Introduction
The biological carbon pump is defined as the downward transfer of
biologically fixed carbon from the ocean surface to the ocean interior
(Volk and Hoffert, 1985). Global estimates of
particulate organic carbon (POC) export cluster between 5 Pg C yr-1
(Moore
et al., 2004; Lutz et al., 2007; Honjo et al., 2008; Henson et al., 2011;
Lima et al., 2014) and 10 Pg C yr-1
(Laws
et al., 2000; Schlitzer, 2004; Gehlen et al., 2006; Boyd and Trull, 2007;
Dunne et al., 2007; Laws et al., 2011). The physical transfer of dissolved
inorganic carbon to the ocean interior during subduction of water masses is
2 orders of magnitude higher (> 250 Pg C yr-1;
Karleskind
et al., 2011; Levy et al., 2013). The global ocean represents a net annual
CO2 sink of 2.5 Pg C yr-1(Le Quéré
et al., 2013), slowing down the increase in the atmospheric CO2
concentration resulting from anthropogenic activity. Although the Southern
Ocean (south of 44∘ S) plays a limited role in the net air–sea
CO2 flux (Lenton et al., 2013), it
is a key component of the global anthropogenic CO2 sink representing
one-third the global oceanic sink (∼ 1 Pg C yr-1) while
covering 20 % of its surface
(Gruber et al.,
2009). The solubility pump is considered to be the major component of this
sink, whereas the biological carbon pump is considered to be inefficient in
the Southern Ocean and sensitive to iron supply.
Following “the iron hypothesis” in the 1990s (Martin, 1990), iron
limitation of high-nutrient, low-chlorophyll (HNLC) areas, including the
Southern Ocean, has been tested in bottle experiments (de
Baar et al., 1990) and through in situ artificial fertilization experiments
(de
Baar et al., 2005; Boyd et al., 2007). Results from these experiments are
numerous and essentially highlight that the lack of iron limits
macronutrient (N, P, Si) utilization
(Boyd
et al., 2005; Hiscock and Millero, 2005) and primary production
(Landry
et al., 2000; Gall et al., 2001; Coale et al., 2004) in these vast HNLC
areas of the Southern Ocean. Due to a large macronutrient repository, the
biological carbon pump in the Southern Ocean is considered to be inefficient
in its capacity to transfer atmospheric carbon to the ocean interior
(Sarmiento and Gruber, 2006). In the context of micronutrient
limitation, sites enriched in iron by natural processes have also been
studied and include the Kerguelen Islands
(Blain et al.,
2001, 2007), the Crozet Islands (Pollard et al.,
2007), the Scotia Sea
(Tarling et al., 2012) and
the Drake Passage (Measures et al.,
2013). Enhanced primary producer biomass in association with natural iron
supply
(Korb
and Whitehouse, 2004; Seeyave et al., 2007; Lefèvre et al.,
2008)
strongly support trace-metal limitation. Furthermore, indirect seasonal
budgets constructed from studies of naturally fertilized systems have been
capable of demonstrating an increase in the strength of the biological
carbon pump
(Blain et
al., 2007; Pollard et al., 2009), although strong discrepancies in carbon to
iron sequestration efficiency exist between systems. To date, direct
measurements of POC export over seasonal cycles from naturally fertilized
blooms in the Southern Ocean are limited to the Crozet Plateau
(Pollard
et al., 2009; Salter et al., 2012). The HNLC Southern Ocean represents a
region where changes in the strength of the biological pump may have played
a role in the glacial–interglacial CO2 cycles
(Bopp et al., 2003; Kohfeld et al., 2005)
and have some significance to future anthropogenic CO2 uptake
(Sarmiento and Le Quéré, 1996). In this context,
additional studies that directly measure POC export from naturally
iron-fertilized blooms in the Southern Ocean are necessary.
POC export can be estimated at short timescales (days to weeks) using the
234Th proxy
(Coale
and Bruland, 1985; Buesseler et al., 2006; Savoye et al., 2006), by optical
imaging of particles (e.g. Picheral et al., 2010, Jouandet et
al., 2011) or by directly collecting particles into surface-tethered
sediment traps (e.g. Maiti et al., 2013 for a
compilation in the Southern Ocean) or neutrally buoyant sediment traps (e.g.
Salter
et al., 2007; Rynearson et al., 2013). Temporal variability of flux in the
Southern Ocean precludes extrapolation of discrete measurements to estimate
seasonal or annual carbon export. However, seasonal export of POC can be
derived from biogeochemical budgets
(Blain et
al., 2007; Jouandet et al., 2011; Pollard et al., 2009) or be directly
measured by moored sediment traps (e.g.
Salter et al., 2012).
