Introduction
Phosphorus (P) is a ubiquitous element found in amino acids, in proteins and
as an integral part of organisms, together with nitrogen (N) and iron (Fe).
It is an essential nutrient that can limit primary production and
nitrogen fixation in aquatic environments and thus significantly influence
carbon storage (Elser et al., 2007). Reviewing experimental data, Moore et al. (2013) proposed two broad regimes of phytoplankton nutrient limitation in
the modern upper ocean: (1) N-limited regimes in most of the low-latitude
oceanic surface and (2) Fe-limited regimes where subsurface nutrient supply is
enhanced (while P may co-limit primary productivity). Moutin et al. (2008)
pointed out the potential importance of phosphate for N2 fixation in
particular in the southeast Pacific under high temperature conditions and Fe
availability, which are favorable for the presence of N2-fixing organisms (like
Trichodesmium spp.) that potentially counteract the N limitation
(Deutsch et al., 2007). However, in some regions like the Mediterranean,
primary productivity is found to be limited by P availability to the marine
ecosystems (Krom et al., 2005). Furthermore, Brahney et al. (2015) and Du
et al. (2016) found that human-driven imbalanced atmospheric deposition of N
and P might have induced or will induce P limitation to the ecosystems
(global alpine lakes and large areas of China's forests, respectively).
Simplified illustration of the atmospheric P cycle showing the
various sources of particulate inorganic and organic P (IPP, OPP) and their
soluble forms (DIP and DOP), the transformation of PP to DP during
atmospheric transport and the deposition of P to the land and to the ocean.
Emissions fractions among atmospheric P forms are those used as input in the
TM4-ECPL model.
The two external-to-the-ocean sources of nutrients are the atmosphere and
rivers. Depending on these inputs and marine dynamics, different nutrients
can limit the marine primary productivity. Riverine inputs of nutrients to
the marine ecosystem are important for coastal regions, while the
atmospheric deposition of nutrients is a more significant source to the open
ocean
(Jickells et al.,
2005; Duce et al., 2008; Mahowald et al., 2008). In contrast to the
atmospheric reactive N pool, the atmospheric soluble P pool is less studied
and remains highly uncertain. Okin et al. (2011) evaluated the impact of Fe and P atmospheric deposition to the ocean
in increasing N2 fixation and found that Fe deposition is more
important than P deposition in supporting N2 fixation, while they
pointed out the large uncertainty in the bioavailability of atmospherically
deposited P. Benitez-Nelson's (2000) review
discussed the importance of discrete pulses of P input to the oligotrophic
seas that have been found to increase the phytoplankton biomass over short
timescales. They also estimated that atmospheric P deposition could be
underestimated by as much as 50 %, when neglecting the P fraction that is
soluble under acidic and high temperature conditions.
In marine ecosystems, the bioavailability of P is found to depend
significantly on its degree of solubility
(Anderson et al., 2010).
Experimentally, bioavailable P is usually considered to be the
“filterable” reactive or total reactive P that passes through a 0.45 µm membrane (Maher and Woo, 1998
and references therein). Although marine organisms, such as cyanobacteria,
have evolved ways to acquire P from mineral sources under P-limited
conditions (Schaperdoth et al., 2007),
phosphate is considered as the most bioavailable form of P
(e.g., Björkman and
Karl, 2003). Experiments have also shown that human-produced P-containing
organics, such as organophosphorus pesticide breakdown products, can also be
utilized by bacteria (Cook et al., 1978). Moreover,
aerosol samples originating from combustion P sources were found to be more
soluble and possibly more bioavailable than those from mineral sources
(Anderson et al., 2010).
Atmospheric P has a variety of sources (Fig. 1), including mineral dust,
combustion products of natural and anthropogenic origin, agricultural
activities (fertilizers and insecticides), bioaerosols, volcanic emissions,
sea spray and phosphine from freshwater wetlands
(Mahowald
et al., 2008; Tipping et al., 2014; Wang et al., 2014; Brahney et al.,
2015). The total P (hereafter TP) found in natural waters can be grouped in
two major forms (Maher and Woo,
1998): (1) the particulate P (PP) and (2) the soluble P, often termed dissolved
P (DP). The PP mainly originates from mineral material (e.g., hydroxyapatite,
brushite, fluorapatite, variscite, strengite and wavellite) as well as P
absorbed to mixed phases (e.g., clay P, clay-organic P and metal
hydroxide P). The DP, on the other hand, includes orthophosphates (i.e.,
H2PO4-, HPO42-, PO43-; hereafter referred
to as PO4) and inorganic condensed P (pyro-, meta- and polyphosphates).
However, both PP and DP can also contain organic P (OP) of both natural and
anthropogenic origin. Naturally emitted OP can be sugar P, inositol P,
phospholipids, phosphoproteins and phosphoamides mainly associated with plants,
animal and bacterial cellular materials
(Maher and Woo, 1998), commonly
present in atmospheric aerosols of biological origin. In addition,
orthophosphate monoesters are known products of ribonucleic acid (RNA) and
lipid degradation that dominate the OP pool in the marine environment,
which also contains orthophosphate diesters and phosphonates
(Paytan et al., 2003).
Mineral dust has been estimated to be the largest external-to-the-ocean
source of bioavailable P
(Mahowald et al., 2008). These
authors estimated a global P mineral source of 1.15 Tg-P yr-1 by
taking into account a typical observed P fraction of 720 ppm in dust
emissions. They also applied a constant solubility fraction of 10 % on the
dust mineral source, based on the observations of
Baker et al. (2006a, b) for
Saharan P-containing aerosols over the Atlantic Ocean, in order to estimate
the soluble P source associated with mineral dust. Recently published
aerosol and deposition observations of African dust in Miami and Barbados
(Zamora et al., 2013) suggest a total P content of
about 880 ppm, which is in the range of P fraction in dust from earlier
studies (roughly 700–1090 ppm, as reviewed by
Mahowald et al., 2008).
Furthermore, based on OP : OC atomic ratios of 0.001–0.009 observed in several
types of soils, Kanakidou et al. (2012) calculated that about 0.03 Tg-P yr-1 of OP (10 % of which
is soluble) is also emitted together with soil dust in the global
atmosphere.
P-containing dust solubilization in deliquesced mineral dust aerosols is
expected to significantly contribute to the soluble inorganic forms of P
(DIP) in the atmosphere. Nenes et al. (2011) suggested that dissolution of apatite minerals
(i.e., Ca5(PO4)3(OH,F,Cl)) under acidic conditions can explain the
observed DIP levels over the eastern Mediterranean, a characteristic region
where Saharan dust can interact with polluted air masses from Europe and the
Middle East. Under acidic atmospheric conditions, H+ can react with the
PO4 and the OH or F groups in the crystal surface, weakening the
Ca2+ bonds and thus causing phosphate to be mobilized from the crystal surface
(Christoffersen and Christoffersen, 1981).
Hence, mineral dust acid dissolution under polluted conditions can
potentially increase the bioavailable P supply into oceanic regions and
further stimulate the net primary production of marine ecosystems
(Nenes et al., 2011).
