. (2010)
showed that LAI and growing season length explained a large proportion of the
variance in NEP and its components across a range of northern peatlands. Our
results from Kopuatai bog are consistent with the relationship between
summertime NEP and LAI established by Lund et al. (2010), given the
relatively high LAI and NEP measured here.
Drought effects on Kopuatai NECB and global context
In years with summer/autumn drought (2013 and 2014), including one of the
most severe and widespread meteorological droughts in New Zealand in the past
70 years (Porteous and Mullan, 2013), NECB at Kopuatai bog was reduced by
roughly 30–40 % compared to the relatively wet or meteorologically
“normal” years (2012 and 2015). However, the bog was still a strong C sink
during early drought months in both 2013 and 2014 (Fig. 5) and overall during
the drought years (Table 1). Total GPP in January 2012 and December 2015 was
higher than for any other months during the study period (Fig. 2), which was
likely caused by the low, but still saturating, PPFD and the associated low
VPD conditions (Goodrich et al., 2015a). This also fit within a general
pattern whereby the largest monthly GPP values occurred during saturating
light levels but reduced VPD (Fig. 7b). Similarly, Aurela et al. (2007)
showed that GPP at a sedge fen in Finland was relatively unchanged during a
drought summer, although rates of uptake during clear-sky afternoons within
drought months were suppressed due to high VPD, contributing a small
percentage of the overall drought-induced reduction in NEP.
Some Northern Hemisphere peatlands shift from annual (or growing season)
sinks to sources of CO2 in response to dry conditions (Alm et al.,
1999; Joiner et al., 1999; Shurpali et al., 1995). Reduction in peatland NEP
during dry conditions can result from reduced GPP, increased ER, or a
combination of both. FCH4 tends to be reduced during dry years
(Brown et al., 2014; Moore et al., 2011), while observations of
FDOC during dry years in different peatland types are less
conclusive (Koehler et al., 2011; Roulet et al., 2007). Our results suggest
that both FCH4 and FDOC were lowered during dry
months but that these changes contributed only slightly to the overall NECB
response to dry conditions.
The relative response of GPP and ER to dry conditions has important
implications for the future C sink status of peatlands under changing
climates (Wu and Roulet, 2014). However, there is no consensus on whether
changes in GPP or ER are more important to the peatland NEP drought response
(Lafleur, 2009), and very few data have yet been obtained in tropical and
Southern Hemisphere systems.
Most studies reporting peatland CO2 fluxes during relatively dry
conditions attribute some portion of the NEP reduction to an increase in ER,
including in the tropics (Hirano et al., 2009), and therefore, the ER
response at Kopuatai was expected. However, the reported effects of a lowered
water table per se on GPP are more varied in the literature. Differences in
GPP among years due to WTD changes at our site were small, indicating
relative insensitivity in photosynthetic uptake of E. robustum-dominated peatlands to summer water table drawdown, even during
droughts. Some researchers have reported relatively low sensitivity of annual
GPP to lowered water tables, often due to compensating factors that allowed
NEP to recover despite either temporary reductions in GPP or increases in ER
(Aurela et al., 2007). The only peatland in which GPP has been reported to
increase in response to drought conditions was a treed, moderately rich fen
in western Canada (Cai et al., 2010; Flanagan and Syed, 2011). However, that
site may have been within a successional phase toward increased tree growth
and more above-ground C allocation (Flanagan and Syed, 2011). The
late-successional New Zealand peatland species Sporadanthus ferrugineous (giant cane rush) has deeper roots (Clarkson et al., 2009) and higher above-ground
biomass than the mid-successional E. robustum (Thompson et al.,
1999), dominant at our study site. The large annual CO2 sink strength
reported from the much drier Moanatuatua bog, dominated by S. ferrugineus (Campbell et al., 2014), highlights the need for future work in
New Zealand peatlands to investigate the potential shift in C allocation from
below ground (accumulating root biomass and peat) to above ground (stem and
shoot biomass) resulting from succession or disturbance, such as long-term
lowering of the water table.
Peatlands in which Sphagnum mosses contribute significantly to
ecosystem GPP are particularly sensitive to dry conditions (Shurpali et al.,
1995; Alm et al., 1999; Arneth et al., 2002; Bubier et al., 2003; Lafleur et
al., 2003), which is likely due to the inability of Sphagnum to
control capitulum moisture content when water tables drop (Laitinen et al.,
2008). Sulman et al. (2010) observed opposite responses of fens and bogs to
inter-annual differences in water tables, conjecturing on the importance of
relative Sphagnum cover in accounting for the observed differences.
However, sites dominated by vascular vegetation can also exhibit reduced GPP
with lowered water tables, and the magnitude of the response may depend on
timing of dry conditions. Joiner et al. (1999) found that a late summer
drought led to early autumn senescence of the vascular vegetation at a fen
site in Manitoba, Canada, while ER remained steady until temperatures
dropped. In contrast, Griffis et al. (2000) showed that dry periods during
the early growing season in a sub-Arctic fen, while plants were developing,
led to substantially reduced GPP relative to wetter years and impacted the
whole growing season CO2 uptake, such that the ecosystem was a source
of CO2. Lund et al. (2012) showed a very similar effect of a
springtime-initiated drought on GPP at a Swedish raised bog, where plant
development and moss biomass accumulation were suppressed, impacting NEP over
the course of the year and causing the ecosystem to act as an annual source
of CO2. In contrast, at the same site, a mid-summer drought did not
have the same effect on the vegetation and only ER was affected (increased)
(Lund et al., 2012).
Dominant vegetation type is clearly a critical factor in determining peatland
NEP and NECB response to dry conditions. Timing of drought is also important,
but the impact of timing seems dependent on vegetation type and few examples
of early spring droughts are available in published peatland NEP records.