Biogeochemical budgets are capable of integrating over large spatial and
temporal scales but may incorporate certain assumptions and lack information
about underlying mechanisms. Direct measurement by sediment traps rely on
fewer assumptions but their performance is strongly related to prevailing
hydrodynamic conditions (Buesseler et al., 2007a), which
can be particularly problematic in the surface ocean. Measuring the
hydrological conditions characterizing mooring deployments is therefore
crucial to address issues surrounding the efficiency of sediment trap
collection.
The ecological processes responsible for carbon export remain poorly
characterized (Boyd and Trull,
2007). There is a strong requirement for quantitative analysis of the
biological components of export to elucidate patterns in carbon and
biomineral fluxes to the ocean interior
(Francois
et al., 2002; Salter et al., 2010; Henson et al., 2012; Le Moigne et al.,
2012; Lima et al., 2014). Long-term deployment of moored sediment traps in
areas of naturally iron-fertilized production, where significant macro- and
micronutrient gradients seasonally structure plankton communities, can help
to establish links between ecological succession and carbon export. For
example, sediment traps around the Crozet Plateau
(Pollard et al., 2009)
identified the significance of Eucampia antarctica var. antarctica resting spores for carbon transfer to
the deep ocean, large empty diatom frustules for Si : C export stoichiometry
(Salter et al., 2012) and
heterotrophic calcifiers for the carbonate counter pump
(Salter et al., 2014).
The increase in primary production resulting from natural fertilization
might not necessarily lead to significant increases in carbon export. The
concept of “high-biomass, low-export” (HBLE) environments was first
introduced in the Southern Ocean
(Lam and
Bishop, 2007). This concept is partly based on the idea that a strong grazer
response to phytoplankton biomass leads to major fragmentation and
remineralization of particles in the twilight zone, shallowing the
remineralization horizon
(Coale et al., 2004).
In these environments, the efficient utilization and reprocessing of
exported carbon by zooplankton leads to faecal-pellet-dominated, low-POC
fluxes
(Ebersbach
et al., 2011). A synthesis of short-term sediment trap deployments,
234Th estimates of upper ocean POC export, and in situ primary
production measurements in the Southern Ocean by
Maiti et al. (2013) highlighted the inverse
relationship between primary production and export efficiency, verifying the
HBLE status of many productive areas in the Southern Ocean. The iron-fertilized bloom above the Kerguelen Plateau exhibits strong
remineralization in the mixed layer compared to the mesopelagic
(Jacquet et al.,
2008) and high bacterial carbon demand
(Obernosterer et al., 2008), features
consistent with a HBLE regime. Moreover, an inverse relationship between
export efficiency and zooplankton biomass in the Kerguelen Plateau region
supports the key role of grazers in the HBLE scenario
(Laurenceau-Cornec
et al., 2015). Efficient grazer responses to phytoplankton biomass following
artificial iron fertilization of HNLC regions also demonstrate increases in
net community production that are not translated to an increase in export
fluxes
(Lam
and Bishop, 2007; Tsuda et al., 2007; Martin et al., 2013; Batten and Gower,
2014).
POC flux attenuation with depth results from processes occurring in the
euphotic layer (setting the particle export efficiency,
Henson et al., 2012) and processes occurring in
the twilight zone between the euphotic layer and ∼ 1000 m
(Buesseler and Boyd, 2009), setting the transfer efficiency
(Francois et al., 2002). These processes are mainly
biologically driven (Boyd and
Trull, 2007) and involve a large diversity of ecosystem components from
bacteria (Rivkin and
Legendre, 2001; Giering et al., 2014), protozooplankton
(Barbeau et al., 1996), mesozooplankton
(Dilling and Alldredge, 2000;
Smetacek et al., 2004) and mesopelagic fishes
(Davison
et al., 2013; Hudson et al., 2014). The net effect of these processes is
summarized in a power-law formulation of POC flux attenuation with depth
proposed by
Martin
et al. (1987) that is still commonly used in data and model applications.
The b exponent in this formulation has been reported to range from 0.4 to
1.7
(Buesseler
et al., 2007b; Lampitt et al., 2008; Henson et al., 2012) in the global
ocean. Nevertheless, a change in the upper mesopelagic community structure
(Lam et al., 2011) and, more precisely, an
increasing contribution of mesozooplankton
(Lam
and Bishop, 2007; Ebersbach et al., 2011) could lead to a shift toward
higher POC flux attenuation with depth.
In this paper, we provide the first annual description of the POC and PON
export fluxes below the mixed layer within the naturally fertilized bloom of
the Kerguelen Plateau, and we discuss the reliability of these measurements
considering the hydrological and biological context. A companion paper
(Rembauville et al., 2015) addresses our final aim: to identify the
ecological vectors that explain the intensity and the stoichiometry of the
fluxes.
Material and methods
Trap deployment and mooring design
As part of the KEOPS2 multidisciplinary programme, a mooring line was deployed
at station A3 (50∘38.3 S–72∘02.6 E) in the
Permanently Open Ocean Zone (POOZ), south of the polar front (PF; Fig. 1).