Primary P sources from combustion processes of anthropogenic and biomass
burning origin are estimated to contribute significantly to global P fluxes
in the atmosphere
(Mahowald
et al., 2008; Tipping et al., 2014; Wang et al., 2014; Brahney et al.,
2015). However, the estimates of global strength of the primary P combustion
source vary by about an order of magnitude on the global scale, due to the
consideration of different forms of the emitted P (i.e., residual or
P-containing ash, gaseous or particulate P produced during combustion
processes; Wang et al., 2014) and different
size distributions in the emitted P-containing particulate matter.
Mahowald et al. (2008) used observed mass ratios of P to black carbon (BC) for fine (< 2 µm) and coarse (2 µm ≤ mean particle diameter
< 10 µm)
particles (Mahowald et al., 2005) and calculated
emission fluxes from biomass burning and anthropogenic fuel (i.e., fossil
fuel and biofuel) combustion of 0.03 and 0.05 Tg-P yr-1,
respectively.
Tipping et al. (2014) estimated a global atmospheric P emission flux of 3.7 Tg-P yr-1 by combining observed deposition rates over land together with
modeled deposition rates over the ocean. This emission flux also accounts
for P deposition fluxes of larger particles (i.e., primary biological
material in the aerosol mode > 10 µm) that are
mainly deposited very close to their source region and thus not long-range
transported. On the other hand, Wang et al. (2014), by assuming that combustion processes emit significant amounts of P
as large particles > 10 µm (hereafter as super-coarse
particles), calculated that P emissions from biomass burning and
anthropogenic combustion processes can contribute about 0.7
and 1.8 Tg-P yr-1, respectively. In contrast to that study, which was
more focused on the impact of anthropogenic combustion on the global P
source, Brahney et al. (2015) extended the methodology of
Mahowald et al. (2008) in a
more explicit aerosol size manner by taking into account also the naturally
emitted super-coarse P-containing particles (i.e., dust, primary biological
material and sea salt). Brahney et al. (2015) showed that considering this
super-coarse fraction as an additional P source, the estimated deposition
fluxes close to the source areas where large particles are emitted (e.g.,
Tipping et al., 2014) can be significantly improved.
The sea surface microlayer can also act as an atmospheric source of P in the
marine environment (Graham et al.,
1979). Correlations between sea-salt fluxes and seawater P concentrations
revealed a 10- to 200-fold enrichment of P content in sea-salt particles
compared to seawater Na concentrations
(Graham
and Duce, 1979; Graham et al., 1979). However, this enrichment was found to
decrease with increasing wind velocity, introducing significant uncertainty
in the strength of the oceanic flux of P on a global scale.
The Vet et al. (2014) review of deposition observations pointed out that sea spray may be a weak contributor to
atmospheric P. Mahowald et al. (2008), taking into account a constant Na concentration in seawater of 10.781 g-Na kg-water-1 and surface seawater phosphate concentrations from the
NOAA Data Center, calculated a global annual flux of soluble P of 0.0049 Tg-P yr-1 (accounting for particles up to 10 µm in diameter;
i.e., PM10). Wang et al. (2014) used a total
oceanic emission P flux of 0.16 Tg-P yr-1 that was calculated as the
average of earlier estimates (ranging from 0.005–0.33). Additionally,
Paytan et al. (2003) found that OP in the seawater particulate matter can be up to 80 % of
total P. Based on an OP / Na mass ratio of 0.02 % as observed by
Graham and Duce (1979), Kanakidou et al. (2012)
estimated that the surface global ocean may also emit 0.19 –0.80 Tg-P yr-1 in the form of OP.
Bioaerosols are P carriers (Mahowald et al., 2008) that can
significantly contribute to the OP budget in the atmosphere
(Kanakidou et al., 2012). These primary biological aerosol particles (hereafter PBAPs) usually range from
10 nm to roughly 100 µm in diameter and, depending on their size, origin
and type, can be transported over long distances. PBAPs can be either alive,
dead, dormant (e.g., bacteria, viruses and fungi spores) or products released
from living organisms such as pollen (Ariya et al., 2009). Mahowald et al. (2008)
calculated that PBAPs contribute 0.165 Tg-P yr-1 to global DP emissions, while Kanakidou et al. (2012),
based on organic carbon (OC) estimates of PBAP emissions and by using a
OP : OC atomic ratio of 0.001, calculated that PBAPs contribute about 0.13 Tg-P yr-1 to global OP emissions. Large uncertainties, however, are
associated with this estimate since it relies on the applied OP : OC ratios of
PBAPs that have been observed to range over 2 orders of magnitude from about
0.0002 up to 0.02 (Kanakidou
et al., 2012 and references therein) and on the simplified approximation of
the density (1–1.2 g cm-3) used for the conversion of the PBAP number
fluxes to mass units
(Burrows et al., 2009a, b). Mahowald et al. (2008)
and Kanakidou et al. (2012)
assumed half of the PBAP source to be hydrophilic, while
Heald and Spracklen (2009) assumed all PBAPs to be totally
hydrophilic particles using an OM : OC ratio of 2.6 that is based on
observations of fungal spores as proposed by
Bauer et al. (2008). However,
bacteria (e.g., P. syringae) are considered as rather insoluble bioaerosols, in contrast
to the water-soluble fractions of highly polar sugars (fructose, glucose,
sucrose, trehalose) and sugar alcohols (arabitol, inositol, mannitol),
mainly contained in pollen grains and fungi spores
(Ariya et al.,
2009). Ageing during atmospheric transport is also expected to increase
bioaerosols' solubility, converting a fraction of their insoluble OP content
to soluble OP (DOP) due to the uptake of oxidants and the formation of
larger chains of soluble multifunctional groups
(Ariya et al.,
2009). Regardless of bioaerosols being hydrophilic or not, because they
consist of biological material, they are expected to be bioavailable
(Bjórkman and Karl, 1994). The degree
of hydrophilicity therefore is more important for determining the relative
importance of dry and wet deposition during their supply to the oceans.
In the present study, the 3-D chemical transport global model TM4-ECPL is
used to integrate current knowledge on the atmospheric P cycle and simulate
the atmospheric concentrations and deposition fluxes of P over land and
oceans, driven by mineral, natural and combustion P emissions. To our
knowledge, this is the first study that accounts for both inorganic and
organic forms of P and their evolution in the atmosphere. Furthermore, we
also present the first global estimate of the PO4 flux due to the
acid solubilization of dust particles. The model description and the
parameterization of atmospheric acidity impact on the P solubilization from
mineral dust aerosol in atmospheric water, together with the OP atmospheric
ageing contribution to the DP global budgets are presented in Sect. 2. The
calculated TP and DP global atmospheric concentrations are shown and
compared to observations in Sect. 3. In Sect. 4, the importance of present-day air pollutants on DP atmospheric deposition is investigated based on
simulations using past and future anthropogenic and biomass burning emission
scenarios. The contribution of bioaerosols to the bioavailable P atmospheric
deposition and implications of the findings concerning the biogeochemistry
of marine ecosystems are also discussed (Sect. 4). Overall, the impacts of
human-driven changes on the calculated DP deposition fluxes to the global
ocean are summarized in Sect. 5.