However, Kupier et al. (2014) used mesocosms from a raised bog to demonstrate
that peatland plant functional types determined when the peatland shifted
from CO2 sink to CO2 source in response to drying.
Vascular plants are better adapted to functioning during dry conditions,
given their ability to control water loss through stomata (Körner, 1995)
and to access water with deeper roots. Furthermore, evergreen species (e.g.,
ericaceous shrubs and restiads) seem to be particularly resilient to drought
stress. E. robustum peatlands may be especially well equipped for
drought given the mulch-like layer of dead stem material that accumulates
above the surface, partially contributing to reduced evaporation rates during
dry conditions (Campbell and Williamson, 1997).
Radiative forcing of the Kopuatai bog greenhouse gas budget
The relative importance of peatland CH4 emission or CO2
uptake to net climate forcing by an ecosystem ultimately depends on the
timescale of interest and the relative flux magnitudes. This is often
assessed using the global warming potential (GWP) approach (IPCC, 1990). If
the standard 100-year GWP factor for CH4 (28, IPCC, 2014) is applied
to the annual C gas fluxes from Kopuatai bog, the result suggests that this
peatland had a net cooling effect on the atmosphere during the 4 measurement
years. GWP fluxes were -78.5, -43.2, -65.5, and
-355.7 gCO2eq. m-2 yr-1 in 2012–2015, respectively (the negative
sign convention here indicates a net greenhouse gas sink). This is not a
surprising conclusion considering the large annual NEP even during drought
years and the very deep peat deposits (as deep as 14 m in places) that have
accumulated over Kopuatai bog's ∼ 10 000-year development (Newnham et
al., 1995), both of which are associated with cooling of the atmosphere
through consistent yearly net removal of atmospheric C. Others have estimated
that annual peatland CH4 fluxes can sometimes offset the C gains from
net CO2 uptake, depending on latitude and environmental conditions in
a given year (Roulet, 2000; Crill et al., 2000; Whiting and Chanton, 2001;
Friborg et al., 2003). However, Frolking et al. (2006) showed that from the
time of peatland formation, the sustained CH4 emissions dominate the
radiative forcing signal for only about 50–100 years before the CH4
effect stabilizes due to a relatively short atmospheric lifetime
(∼ 12 years), while the CO2 uptake effect continues to
accumulate, leading to net atmospheric cooling.
Although abrupt changes in peatland radiative forcing may be possible through
high CH4 emissions because of its high GWP, changes in CO2
dynamics, while dampened in the short term, are much longer-lasting (Frolking
and Roulet, 2007). Losses of DOC may, however, add to the climate warming
potential of a peatland depending on the fate of this C. The contribution of
Kopuatai's DOC losses to the GWP would depend on the eventual loss as either
downstream CO2 or CH4, but is likely small compared to the
measured CH4 fluxes. In some peatland catchments, DOC lost from
peatland margins is bubbled to the atmosphere as CO2 during transport
in neighboring streams, adding substantially to the overall C loss (Billet et
al., 2015), which would contribute to a warming effect.
Radiative forcing considerations have important implications for peatland
restoration efforts because the C balance, and thus the ratio of CH4
emission to CO2 uptake, should be considered a key aspect of a
functioning peatland. For example, loss of C due to peatland drainage for
cultivation in New Zealand (e.g., Pronger et al., 2014) likely has a profound
impact on the relationship between those peatlands and the climate system
(Frolking et al., 2006, 2014; Frolking and Roulet, 2007).
Furthermore, Campbell et al. (2015) showed that grazing on drained peatlands
can result in 190 gC m-2 yr-1 as CO2 lost to the
atmosphere and potentially near 300 gC m-2 yr-1 if the full NECB
is accounted for. Thus the transient warming impact of CH4 emissions
upon re-wetting/re-establishing a peatland during restoration is trivial over
the long run compared to the need to restore the peatland's ability to
accumulate C. Furthermore, Shoemaker and Shrag (2013) illustrated the dangers
of over-valuing the climate impact of CH4 compared to CO2 if
the ultimate goal is to slow the warming effects of anthropogenic activities.
Conclusions
We have shown that a warm temperate bog in New Zealand dominated by the
vascular plant E. robustum was a strong C sink even during drought
years. Our results from Kopuatai bog extend the coverage of ecosystem-scale C
response to a globally unique peatland plant functional type and provide
insight into the role of plants in the drought response of peatlands in
general. Although peak GPP was reduced during dry summer days and ER was
enhanced during drought months, the overall effect was not large enough to
shift the ecosystem to being a CO2 source over the course of a dry
summer/autumn. Furthermore, the importance of summer NEP to annual totals was
reduced due to the year-round growing conditions. The drought resilience of
Kopuatai bog in terms of reduced, but still relatively large, annual carbon
uptake, also provides insight into the existence of these peatlands in a
climatic setting that would not generally be considered conducive to peatland
development and persistence given the often negative summer water balance and
warm annual temperatures (McGlone, 2009). The negative feedback between the
dry conditions and lower evaporation rates (Campbell and Williamson, 1997),
while reducing GPP, helps maintain high water tables, which may limit
respiration losses of C and maintain plant functioning.
Non-CO2–C losses did not contribute to the drought-induced decreases
in C sink strength of Kopuatai bog, as both FCH4 and
FDOC were lowest during dry months. While FCH4 at
Kopuatai is large relative to Northern Hemisphere peatlands and should be
considered an important component of the greenhouse gas balance of the bog,
the ecosystem persisted as a net greenhouse gas sink, according to the GWP
approach, during both relatively wet and dry years covered in this study.