The mooring line was instrumented with a Technicap PPS3 (0.125 m2
collecting area, 4.75 aspect ratio) sediment trap and inclinometer (NKE
S2IP) at a depth of 289 m (seafloor depth 527 m; Fig. 2). A
conductivity–temperature–pressure (CTD) sensor (Sea-Bird SBE 37) and a
current meter (Nortek Aquadopp) were placed on the mooring line 30 m beneath
the sediment trap (319 m). The sediment trap collection period started on 21
October 2011 and continued until 7 September 2012. The sediment trap was composed of
12 rotating sample cups (250 mL) filled with a 5 % formalin
hypersaline solution buffered with sodium tetraborate at pH = 8. Rotation
of the carousel was programmed to sample short intervals (10–14 days)
between October and February to optimize the temporal resolution of export
from the bloom, and long intervals (99 days) between February and September.
All instruments had a 1 h recording interval. The current meter failed on
7 April 2012.
Localization of the Kerguelen Plateau in the Indian sector
of the Southern Ocean and detailed map of the satellite-derived surface
chlorophyll a concentration (MODIS level 3 product) averaged over the
sediment trap deployment period. Sediment trap location at station A3 is
represented by a black dot, whereas the black circle represents the 100 km
radius area used to average the surface chlorophyll a time series. Arrows
represent surface geostrophic circulation derived from the absolute dynamic
topography (AVISO product). Positions of the Antarctic Circumpolar Current
core (AAC core), the polar front (PF) and the Fawn Trough Current (FTC) are
shown by thick black arrows. Grey lines are 500 and 1000 m isobaths.
Surface chlorophyll data
The MODIS AQUA level 3 (4 km grid resolution, 8-day averages) surface
chlorophyll a product was extracted from the NASA website (http://oceancolor.gsfc.nasa.gov/) for the sediment trap deployment period.
An annual climatology of surface chlorophyll a concentration, based on
available satellite products (1997–2013), was calculated from the
multisatellite GlobColour product. The GlobColour level 3 (case 1 waters,
4.63 km resolution, 8-day averages) product merging SeaWiFS, MODIS and MERIS
data with GSM merging model
(Maritorena and Siegel, 2005) was
accessed via http://www.globcolour.info. Surface chlorophyll a concentrations derived from GlobColour (climatology) and MODIS data
(deployment year) were averaged across a 100 km radius centred on the
sediment trap deployment location (Fig. 1).
Schematic of the instrumented mooring line against vertical
temperature profiles. The sediment trap and the current meter/CTD sensor
location on the mooring line are shown by white circles. Temperature
profiles performed during the sediment trap deployment (20 October 2011) are
represented by grey curves. The solid black curve
is the median temperature profile
from 12 casts realized on the 16 November 2011. Dashed black lines are the
first and third quartiles from these casts. The grey rectangle represents
the Kerguelen Plateau seafloor. The different water masses are Antarctic
Surface Water (AASW), Winter Water (WW) and Upper Circumpolar Deep Water
(UCDW).
Time series analyses of hydrological parameters
Fast Fourier transform (FFT) analysis was performed on the annual time
series data obtained from the mooring, depth and potential density anomaly
(σθ) that were derived from the CTD sensor. Significant
peaks in the power spectrum were identified by comparison to red noise, a
theoretical signal in which the relative variance decreases with increasing
frequency (Gilman et al., 1963). The red noise signal
was considered as a null hypothesis, and its power spectrum was scaled to the
99th percentile of χ2 probability. Power peaks higher than
99 % red noise values were considered to be statistically significant
(Schulz and Mudelsee, 2002), enabling the identification
of periods of major variability in time series. In order to identify the
water masses surrounding the trap, temperature and salinity recorded by the
mooring CTD were placed in context to previous CTD casts conducted at station A3
during KEOPS1 (39 profiles, 23 January 2005–13 February 2005) and KEOPS2
(12 profiles, from 15 to 17 November).
Sediment trap material analyses
Hydrological properties recorded by the instrument mooring at station A3. (a) Depth of the CTD sensor, (b) salinity, (c) potential temperature, (d) line angle, and (e) current speed. In (a)–(e), grey lines are raw data, and black lines are low-pass-filtered data with a Gaussian filter (40 h window as suggested by the spectral analysis). (f) Direction and speed of currents
represented by vectors (undersampled with a 5 h interval) and (g) wind
rose plot of current direction and intensities (dotted circles are
directions relatives frequencies and colours refer to current speed (m s-1)).
Potential temperature–salinity diagram at station A3. Data
are from the moored CTD (black dots), KEOPS1 (blue line) and KEOPS2 (red
line). Grey lines are potential density anomaly. The different water masses
are Antarctic Surface Water (AASW), Winter Water (WW) and Upper Circumpolar
Deep Water (UCDW).