Model description
The TM4-ECPL global chemistry transport model
(Myriokefalitakis et al., 2015) simulates the oxidant chemistry accounting for non-methane
volatile organics and all major aerosol components, including secondary
inorganic aerosols like sulfate (SO42-), nitrate
(NO3-) and ammonium (NH4+) calculated using the ISORROPIA II
thermodynamic model (Fountoukis and Nenes, 2007) and secondary
organic aerosols (Tsigaridis and Kanakidou, 2007; Tsigaridis et al., 2014). The atmospheric cycles of Fe
and N in TM4-ECPL have been parameterized and evaluated in
Myriokefalitakis et al. (2015) and Kanakidou et al. (2016), respectively, while uncertainties in the computed atmospheric
composition associated with different emission parameterizations have been
calculated in Daskalakis et al. (2015). The
model's ability to reproduce distributions of organic aerosols
(Tsigaridis et al., 2014) and tropospheric ozone, ozone's precursors and aerosols have
been also evaluated against satellite and in situ observations
(Eckhardt et al., 2015; Stohl et al., 2015; Quennehen et al., 2016; Myriokefalitakis et al., 2016).
TM4-ECPL is driven by the ECMWF (European Centre for Medium-Range Weather
Forecasts) Interim Reanalysis project (ERA-Interim) meteorology
(Dee et al., 2011). The current model configuration has a horizontal resolution
of 3∘ in longitude by 2∘ in latitude and 34 hybrid layers in the
vertical, from the surface up to 0.1 hPa, with a model time step of 30 min.
TM4-ECPL uses modal size (lognormal) distributions to describe the evolution
of fine and coarse aerosols in the atmosphere. In the model, phosphorus
is parameterized using 32 P-containing aerosol tracers of different sizes and solubilities. In TM4-ECPL, different sources emit
P-containing aerosols of different sizes represented by lognormal
distributions as outlined in Sect. 2.1. For each aerosol mode and source
(Fig. 1), the model accounts for total P, phosphate, insoluble and soluble
OP. For the dust source, it also accounts for the two P-containing minerals
(fluorapatite and hydroxyapatite) as further described in Sect. 2.1.1. These are individually transported, aged and deposited in the
atmosphere. The “dry” aerosol hygroscopic growth in the model is treated as
a function of ambient relative humidity and the composition of soluble
aerosol components based on experimental work by Gerber (1985) and this uptake of water on aerosols changes the particle size.
In addition, during atmospheric transport, there are major changes in the
size distribution of aerosols as a consequence of the removal of larger
particles due to gravitational settling. The P-containing aerosols follow
the same parameterizations, hygroscopic growth and removal processes that affect the mass median radius (i.e., size).
TM4-ECPL uses anthropogenic (including ship and aircraft) emissions and
biomass burning emissions from the historical Atmospheric Chemistry and
Climate Model Intercomparison Project (ACCMIP) database
(Lamarque et al.,
2013) for the years 1850 (hereafter PAST), 1999 and 2000 from the
Representative Concentration Pathways 6.0 (RCP 6.0) emission scenario
(van Vuuren et al., 2011) for the
years 2001 to 2010 (year 2008 is hereafter called PRESENT) and for the year
2100 (hereafter FUTURE) that have been used for the sensitivity simulations.
Details on anthropogenic and natural emissions used for this work are
provided in Myriokefalitakis et al. (2015) with the exception of mineral dust that, for the present study,
is calculated online as in van Noije et al. (2014),
based on the dust source parameterization of Tegen et al. (2002). The three base simulations (PAST, PRESENT and FUTURE) have been
performed with meteorology for the year 2008. Note, however, that we have
extended the present-day simulation to the 11-year period from 2000 to 2010
with a spin-up time of 1 year (i.e., with 1999 meteorology and emissions),
to cover the majority of the dates of the available atmospheric observations
used for model evaluation (see Sects. 2.4 and 3.2).
Phosphorus emissions
Phosphorus emissions from mineral dust
Apatite is the most abundant primary natural source of P in soils
(Newman, 1995), compared to other low solubility P
forms such as secondary metal phosphate precipitates and organic phosphate.
For the present study, apatite is assumed to be the only mineral in dust
that contains P. The spatially distributed fraction of P in soils
(fP) from the global soil mineralogy dataset developed by
Nickovic et al. (2012) is used to calculate the
inorganic P-containing mineral (i.e., apatite) emissions as
EP=F880⋅fP⋅EDu,
where EDu is the online calculated dust emissions in the model,
F880 is a factor applied to adjust the P emissions to the global mean P
content of mineral dust in the model domain of 880 ppm per weight as
observed by Zamora et al. (2013) and EP are
the resulting inorganic P emissions from mineral dust. P-containing minerals
associated with dust particles are emitted in the fine and coarse mode with
mass median radii (lognormal standard deviation) of 0.34 (1.59) and
1.75 µm (2.00), respectively. The P-containing dust aerosol emissions
are treated as a lognormal distribution with a dry mass median radius and sigma
same as that of dust particles. Particle sizes are changing due to
hygroscopic growth as a function of ambient relative humidity and the
composition of soluble aerosol components (Gerber, 1985).
Note, however, that no coagulation among different dust modes is considered
for the current study.
Emissions of TP and DP (in Tg-P yr-1) taken into account in
the TM4-ECPL model for PAST, PRESENT and FUTURE simulations. In parenthesis,
the average values of TP and DP emissions for the years 2000–2010 are also
provided.
TM4-ECPL
Biomass
Anthropogenic
Volcanoes
PBAPs
Sea spray
Soils
Total
burning
combustion
TP
PAST
0.014
0.008
1.289
PRESENT
0.018
0.043
0.006
0.156
0.008
1.097
1.328
(0.018)
(0.042)
(0.006)
(0.156)
(0.008)
(1.095)
(1.326)
FUTURE
0.022
0.009
1.298
DP
PAST
0.007
0.004
0.254
PRESENT
0.009
0.021
0.006
0.123
0.008
0.106
0.272
(0.009)
(0.021)
(0.006)
(0.123)
(0.007)
(0.105)
(0.271)
FUTURE
0.011
0.004
0.258
Although in most relevant modeling studies airborne P-containing dust
particle emissions are assumed to have an average P content of 720 ppm
(Mahowald et al., 2008; Wang et al., 2014; Brahney et al., 2015) in the
atmosphere due to transport, ageing and deposition processes; the overall mineralogy may
change the chemical composition and size of the dust aerosol population. In a
recent iron modeling study (Perlwitz et al., 2015), however, a significant
effort has been made to model the mineral composition of dust considering
the differences from the original soil composition.
Perlwitz et al. (2015) have found a significant overestimate (a factor of 10–30) mainly in the fine aerosol
emissions that are the smallest part of dust emissions (e.g., about 7 % of
the total emissions in our model) and an underestimate in the larger
particle emissions both for total dust and for individual minerals when the
mineralogy of dust aerosol is assumed to be the same as that of the soil.
However, for the present study, we did not account for different P content
for dust particles in the fine and the coarse mode, since the global soil
mineralogy dataset used by Nickovic et al. (2012) does
not provide any information of P content in silt and clay soil particles
separately. Note also that recent studies indicate that dust super-coarse
particles can be very important for the biogeochemistry over land, since
they can represent the dominant fraction of dust close to source regions
(Lawrence and Neff, 2009; Neff et al., 2013). The Brahney et al. (2015)
modeling study that focused on the atmospheric phosphorus deposition over
global alpine lakes, based on Neff et al. (2013) observations, estimated that only 10 % of the mass that travels in
the atmosphere is within the < 10 µm size fraction. In our
study, we do not account for super-coarse dust particles because due to their
short atmospheric lifetime, they are emitted and deposited in the same model
grid box (Brahney et al., 2015).