Upon recovery of the sediment trap the pH of the supernatant was measured in
every cup and 1 mL of 37 % formalin buffered with sodium tetraborate
(pH = 8) was added. After allowing the particulate material to settle to the
base of the sample cup (∼ 24 h), 60 mL of supernatant was
removed with a syringe and stored separately. The samples were transported
in the dark at 4 ∘C (JGOFS Sediment Trap Methods, 1994)
and stored under identical conditions upon arrival at the laboratory until
further analysis. Nitrate, nitrite, ammonium and phosphate in the
supernatant were analysed colorimetrically (Aminot and Kerouel,
2007) to check for possible leaching of dissolved inorganic nitrogen and
phosphorus from the particulate phase.
Samples were first transferred to a Petri dish and examined under
stereomicroscope (Leica MZ8, ×10 to ×50 magnification) to determine and
isolate swimmers (i.e. organisms that actively entered the cup). All
swimmers were carefully sorted, cleaned (rinsed with preservative solution),
enumerated and removed from the cups for further taxonomic identification.
The classification of organisms as swimmers remains subjective, and there is
no standardized protocol. We classified zooplankton organisms as swimmers if
organic material and preserved structures could be observed. Empty shells,
exuvia (exoskeleton remains) and organic debris were considered part of the
passive flux. Sample preservation prevented the identification of smaller
swimmers (mainly copepods), but, where possible, zooplankton were identified
following Boltovskoy (1999).
Following the removal of swimmers, samples were quantitatively split into
eight aliquots using a Jencons peristaltic splitter. A splitting precision
of 2.9 % (coefficient of variation) was determined by weighing the
particulate material obtained from each of four 1/8th aliquots (see
below). Aliquots for chemical analyses were centrifuged (5 min at 3000 rpm)
with the supernatant being withdrawn after this step and replaced by
Milli-Q-grade water to remove salts. Milli-Q rinses were compared with
ammonium formate. Organic carbon content was not statistically different
even though nitrogen concentrations were significantly higher; as a
consequence, Milli-Q rinses were routinely performed. The rinsing step was repeated three
times. The remaining pellet was freeze-dried (SGD-SERAIL, 0.05–0.1 mbar, -30 to 30 ∘C, 48 h run) and weighed three times
(Sartorius MC 210 P balance, precision × 10-4 g) to calculate the total
mass. The particulate material was ground to a fine powder and used for
measurements of particulate constituents.
For particulate organic carbon (POC) and particulate organic nitrogen (PON)
analyses, 3 to 5 mg of the freeze-dried powder was weighed directly into
pre-combusted (450 ∘C, 24 h) silver cups. Samples were decarbonated
by adding 20 µL of 2 M analytical-grade hydrochloric acid
(Sigma-Aldrich). Acidification was repeated until no bubbles could be seen,
ensuring all particulate carbonate was dissolved
(Salter et al., 2010). Samples were dried overnight at
50 ∘C. POC and PON were measured with a CHN analyser (Perkin
Elmer 2400 Series II CHNS/O elemental analyser) calibrated with glycine.
Samples were analysed in triplicate with an analytical precision of less
than 0.7 %. Due to the small amount of particulate material in sample
cups #5 and #12, replicate analyses were not possible. Uncertainty
propagation for POC and PON flux was calculated as the quadratic sum of
errors on mass flux and POC/PON content in each sample. The annual flux
(± standard deviation) was calculated as the sum of the
time-integrated flux.
Discussion
Physical conditions of trap deployment
Moored sediment traps can be subject to hydrodynamic biases that affect the
accuracy of particle collection (Buesseler et al.,
2007a). The aspect ratio, tilt and horizontal flow regimes are important
considerations when assessing sediment trap performance. Specifically, the
line angle and aspect ratio of cylindrical traps can result in oversampling
(Hawley, 1988). Horizontal current velocities of 12 cm s-1 are often invoked as a critical threshold over which particles are
no longer quantitatively sampled (Baker et al., 1988). During
the sediment trap deployment period we observed generally low current speeds
(mean < 10 cm s-1), with 75% of the recorded data lower than
12 cm s-1. Despite the high aspect ratio of the PPS3 trap (4.75), and
the small mooring line angle deviations, it is likely that episodic
increases in current velocities (> 12 cm s-1) impacted
collection efficiency. When integrated over the entire current meter record
(October 2011 to April 2012), the resulting flow is consistent with the
annual northeastward, low-velocity (∼ 1 cm s-1) geostrophic flow
previously reported over the central part of the Kerguelen Plateau
(Park et al., 2008b).
The depth of the winter mixed layer (WML) on the Kerguelen Plateau is
usually shallower than 250 m
(Park
et al., 1998; Metzl et al., 2006). The sediment trap deployment depth of
∼ 300 m was selected to sample particle flux exiting the WML.