This omission is not expected to have a significant impact on our results,
since the present work is focused on the P-solubilization mechanisms
occurring via atmospheric long-transport mixing and on the bioavailable P
deposition over the marine environment.
For the year 2008, the mineral dust emissions calculated in TM4-ECPL amount
to 1181 Tg yr-1 and the corresponding apatite emissions to 1.034 Tg-P yr-1 with 10 % of it (0.103 Tg-P yr-1) in the soluble form
(Table 1). The soluble fraction used in our model is based on the
measurements of leachable inorganic phosphorus (LIP) for Saharan soil dust,
as presented by Nenes et al. (2011). These
authors found that LIP represented up to 10 % of total inorganic P in
Saharan soil samples and dry fallout collected during Sahara dust storms
before acid treatment. Moreover, Yang et al. (2013)
estimated the labile inorganic P in the top soil on the global scale at
about 3.6 Pg-P, which corresponds to about 10 % of the estimates of total
soil P on the global scale 30.6–40.6 Pg-P
(Smil, 2000; Wang et al., 2010; Yang et al., 2013). To further investigate uncertainties
associated with the soluble fraction of P-containing dust aerosol emissions
in our model, an additional simulation has been performed neglecting any
soluble fraction on initial emissions.
In addition to the desert dust inorganic P source, we account for the OP
present in soil's organic matter, following the method developed by
Kanakidou et al. (2012, and
references therein). Thus, using a mean OP : OC molar ratio of 0.005, a mean
OM content of soil dust of 0.25 % and an OM : OC molar ratio of 1.76, we
evaluate the dust source of OP here at 0.022 Tg-P yr-1 for the year
2008. This flux is in good agreement with the 0.03 Tg-P yr-1 calculated
for 2005 by Kanakidou et al. (2012) using the same methodology but with the
AEROCOM database for dust emission fluxes
(Dentener et al., 2006). Note that similarly to that earlier study, a solubility of 10 % is
applied here to the OP dust emissions.
Phosphorus emissions from combustion sources
For the present study, the P / BC mass ratios of combustion sources as
estimated by Mahowald et al. (2008; i.e., 0.0029 for fine aerosols and 0.02 for coarse aerosols) are
applied to the inventories of monthly BC emissions of anthropogenic (i.e., for
fossil fuel, coal, waste and biofuel) and biomass burning origin, as
provided by the historical ACCMIP database for 1850 and from the RCP 6.0 for
2008 and 2100. In the model, a number mode radius of 0.04 µm and a
lognormal standard deviation of 1.8 are assumed for fine P emissions, while
for coarse P a number mode radius of 0.5 µm and lognormal standard
deviation of 2.00 are used as proposed for combustion aerosols by
Dentener et al. (2006). BC emissions from anthropogenic combustion in the coarse mode are
assumed to be 25 % of those in the fine mode (Jacobson and
Streets, 2009), while biomass burning emissions in the coarse mode are
assumed equal to 20 % of those of fine aerosols
(Mahowald et al., 2008). Thus,
the computed anthropogenic combustion and biomass burning annual mean
sources of TP are calculated to be 0.043 (by about 70 % in
the coarse mode) and 0.018 Tg-P yr-1 (by about 66 % in the coarse
mode), respectively, all corresponding to the year 2008. Despite the
different emission databases and the aerosol size parameterization, the
computed present-day TP sources for the year 2008 are comparable to those of
Mahowald et al. (2008) for the
year 2000 (i.e., 0.045 and 0.025 Tg-P yr-1 for
anthropogenic combustion and biomass burning, respectively). PAST, PRESENT
and FUTURE combustion emissions calculated for this study based on the
ACCMIP and RCP 6.0 database are presented in Table 1.
Half of TP emissions from combustion sources are considered to be in the
form of OP following the approach of
Kanakidou et al. (2012). All P-containing particles from combustion emissions are initially treated here
as 50 % soluble (Mahowald et al., 2008). The insoluble fraction of OP associated with combustion
emissions can be further converted to soluble OP (DOP) during atmospheric
ageing, using the ageing parameterization for primary hydrophobic organic
aerosols in the model (Tsigaridis and
Kanakidou, 2003; Tsigaridis et al., 2006), but for the respective size and
lognormal distribution of OP aerosols (with the larger particles experiencing
the smallest conversion rates).
To further investigate uncertainties in the P combustion emissions in our
model, an additional present-day simulation was performed taking into
account the total (bulk) mass of anthropogenic combustion and biomass
burning P emissions, as developed by Wang et al. (2014; R. Wang, personal communication, 2016). According to that
database, global anthropogenic emissions from fossil fuels, biofuels and
deforestation fires amount to 1.079 Tg-P yr-1 and natural fire
emissions are equal to 0.808 Tg-P yr-1. For this sensitivity simulation, we
apply the size distribution as described in
Wang et al. (2014); i.e., by dividing total
emissions into three modes – one fine (2 % of P) and two coarse modes (25
and 73 % of P) – with mass mode dry diameters of 0.14, 2.5
and 10 µm and lognormal standard deviations of 1.59 and 2.00 for fine
and coarse modes, respectively.
Phosphorus emissions from primary biological aerosol
particles
Three types of P-containing PBAPs are considered for the present study:
bacteria (BCT), fungal spores (FNG) and pollen grains (PLN). PBAPs from
other sources, such as insect fragments and plant debris
(e.g., Després et al., 2012), are however neglected in the present study.
Omission of these super-coarse particles is expected to lead to an
underestimate in the PBAP contribution to P deposition over land that
requires evaluation with targeted observations. The BCT fluxes are
parameterized based on the Burrows
et al. (2009b) best-fit estimates for particles of 1 µm diameter flux
rates (f) and for six different ecosystems: coastal: 900;
crops: 704; grassland: 648; land ice:
7.7; shrubs: 502; and wetlands: 196 m-2 s-1. For the present study, the Olson Global Ecosystem
Database (Olson, 1992), originally available for
74 different land types on a spatial scale of 0.5∘ × 0.5∘, is lumped into 10 ecosystem groups as proposed by
Burrows et al. (2009). The total
BCT flux (FBCT; s-1) in the model is calculated based on the
aforementioned fluxes (fi; m-2 s-1) per ecosystem (i), weighted
by the respective ecosystem area fraction in the model grid box (ai; m2) as
FBCT=∑i=16ai⋅fi.
Heald and Spracklen (2009) proposed that FNG fluxes
linearly depend on the leaf area index (LAI; m2 m-2) and the
specific humidity (q; kg kg-1), based on near-surface mannitol
observations. For the present study, however, we use a recently published
emission parameterization proposed by
Hummel et al. (2015), as derived based on fluorescent biological aerosol particles field
measurements at various locations across Europe and for spores with a mean
dry diameter of 3 µm (Eq. 3):
FFNG=20.426⋅T-275.82K+3.93⋅104⋅q⋅LAI.
In the TM4-ECPL, that parameterization (Eq. 3) is used to calculate FNG
emissions online, using monthly averaged LAI distributions and 3-hourly
averaged specific humidity (q) and temperature (T) data, as provided by
ERA-Interim.
PLN emissions maximize when plant surfaces are dry and are under high turbulence
during the morning hours and during spring months (Jacobson
and Streets, 2009). Hoose et al. (2010)
parameterized the pollen flux rate as linearly dependent on LAI, assuming
particles with a mean dry diameter of 30 µm, by simplifying the more
sophisticated parameterization developed by Jacobson and
Streets (2009) for a global model. Here, we use the
Jacobson and Streets (2009) pollen parameterization
(particle mean dry diameter of 30 µm) with the pollen flux
(FPLN; s-1) calculated by the following equation:
FPLN=fPLN⋅LAI⋅Rmonth⋅Rhour,
where fPLN= 0.5 m-2s-1, the factor Rmonth accounts
for the seasonal and Rhour the hourly pollen flux variation.
PBAPs are assumed here to be monodisperse spherical particles
(Hoose et al., 2010; Hummel et al., 2015) of 1 g cm-3 density
(Sesartic and Dallafior, 2011) with an organic matter to
organic carbon (OM : OC) ratio set equal to 2.6 (i.e., that of mannitol)
corresponding to a molecular weight equal to 31 g mol-1, as suggested
by Heald and Spracklen (2009). According to our model
estimates, roughly 60 Tg-C yr-1 are emitted as PBAPs. Bacterial emissions
are assumed to be completely insoluble (Ariya et al., 2009), fungal spores are emitted as 50 % soluble aerosols
(Mahowald et al., 2008; Kanakidou et al., 2012), while pollen is emitted as totally
soluble aerosols (Hoose et al., 2010). A constant mean P : C atomic ratio of 0.001 is used for PBAPs, as suggested by
Kanakidou et al. (2012) and all P is assumed in the form of OP. Based on the above parameterizations, the
model calculates an OP emission flux associated with a PBAP value equal to 0.156 Tg-P yr-1, of
which 0.123 Tg-P yr-1 (about 80 %) is considered
to be in the form of DOP (Table 1). However, because PBAPs consist of
biological material, they are considered here to be bioavailable for marine
ecosystems, as further discussed in Sect. 4.1 and 4.2. In addition, in
TM4-ECPL, upon emission, the insoluble fraction of PBAPs becomes
progressively soluble due to atmospheric ageing. This process that has been
seen to occur, for instance, by degradation of RNA
(Paytan et al., 2003), in TM4-ECPL is parameterized based on oxidant levels as for all organic
aerosols (Tsigaridis and Kanakidou, 2003; Tsigaridis et al., 2006).
Phosphorus emissions from sea spray
Oceanic P emissions associated with sea spray are computed online here
based on a sea-salt emission flux parameterization of
Vignati et al. (2010), accounting for fine and coarse modes, with number mode dry
radii of 0.09 and 0.794 µm, and lognormal standard deviations
of 1.59 and 2.00 for accumulation and coarse particles, respectively.
Sea-spray emissions are driven by the model's meteorology, and for the year
2008, the model calculates a total of about 8284 Tg yr-1 of sea-salt
emissions (of which 41 Tg yr-1 are in the fine mode). These numbers
compare well with the AEROCOM recommendation of 7925 Tg yr-1 by
Dentener et al. (2006) and are within the range of 2272–12 462 Tg yr-1 computed by
Tsigaridis et al. (2013) using several
different parameterizations. Note that our sea-salt source estimation is,
however, much lower than the one used in the modeling study by
Wang et al. (2014; i.e., 25 300 Tg yr-1),
since super-coarse sea-salt particles are not considered in the current
parameterization.
The oceanic P emissions in TM4-ECPL are calculated as
EPO4=[P]/MWP[Na]/MWNa⋅ENa,
where [P] is the P seawater concentrations in µM, [Na] is Na seawater
concentration in µM and ENa is the sea-salt emission flux from the
ocean surface in kg-Na m-2 s-1. MW is the corresponding molecular
weight of P and Na, used to convert molar to mass ratios. In TM4-ECPL,
sea-salt particles are emitted from the ocean's surface every time step
using surface wind-speed data from the ERA-Interim database (updated every
3 h). Surface seawater PO4 concentrations come from the LEVITUS94
World Ocean Atlas (Conkright et al., 1994;
http://iridl.ldeo.columbia.edu/SOURCES/.LEVITUS94/.ANNUAL/.PO4/) ranging
up to about 3 µM of PO4 in the global ocean. Taking into
account that the average Na concentration in seawater is about 10.781 g-Na kg-water-1, as well as an average seawater salinity
of 35.5 g kg-water-1, the spatial distribution of surface oceanic Na
concentrations can be derived from the distribution of the surface salinity
concentrations as provided by the LEVITUS94 World Ocean Atlas
(Levitus et al., 1994; http://iridl.ldeo.columbia.edu/SOURCES/.LEVITUS94/.ANNUAL/.sal/). Note
that surface concentrations, both for seawater PO4 and salinity,
correspond to the data available for 0 m depth (with the next available depth
in the LEVITUS94 database at 10 m).
We additionally take into account the OP oceanic emissions, as described in
Kanakidou et al. (2012; see Supplement and references
therein). For this, the model accounts for a mean seawater
OP concentration of 0.2 µM of P, based on Björkman and
Karl (2003) observations. Since, to our knowledge, no spatial distribution of
seawater OP concentrations is available, the monthly mean surface chlorophyll
a (Chl a) concentrations from MODIS retrievals, used in the model to derive
marine primary organic aerosol emissions (Myriokefalitakis et al., 2010), are
used as a proxy to geographically distribute the mean seawater OP
concentrations. Overall, the model calculates an emission flux of TP equal to
0.008 Tg-P yr-1 from the global ocean (Table 1), of which
0.001 Tg-P yr-1 is in the form of OP. Note that the insoluble fraction
of oceanic OP in the model can be transferred to the soluble mode (DOP) due
to atmospheric ageing processes. The omission of the super-coarse sea-salt
aerosol might affect our estimates of P deposition to the ocean. Brahney et
al. (2015) evaluated this source at 0.0046 Tg-P yr-1, an amount that
introduces a 3 % underestimate to the present-day P
deposition flux to the oceans calculated here.
Phosphorus emissions from volcanic aerosols
Mahowald et al. (2008)
estimated that about 0.006 Tg-P yr-1 is associated with volcanic
aerosols on a global scale, based on volcanic plume observations. Although,
on a global scale, volcanic ash is a small source of TP, it is found to
impact, at least regionally, the ocean nutrient distributions and marine
productivity (Uematsu et al.,
2004; Henson et al., 2013; Olgun et al., 2013). For the present study, we
applied that global annual mean volcanic flux (see also Table 1), using the distribution of sulfur volcanic emissions by
Andres and Kasgnoc (1998), as updated by
Dentener et al. (2006). Volcanic phosphorus is assumed here to reside in the fine
particulate mode and is treated in the model as totally soluble aerosol
(i.e., DIP), as proposed by
Mahowald et al. (2008). The
lognormal size-distribution parameters used for volcanic P aerosol have a
number mode radius of 0.04 µm and a lognormal standard deviation of 1.8
as for sulfate fine aerosols from continuous volcanic eruptions (Dentener et al., 2006).