The moored CTD sensor did not record any evidence of a winter water
incursion during the deployment period, confirming that the WML did not reach the
trap depth. The small depth variations observed during the deployment period
resulted from vertical displacement of the trap. Variations in σθ may have resulted from both vertical displacement of the CTD
sensor and possible isopycnal displacements due to strong internal waves
that can occur with an amplitude of > 50 m at this depth
(Park et al., 2008a). Our measurements
demonstrate that isopycnal displacements are consistent with the M2 (moon 2,
12.4 h period) tidal forcing described in physical modelling studies
(Maraldi et al., 2009,
2011). Spectral analysis indicates that high-frequency tidal currents are
the major circulation components. Time-integrated currents suggest that
advection is weak and occurs over a longer timescale (months). Assuming the
current flow measured at the sediment trap deployment depth is
representative of the prevailing current under the WML, more than three
months are required for particles to leave the plateau from station A3,
a timescale larger than the bloom duration itself. Therefore we consider
that the particles collected in the sediment trap at station A3 were
produced in the surface waters located above the plateau during bloom
conditions.
Swimmers and particle solubilization
Aside from the hydrodynamic effects discussed above, other potential biases
characterizing sediment trap deployments, particularly those in shallow
waters, are the presence of swimmers and particle solubilization. Swimmers
can artificially increase POC fluxes by entering the cups and releasing
particulate organic matter or decrease the flux by feeding in the trap
funnel (Buesseler et al., 2007a). Some studies have
focused specifically on swimmer communities collected in shallow sediment
traps (Matsuno et al., 2014, and references
therein), although trap collection of swimmers is probably selective and
therefore not quantitative. Total swimmer intrusion rate was highest in cups
#6 to #9 (December 2011 to February 2012) generally through the
representation of copepods and amphipods (Table 2). The maximum swimmer
intrusion rate in mid-summer as well as the copepod dominance is consistent
with the 4-fold increase in mesozooplankton abundance observed from winter
to summer (Carlotti et al., 2015). Swimmer abundance was not correlated with
mass flux, POC or PON fluxes, suggesting that their presence did not
systematically affect particulate fluxes inside the cups. Nevertheless such
correlations cannot rule out the possibility of swimmers feeding in the trap
funnel modifying particle flux collection.
Number of swimmer individuals found in each cup and swimmer
intrusion rate (number per day, bold numbers) for each taxa and for the total
swimmers.
Cup
Copepod
Pteropod
Euphausid
Ostracod
Amphipod
Cnidarian
Polychaete
Ctenophore
Siphonophore
Salp
Total
1
166
13
1
2
1
0
0
0
0
0
183
12
1
< 1
< 1
< 1
0
0
0
0
0
13
2
55
0
0
0
0
0
0
0
0
0
55
4
0
0
0
0
0
0
0
0
0
4
3
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
4
113
0
0
0
0
0
0
0
0
0
113
11
0
0
0
0
0
0
0
0
0
11
5
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
6
540
0
1
0
2
5
1
4
1
0
554
54
0
< 1
0
< 1
< 1
0
0
0
0
55
7
583
0
0
0
0
2
2
3
0
0
590
58
0
0
0
0
< 1
< 1
< 1
0
0
58
8
686
33
2
2
8
5
1
4
0
0
741
49
2
< 1
< 1
1
< 1
< 1
< 1
0
0
52
9
392
14
4
3
121
4
2
0
0
0
540
28
1
< 1
< 1
9
< 1
< 1
0
0
0
38
10
264
69
1
2
18
11
0
2
0
0
367
19
5
< 1
< 1
1
1
0
< 1
0
0
26
11
54
0
0
0
29
4
1
0
0
0
88
1
0
0
0
< 1
< 1
< 1
0
0
0
1
12
1481
44
5
7
2
3
2
0
0
1
1544
15
< 1
< 1
< 1
< 1
< 1
< 1
0
0
< 1
15
Particle solubilization in preservative solutions may also lead to an
underestimation of total flux measured in sediment traps. Previous analyses
from traps poisoned with mercuric chloride suggest that ∼ 30 % of total organic carbon flux can be found in the dissolved phase and
much higher values of 50 and 90 % may be observed for nitrogen and
phosphorous, respectively
(Antia, 2005; O'Neill et al.,
2005). Unfortunately the use of a formaldehyde-based preservative in our
trap samples precludes any direct estimate of excess of dissolved organic
carbon in the sample cup supernatant. Furthermore, corrections for particle
leaching have been considered problematic in the presence of swimmers since
a fraction of the leaching may originate from the swimmers themselves
(Antia, 2005), potentially leading to
overcorrection. Particle solubilization may have occurred in our
samples, as evidenced by excess PO43- in the supernatant. However
the largest values were measured in sample cups where total swimmers were
abundant (cups #8 to #12; data not shown). Consequently, it was not
possible to discriminate solubilization of P from swimmers and passively
settling particles, and it therefore remains difficult to quantify the effect
of particle leaching. However, leaching of POC should be less problematic in
formalin-preserved samples because aldehydes fix organic matter, in addition
to poisoning microbial activity.