Phosphorus acid-solubilization mechanism
Phosphorus solubilization from mineral dust under acidic atmospheric
conditions is assumed here to occur for the least- and the most-soluble
member of apatite minerals as proposed by
Nenes et al. (2011): the fluorapatite
(Ca5(PO4)3(F); hereafter FAP) and the hydroxyapatite
(Ca5(PO4)3(OH); hereafter HAP), respectively. FAP is
considered as a geologically abundant apatite, usually present in the form
of igneous or sedimentary carbonate FAP
(Guidry and Mackenzie,
2003). However, due to lack of information on the relative abundance and
geographic distribution of FAP and HAP in soils, we assume equal mass
fractions of FAP and HAP in apatite-containing soils.
The dissolution of FAP and HAP here is treated as a kinetic process, the
rate of which depends on the H+ activity of atmospheric water (i.e.,
aerosol water and cloud droplets), the reactivity of P species, the ambient
temperature and the degree of solution saturation. For aerosol water, the
activity of H+ is calculated online in the model by the thermodynamic
module ISORROPIA II (Fountoukis and Nenes, 2007). For cloud
water, the H+ concentration is calculated by the aqueous-phase
chemistry module as presented in
Myriokefalitakis
et al. (2011, 2015). The phosphate dissolution rate (R), as moles of
HPO4-2 per second per gram of apatite, is obtained using the
empirical formulation of Lasaga et al. (1994):
R=K(T)⋅a(H+)m⋅f⋅A,
where K is the reaction constant in moles m-2 s-1, a(H+) is the
H+ activity, m is the experimentally derived reaction order with respect
to the solution H+ concentration and A is the specific surface area of
each apatite-containing particle in m2 g-1. The function f
(Cama et al., 1999) depends on the solution
saturation state (0 ≤f≤ 1) and is given by
f=1-Q/KEq,
where Q is the reaction activity quotient, KEq is each apatite
equilibrium constant and Q/Keq is the fraction that expresses the state of
saturation of the solution (with respect to the apatite), calculated every
time step in the model. Thus, when f= 1, the solution is far from
equilibrium; therefore, the dissolution rate becomes maximum, while as f approaches
0, the solution approaches equilibrium with any remaining undissolved FAP
and HAP.
Fluorapatite (FAP) acid dissolution constants used for this study.
Minerale
pH
K (T)
m
AMIN
Keq
(mol m-2 s-1)
(m2 g-1)
FAP
< 5.5
5.75 × 10-6exp[4.1 × 103(1/298-1/T)]a
0.81a
10.7b
10-23.12d
(FAP)
(FAP)
5.5–6.5
6.91 × 10-8exp[4.1 × 103(1/298-1/T)]a
0.67a
> 6.5
6.53 × 10-11exp[4.1 × 103(1/298-1/T)]a
0.01a
80.5c
10-20.47d
(HAP)
(HAP)
a Guidry and Mackenzie (2003). b Bengtsson et al. (2007). c Bengtsson et al. (2009).
d van Cappellen and Berner (1991).
e For HAP dissolution constants, we assume those of FAP, as adopted from Guidry and Mackenzie (2003) and
corrected based on Palandri and Kharaka (2004) reviewed data (see Sect. 2.2).
HAP is experimentally found to be roughly 3 orders of magnitude more soluble
than FAP (KEq(HAP) = 10-20.47 vs. KEq(FAP) = 10-23.12), as reported by Nenes et al. (2011) based on van
Cappellen and Berner (1991). According to the compilation of experimental
determinations of P-dissolution rates of HAP and FAP by
Palandri and Kharaka (2004), the dissolution
rate of HAP is found to be about an order of magnitude slower than that of
FAP under highly acidic conditions (K(HAP) = 10-4.29 and K(FAP) = 10-3.73 for pH = 0), while under neutral conditions, HAP is found to
dissolve 2 orders of magnitude faster than FAP (K(HAP) = 10-6 and
K(FAP) = 10-8 for pH = 7). Moreover, HAP is measured to have an almost 8
times larger specific surface area (80.5 m2 g-1,
Bengtsson et al., 2009) compared to that of FAP (10.7 m2 g-1),
which is in agreement with the measured specific surface areas of 8.1–16 m2 g-1 for sedimentary FAP
(Guidry and Mackenzie, 2003). Guidry and
Mackenzie (2003) have experimentally derived different rate constants (K)
for FAP dissolution ranging from 5.75 × 10-6 to
6.53 × 10-11 mol m-2 s-1, with a pH ranging from 2 to 8.5.
They further derived the respective reaction orders (m) for each pH region,
between 0.01 (for neutral to basic conditions) and 0.81 (for acidic
conditions), while the activation energy of the FAP dissolution (Ea)
was calculated as equal to 8.3 kcal mol-1. For the present study, the
dissolution reaction coefficient K for FAP (Table 2) is based on the
dissolution experiments by Guidry and Mackenzie (2003) for a range of pH values (2–12), temperatures (25–55 ∘C) as well
as for various solution saturation states and ionic strengths.
Bengtsson et al. (2009) have experimentally studied the solubility and the surface
complexation of non-stoichiometric synthetic HAP, identifying three distinct
pH regions for their batch dissolution experiments: (1) under acidic pH
(< 4.5), where HAP dissolution is relatively high, producing high
concentrations of Ca2+ and H2PO4-; (2) under basic pH
(> 8.2), where surface complexation is the main process; and
(3) under intermediate pH (4.5–8.2), where both dissolution and surface complexation
occur. However, they do not provide sufficient information to enable
parameterizing HAP dissolution similarly to FAP dissolution. Therefore, for
HAP dissolution kinetics, we use the dissolution rates of FAP after
correcting them to account for the differences between HAP and FAP
dissolution kinetics as a function of pH and T, as reported by
Palandri and Kharaka (2004). For this, we
consider the different dissolution rates for a pH range of 0 to 7–8, which
is the range of acidity encountered by atmospheric particles, including dust
(e.g., Bougiatioti et al., 2016; Weber et al., 2016). At the strongly acidic limit
(25 ∘C and pH = 0), the dissolution rate of HAP is assumed here
to be about 27 % (i.e., 10-0.56 times) slower than that of FAP,
but for neutral and basic conditions (and 25 ∘C) HAP dissolves 2
orders of magnitude faster than FAP (Palandri
and Kharaka, 2004). The dissolution rate also changes with temperature; we
assume that HAP dissolution has a similar activation energy to FAP
(Palandri and Kharaka, 2004; Guidry and
Mackenzie, 2003). Additional details for the FAP and HAP mineral dissolution
rate parameters are presented in Table 2.
Observation data for model evaluation
The evaluation of the global atmospheric P cycle for the present study has
been performed based on available observations of aerosol concentrations
(Table S1 in the Supplement) and deposition fluxes (Table S2) from various locations around
the globe (cruises and land-based stations). The methodological details of
the observations used for this study are well documented in the literature
and thus are not reviewed here in detail. For DP concentrations in ambient
aerosols, we compiled cruise observations of PO4 over the Atlantic
Ocean (50∘ N–50∘ S) from Baker et al. (2010), over the western Pacific
(25∘ N–20∘ S) from Martino et al. (2014) and over the eastern
tropical North Atlantic Ocean (58∘ S–35∘ N, 14–38∘ W) from
Powell et al. (2015). For these oceanic cruise observations, samples were either collected and separated into
fine-mode (aerodynamic particle diameter < 1 µm) and coarse-mode
(1 µm < aerodynamic particle diameter) particles using cascade
impactors that may include or exclude particles with diameters larger than
10 µm, or using a single bulk filter. We additionally use average PO4
concentrations (aerodynamic particle diameter < 10 µm) from
cruise measurements over Bay of Bengal and the Arabian Sea
(Srinivas and Sarin, 2012, 2013, 2015). Finally, we also took into account land-based TP and PO4
aerosol concentration measurements from two sites in the Mediterranean: (i) the Finokalia monitoring station (35∘20′ N, 25∘40′ E)
located in the eastern Mediterranean (Crete, Greece) and (ii) Ostriconi
(42∘40′ N, 09∘04′ E) located in the western Mediterranean
(Corsica, France). The samples at both sites were collected either
separating the fine (aerodynamic particle diameter < 1.3 µm) and coarse mode (10 µm > aerodynamic particle
diameter > 1.3 µm; Koulouri et al.,
2008; Mihalopoulos and co-workers, unpublished data) or as bulk
(Markaki et al., 2010). Details about the characteristics of these Mediterranean sampling
sites can be found in
Markaki et al. (2010), while the methodology for aerosol sampling and analysis is described
in detail in
Koulouri et al. (2008).
Although P deposition flux data are rather limited on a global scale, for
the present study, we use the wet and dry deposition fluxes (both for TP and
DP) compiled by Vet et al. (2014; R. Vet, personal communication, 2016). For wet deposition of
DP, we use available filtered (i.e., analyzed as orthophosphates with no
digestion as DIP) and unfiltered (i.e., analyzed as orthophosphates following
digestion as total DP) annual measurements (Fig. 8.2 in
Vet et al., 2014).
For the TP wet deposition measurements, we use annual wet deposition
measurements (Fig. 8.3 in Vet et al., 2014) of
unfiltered samples. The compilation of the phosphorus dry deposition fluxes
by Vet et al. (2014) is based on airborne phosphorus (TP and PO4) concentrations from
around the world and gridded annual dry deposition velocities from the
Mahowald et al. (2008)
modeling study (Fig. 8.6 and 8.7 in Vet et al., 2014).
The size distribution used in these dry deposition calculations is the same
as in the modeling study by
Mahowald et al. (2008); thus,
the derived dry deposition fluxes account for particles with diameter up to
10 µm. Finally, we also take into account DP wet and dry deposition
observations from the Finokalia station in the eastern Mediterranean
(Markaki et al.,
2010; Mihalopoulos and co-workers, unpublished data), based on rain water
samplings (wet-only collector) and glass-bead devices, respectively. Further
details on the methodology of the deposition measurements at Finokalia can
be found in
Markaki et al. (2010).
Secondary DP sources (in Tg-P yr-1) due to OP ageing contained
in biomass burning, anthropogenic combustion, sea spray and dust as well as
due to dust (apatite) dissolution via the acid-solubilization mechanism, as
calculated by the TM4-ECPL model for PAST, PRESENT and FUTURE simulations.
Dust ageing
TM4-ECPL
Biomass burning
Anthropogenic
PBAPs
Sea spray
OP
Apatite
ageing
combustion ageing
ageing
ageing
ageing
dissolution
DP
PAST
0.002
0.001
0.016
0.0001
0.007
0.085
PRESENT
0.003
0.005
0.016
0.0001
0.008
0.144
FUTURE
0.003
0.001
0.016
0.0001
0.008
0.100
Annual averaged column distributions (in ng-P m-2 s-1)
of the (a) TP emissions, (b) DP emissions, (c) DIP flux from apatite
dissolution and (d) DOP production due to OP atmospheric ageing, as
calculated by the TM4-ECPL model for the present atmosphere (year 2008).
Results and discussion
Sources of atmospheric phosphorus
Figure 2 presents the annual mean primary TP and DP emissions from the
various sources taken into account in the model (the
emission distribution per source for TP and DP are also presented in Figs. S1
and S2 in the Supplement, respectively). TP emissions (Fig. 2a) maximize over the major
deserts of the world (e.g., Sahara, Gobi, Arabian, Kalahari, North American
and Australian deserts) with simulated P fluxes up to 100 ng-P m-2 s-1 (Figs. 2a and S1a). Secondary maxima of TP emission fluxes of
about 0.1–1 ng-P m-2 s-1 are also calculated over the
midlatitudes of the Northern Hemisphere (NH), such as in China, Europe and the
USA, due to release of TP to the atmosphere in ash produced during combustion
processes of anthropogenic origin (Fig. S1b) and over forested areas in
equatorial America. Additionally, during biomass burning episodes, TP is
further released to the atmosphere (Fig. S1c), however, at rates about 1
order of magnitude lower than those of combustion of anthropogenic origin
(roughly 0.01 ng-P m-2 s-1).
The same pattern (as for TP emissions) is simulated for the P-soluble
fraction (Fig. 2b), but with lower emission fluxes (e.g., about 1 ng-P m-2 s-1 over the Sahara). This is attributed to the
solubility of P-containing mineral dust at emission that corresponds to the
DP present in the desert soil due to weathering. As discussed in Sect. 2.1.2, this fraction is taken equal to 10 % for the present study.
Associated mineral DP emissions (Fig. S2a) of 0.106 Tg-P yr-1 (as PO4
and/or DOP) occur mainly over the Saharan desert region, but significant
fluxes are also calculated to occur over other important deserts of the
globe. Anthropogenic DP emissions (0.021 Tg P yr-1) occur mainly over
densely populated regions of the globe (e.g., the midlatitudes of the NH
such as China, Europe and the USA), with simulated fluxes up to 0.1 ng-P m-2 s-1 (Fig. S2b). DP emissions from biomass burning contribute
about 0.009 Tg-P yr-1, peaking over intense biomass burning areas,
e.g., tropical and high-latitude forests and showing maxima over central Africa,
Indonesia and Amazonia (Fig. S2c).
DIP annual fluxes (in ng-P m-2 s-1) from apatite
dissolution (a) in aerosol water and (b) in cloud droplets, as calculated by
the TM4-ECPL model for the present atmosphere (year 2008).
The present-day annual apatite dissolution flux is calculated as equal to 0.144 Tg-P yr-1 (Table 3, Fig. 2c). Most of the apatite dissolution fluxes
occur downwind of the major dust source regions (i.e., Nigeria downwind of
the Sahara, Pakistan downwind of the Thar Desert and China downwind
of the Gobi Desert). Over these regions, the long- and regional-range
transport of natural and anthropogenic pollutants enhance atmospheric
acidity, and subsequently, P is mobilized from mineral apatite. The model
calculates maximum dissolution fluxes downwind of the Sahara and the Gobi
Desert, over the Persian Gulf, the whole Middle East and the Mediterranean
basin as well as over the equatorial Atlantic. In addition, enhanced apatite
dissolution is calculated over the tropical Atlantic Ocean, India and the
outflow of Asia to the Pacific Ocean, in line with observations of changes
in solubility during transport of dust across the tropical Atlantic Ocean by
Baker et al. (2006a).
As explained in Sect. 2, for the present study the apatite dissolution (Fig.