Seasonal dynamics of POC export
The sediment trap record obtained from station A3 provides the first direct
estimate of POC export covering an entire season over the naturally
fertilized Kerguelen Plateau. We observed a temporal lag of 1 month
between the two surface chlorophyll a peaks and the two export events. Based
on a compilation of annual sediment trap deployments,
Lutz et al. (2007) reported that export
quickly follows primary production at low latitudes, whereas a time lag up to
2 months could occur at higher latitudes. A 1–2-month lag was observed
between production and export in the pacific sector of the Southern Ocean
(Buesseler et al., 2001), as well
as along 170∘ W
(Honjo et al., 2000) and in
the Australian sector of the Subantarctic Zone
(Rigual-Hernández et al., 2015). The
temporal lag between surface production and measured export in deep traps
can originate from ecological processes in the upper ocean (e.g. carbon
retention in the mixed layer) as well as slow sinking velocities
(Armstrong et al., 2009),
and one cannot differentiate the two processes from a single deep trap
signal. A global-scale modelling study suggests that the strongest temporal
decoupling between production and export (more than 1 month) occurs in
areas characterized by a strong seasonal variability in primary production
(Henson et al., 2014). The study attributes this
decoupling to differences in phenology of phytoplankton and zooplankton and
evokes zooplankton ejection products as major contributors to fast-sinking
particles sedimenting post-bloom.
On the Kerguelen Plateau there is evidence that a significant fraction of
phytoplankton biomass comprising the two chlorophyll peaks is remineralized
by a highly active heterotrophic microbial community
(Obernosterer
et al., 2008; Christaki et al., 2014). Another fraction is likely channelled
toward higher trophic levels through the intense grazing pressure that
supports the observed increase in zooplankton biomass
(Carlotti et al., 2008, 2015).
Therefore an important fraction of phytoplankton biomass increases observed
by satellite may not contribute to export fluxes. Notably, the POC : PON ratio
measured in our trap material is close to values reported for marine diatoms
(7.3 ± 1.2;
Sarthou et al., 2005)
compared to the C : N ratio of zooplankton faecal pellets, which is typically
higher (7.3 to > 15,
Gerber
and Gerber, 1979; Checkley and Entzeroth, 1985; Morales, 1987). Simple mass
balance would therefore suggest a significant contribution of
phytoplanktonic cells to the POC export, which is indeed corroborated by
detailed microscopic analysis
(Rembauville et al.,
2015).
Although we observed increasing contributions of faecal pellet carbon
post-bloom
(Rembauville et al.,
2015), in line with the model output of Henson et al. (2014), differences in phytoplankton and zooplankton phenology do not
fully explain the seasonality of export on the Kerguelen Plateau.
Considering the shallow trap depth (289 m) and typical sinking speed of
100 m d-1 for phyto-aggregates
(Allredge
and Gotschalk, 1988; Peterson et al., 2005; Trull et al., 2008a),
aggregate-driven export following bloom demise would suggest a short lag of
a few days between production and export peaks. The temporal lag of 1
month measured in the present study suggests either slow sinking rates
(< 5 m d-1) characteristic of single phytoplanktonic cells or
faster-sinking particles that originate from subsurface production peaks
undetected by satellite. It is generally accepted that satellite detection
depth is 20–50 m (Gordon and McCluney, 1975), and can
be less than 20 m when surface chlorophyll a exceeds 0.2 µg L-1
(Smith, 1981), which prevents the detection of deep
phytoplanktonic biomass structures
(Villareal et al., 2011). Although
subsurface chlorophyll maxima located around 100 m have been observed over
the Kerguelen Plateau at the end of the productive period, they have been
interpreted to result from the accumulation of surface production at the
base of the mixed layer rather than subsurface productivity features
(Uitz et al., 2009). In support of
this, detailed taxonomic analysis of the exported material highlights diatom
resting spores as major contributors to the two export fluxes rather than a
composite surface community accumulated at the base of the mixed layer. The
hypothesis of a mass production of nutrient-limited resting spores
post-bloom with high settling rates explains the temporal patterns of export
we observed
(Rembauville et al.,
2015). However a better knowledge of the dynamics of factors responsible for
resting spore formation by diatoms remains necessary to fully validate this
hypothesis.
Evidence for significant flux attenuation over the Kerguelen
Plateau
The Kerguelen Plateau annual POC export (98.2 ± 4.4 mmol m-2 yr-1) approaches the median global ocean POC export value comprising
shallow and deep sediment traps (83 mmol m-2 yr-1;
Lampitt and
Antia, 1997), but is also close to values observed in HNLC areas of the POOZ
(11–43 mmol m-2 yr-1 at 500 m;
Fischer et al.,
2000). Moreover, the magnitude of annual POC export measured at
∼ 300m on the Kerguelen Plateau is comparable to deep-ocean
(> 2 km) POC fluxes measured from the iron-fertilized Crozet
(60 mmol m-2 yr-1; Salter et
al., 2012) and South Georgia blooms (180 mmol m-2 yr-1;
Manno et al., 2015).
We first compared the sediment trap export fluxes with short-term estimates
at 200 m in spring (KEOPS2) and summer (KEOPS1). The POC flux recorded in
the moored sediment trap represents only a small fraction (3–8 %) of the
POC flux measured at the base of the winter mixed layer (200 m) by different
approaches during the spring KEOPS2 cruise (Table 3). The same conclusion
can be drawn when considering the comparison with different estimates made
at the end of summer during KEOPS1. Moreover, the annual POC export of
∼ 0.1 mol m-2 yr-1 at 289 m (Table 1) represents only
2 % of the indirect estimate of POC export (5.1 mol m-2 yr-1) at
the base of the WML (200 m) on the Kerguelen Plateau based on a seasonal dissolved inorganic carbon
(DIC) budget (Blain et al., 2007). The short-term estimates
are derived from a diverse range of methods. The 234Th proxy is based
on the 234Th deficit relative to the 238U due to its adsorption on
particles, and its subsequent conversion to carbon fluxes using measured
POC : 234Th ratios.
(Coale
and Bruland, 1985; Buesseler et al., 2006; Savoye et al., 2006). The UVP
(underwater video profiler) provides high-resolution images of particles
(> 52 µm), and the particle size distribution is then
converted to carbon fluxes using an empirical relationship
(Guidi et al., 2008;
Picheral et al., 2010). Drifting gel traps allow for the collection,
preservation and imaging of sinking particles (> 71 µm)
that are converted to carbon fluxes using empirical volume–carbon
relationship
(Ebersbach
and Trull, 2008; Ebersbach et al., 2011; Laurenceau-Cornec et al., 2015).
Finally, drifting sediment traps are conceptually similar to moored sediment
traps but avoid most of the hydrodynamic biases associated with this
technique (Buesseler et al., 2007a). The diversity of
the methods and differences in depth where the POC flux was estimated render
quantitative comparisons challenging. Nevertheless, POC fluxes measured at
289 m with the moored sediment trap are considerably lower than other
estimates made at 200 m. This result indicates either extremely rapid
attenuation of flux between 200 and 300 m or significant sampling bias by
the sediment trap.
Summary of estimates of POC fluxes at the base of, or under,
the mixed layer at station A3 from the KEOPS cruises.
Author
Method
Period
Depth (m)
POC flux
(mmol m-2 d-1)
KEOPS1
23 Jan–12 Feb 2005
100
23 ± 3.6
Savoye et al. (2008)
234Th deficit
150
25.7 ± 3.6
200
24.5 ± 6.8
Drifting gel trap,
4 Feb 2005
200
23.9
Ebersbach and Trull (2008)
optical measurements, and
12 Feb 2005
100
5.3
both constant and power-
200
5.2
law C conversion factor
330
0.7
430
1
Jouandet et al. (2008)
Annual DIC budget
Annual
MLD base
85
Trull et al. (2008b)
Drifting sediment trap
4 Feb 2005
200
7.3–10
12 Feb 2005
200
3–3.1
22 Jan 2005
200
72.4
330
27.2
In situ optical
400
21.6
Jouandet et al. (2011)
measurement (UVP) and power
23 Jan 2005
200
29.8
function C conversion factor
330
26.8
400
15.9
12 Feb 2005
200
4.8
330
5.6
400
7.9
KEOPS2
20 Oct 2011
100
3.5 ± 0.9
150
3.9 ± 0.9
234Th deficit,
200
3.7 ± 0.9
Planchon et al. (2014)
steady-state model
16 Nov 2011
100
4.6 ± 1.5
150
7.1 ± 1.5
200
3.1 ± 0.6
234Th deficit,
16 Nov 2011
100
7.3 ± 1.8
non-steady-state model
150
8.4 ± 1.8
200
3.8 ± 0.8
Laurenceau-Cornec et al. (2015)
Drifting gel trap, optical
16 Nov 2011
210
5.5
measurement of particles
Drifting sediment trap
210
2.2
In situ optical measurement
21 Oct 2011
200
0.2
Jouandet et al. (2014)
(UVP) and power function C
350
0.1
conversion factor
16 Nov 2011
200
1.9
350
0.3
We note that low carbon export fluxes around 300 m have been previously
reported on the Kerguelen Plateau. In spring 2011, UVP-derived estimates of
POC export at 350 m were 0.1 to 0.3 mmol m-2 d-1 (Table 3), values
close to our reported value of 0.15 mmol m-2 d-1. In summer 2005,
POC export at 330 m from a gel trap was 0.7 mmol m-2 d-1 (Ebersbach and Trull 2008), which is also close to
our value of 1.5 mmol m-2 d-1. Using the
Jouandet et al. (2014) data at 200 m (1.9 mmol m-2 d-1) and 350 m (0.3 mmol m-2 d-1) and the
Ebersbach and Trull (2008) data at 200 m (5.2 mmol m-2 d-1) and 330 m (0.7 mmol m-2 d-1) leads to Martin
power-law exponent values of 3.3 and 4, respectively. These values are high
when compared to the range of 0.4–1.7 that was initially compiled for the
global ocean (Buesseler et al., 2007b).
However, there is increasing evidence in support of much higher b values in
the Southern Ocean that fall in the range of 0.9–3.9
(Lam
and Bishop, 2007; Henson et al., 2012; Cavan et al., 2015). Our calculations
are thus consistent with emerging observations of significant POC flux
attenuation in the Southern Ocean.
Using the aforementioned b values (3.3 and 4) and the POC flux derived from
234Th deficit at 200 m in spring
(Planchon
et al., 2014), we estimate POC fluxes at 289 m of 0.7 to
1.1 mmol m-2 d-1. The flux measured in our sediment trap
(0.15 mmol m-2 d-1) data represents 14 to 21 % of this calculated flux. Very
similar percentages (21 to 27 %) are found using the POC fluxes
derived from the 234Th deficit in summer
(Savoye et
al., 2008). Therefore we consider that the moored sediment trap collected
∼ 15–30 % of the 234Th-derived particle flux
equivalent throughout the year. Trap-derived particle fluxes can represent
0.1 to > 3 times the 234Th-derived particles in shallow
sediment traps
(Buesseler,
1991; Buesseler et al., 1994; Coppola et al., 2002; Gustafsson et al., 2004),
and this difference is largely attributed to the sum of hydrodynamic biases
and swimmer activities (Buesseler, 1991), although
it probably also includes the effect of post-collection particle
solubilization. In the Antarctic Peninsula, 234Th-derived POC export
was 20 times higher than the fluxes collected by a shallow, cylindrical,
moored sediment trap at 170 m (Buesseler et
al., 2010). The present deployment context is less extreme (depth of 289 m,
mean current speed < 10 cm s-1, low tilt angle, high aspect
ratio of the cylindrical PPS3 trap) but we consider that hydrodynamics
(current speed higher than 12 cm s-1 during short tidal-driven events)
and possible zooplankton feeding on the trap funnel are potential biases
that may explain in part the low fluxes recorded by the moored sediment
trap. Therefore the low fluxes observed likely result from a combined effect
of collection bias (hydrodynamics and swimmers) and attenuation of the POC
flux between the base of the WML and 300 m. However, it is not possible with
the current data set to isolate a specific explanation for low flux values.
Strong POC flux attenuation over the Kerguelen Plateau compared to the open
ocean is also reported by
Laurenceau-Cornec
et al. (2015), who associated this characteristic with an HBLE scenario and
invoked the role of mesozooplankton in the carbon flux attenuation. Between
October and November 2011, mesozooplankton biomass in the mixed layer
doubled (Carlotti et al., 2014) and summer biomass was a further 2-fold higher (Carlotti et al., 2008). These
seasonal patterns are consistent with the maximum swimmer intrusion rate and
swimmer diversity observed in summer (Table 2). It has previously been
concluded that zooplankton biomass is more tightly coupled to phytoplankton
biomass on the plateau compared to oceanic waters, leading to higher
secondary production on the plateau
(Carlotti et al., 2008, 2014). The findings of Cavan et al. (2015) that document the lowest export ratio (exported
production/primary production) in the most productive, naturally fertilized
area downstream of South Georgia provide further
support linking zooplankton dynamics to HBLE environments of iron-fertilized
blooms. Another important ecosystem feature
associated with the HBLE environment of the Kerguelen Plateau, and likely
shared by other island-fertilized blooms in the Southern Ocean, is the
presence of mesopelagic fish (myctophid spawning and larvae foraging site;
Koubbi et al., 1991, 2001). Mesopelagic fish can be
tightly coupled to lower trophic levels
(Saba and Steinberg, 2012) and can play a
significant role in carbon flux attenuation (Davison et al., 2013). Although
important for carbon budgets, mesopelagic fish represents a compartment often neglected due to the
challenge of quantitative sampling approaches. We suggest that the HBLE
scenario and large attenuation of carbon flux beneath the WML at Kerguelen
may reflect the transfer of carbon biomass to higher and mobile trophic
groups that fuel large mammal and bird populations rather than the classical
remineralization-controlled attenuation characterizing open-ocean
environments. Although technically challenging, testing this hypothesis
should be a focus for future studies in this and similar regions.