2c) is due to the respective FAP and HAP solubilizations that occur both in
aerosol water and cloud droplets (Fig. S3). The model calculates that most
of the apatite dissolution (0.111 Tg-P yr-1) is occurring in
deliquesced particles (Figs. 3, S3a and b), mainly attributed to the higher
aerosol acidity, while only 0.034 Tg-P yr-1 is calculated to occur in
cloud droplets (Figs. 3b, S3c and d). Note that the model-calculated
global mean pH in clouds is about 4.5
(Myriokefalitakis
et al., 2015). In addition, the distributions of aerosol and cloud
dissolution of apatite are rather different (Fig. 3a,b). In-cloud
dissolution is calculated to maximize (i) offshore from the African continent
(i.e., over Cote d'Ivoire, Nigeria and Cameroon) over the equatorial Atlantic
Ocean and (ii) over China and India, where dust aerosols downwind of major
desert regions (i.e., the Sahara and the Gobi Desert, respectively) meet polluted and
acidic cloud droplets, while dissolution in aerosol water also shows high
rates over the USA, Europe and Saudi Arabia.
Location of observational data for (a) concentrations of
P-containing aerosols (bulk, fine and coarse) and (b) deposition fluxes (wet
and dry deposition), (c–f) log scatterplots between model (y axis) and all
observations (x axis) for surface (c) PO4 and (e) TP aerosol concentrations
(µg-P m-3) measured in cruises (blue dots) and stations (red
stars), as well as for (d) PO4 and (f) TP dry deposition fluxes (mg-P m-2 s-1)
over oceans (blue dots) and inland sites (red stars). The
continuous black line shows the 1:1 correlation and the dashed black lines
show the 10:1 and 1:10 relationships, respectively.
DIP fluxes from HAP dissolution in the cloud droplets (Fig. S3d) are
calculated to be roughly 60 % higher than those of FAP (Fig. S3c; 0.021 against 0.013 Tg-P yr-1). However, for the DIP
dissolution fluxes from the FAP and the HAP in aerosol water, no differences
are calculated (Fig. S3a and b) for the more acidic environmental in the fine
aerosol water (0.015 Tg-P yr-1 for each of them). On the
contrary, the HAP is more soluble than FAP in the less acidic coarse aerosol
water (0.041 Tg-P yr-1 for the HAP compared to 0.039 Tg-P yr-1
for the FAP; see also Fig. S2 in the Supplement of Myriokefalitakis
et al., 2015 for pH calculations in the model). The
changes in the saturation factor (f) in the aerosol water are also of
importance. Under conditions where HAP is more soluble than FAP, the respective
mobilized PO4 concentrations increase faster in the aerosol solution and
react with the soluble Ca2+ present in dust, ultimately forming
amorphous apatite that precipitates from the solution (i.e., f= 1; thus, the
dissolution process stops). In the presence of soluble Ca2+ and PO4,
other salts, such as monenite (CaHPO4; Somasundaran et al., 1985), can also be formed
and further impact the solution's degree of saturation. These results
suggest that the solution saturation effect in dust aerosol water can be a
critical control on the observed PO4 enhancement in acidic atmosphere
conditions.
Annual mean concentrations (in ng-P m-3) of TP (a, c) and DP
(b, d) for the surface (a, b) and in the troposphere as zonal mean (c, d),
as calculated by the TM4-ECPL model for the present atmosphere (year 2008).
Finally, a significant amount of DOP (0.032 Tg-P yr-1) is added to the
total DP sources due to the ageing of OP-containing aerosols during
atmospheric transport (Table 3). This amount corresponds to about 12 % of
the global DP primary emission sources and to roughly 22 % of the total
dust-P acid-solubilization flux on a global scale. The ageing of organic-aerosol-carrying
P presents maxima over forested areas (about 0.1 ng-P m-2 s-1) due
to the high oxidation of PBAPs (Fig. 2d). Secondary maxima are also
calculated over China (0.01–0.1 ng-P m-2 s-1) and attributed to ageing
of primary OP of anthropogenic origin. Downwind of desert source regions,
significant DOP production rates, up to 0.1 ng-P m-2 s-1, are
calculated over the Sahara, the Thar and Gobi deserts; however, these DP
formation rates are more localized over continental regions than those due
to acid-solubilization mechanism of the dust mineral content (Fig. 2c).
However, non-negligible production of DOP is also calculated over the coastal
oceans, owing to the OP ageing under the long-range transport in the
atmosphere.
Evaluation of phosphorus simulations
Figure 4 presents the evaluation of present-day model simulation at various
locations around the globe (Fig. 4a, b; see also Sect. 2.4) against (1) P-containing aerosol airborne concentrations (Fig. 4c, e) and (2) dry
deposition fluxes (Fig. 4d, f). PO4 and TP aerosol concentrations are
provided in a daily resolution (except for TP concentrations from the
Corsica island which are provided as monthly means) and for different sizes;
i.e., fine (PM1 or PM1.3) and coarse (PM1 to PM10) or
PM10 aerosols or as bulk concentrations (Table S1). For this model
evaluation, a point-by-point comparison has been performed accounting for
the respective daily (or monthly) outputs and aerosol size of each
P-containing aerosol component of our model to the corresponding observation
database. The normalized mean bias (NMB) for the statistical analysis is
calculated as
NMB=∑i=1NMi-Oi∑i=1NOi×100%,
where Oi and Mi stand for observations and model predictions,
respectively, with N to represent the number of pairs (observations, model
predictions) that are compared. More information about the model performance
per database (cruise and stations) and aerosol size can be found in Fig. S4.
The comparison of all available DP aerosol measurements (fine, coarse and
bulk) with the respective model results is presented in Fig. 4c. DP aerosol
concentrations from cruise observations are in the range of about
3.1 × 10-6 – 4.03 × 10-2 µg-P m-3, while from station
observations, this range is about 3.23 × 10-4–1.37 × 10-1 µg-P m-3.
The model overestimates the DP cruise observations (NMB = 21 %) and underestimates the DP concentrations measured at
stations (NMB = -84 %). Focusing, however, on the size-segregated comparison of aerosol
DP (Fig. S4), the model underpredicts the observed concentrations at the
Finokalia station, both for fine and coarse particles, implying thus a
respective underestimation of P sources over land. On the contrary, for
cruise measurements, the model performs much better both for fine and coarse
aerosols as well as bulk observations. Note that the station observations
correspond to those of Finokalia and Corsica. Furthermore, only few cruise
TP observations are available (Graham and
Duce, 1982; Baker et al., 2006a, b) that
are discussed later in Sect. 3.5. The comparison presented here also
indicates that the model underpredicts (NMB = -59 %) the observed TP
concentrations at Finokalia (eastern Mediterranean; see also Fig. S4);
however, it simulates the bulk TP aerosol concentrations better at Corsica
(western Mediterranean). This implies that our model lacks TP sources in the
eastern Mediterranean atmosphere, which is strongly affected by air masses
from surrounding regions and by sources other than local ones.
As in the case of DP aerosol concentrations, the model simulates the DP dry
deposition fluxes better (NMB = 52 %) over oceanic regions (airborne cruise measurements compiled by
Vet et al., 2014)
than the observations (NMB = -93 %) at the Finokalia station (Fig. 4d).
Note that the same pattern is also calculated for the TP dry deposition
fluxes (Fig. 4f). The omission of super-coarse marine DP sources associated
with sea-salt particles can explain some discrepancies between model results
and observations only when these later concern bulk aerosols in oceanic
regions (so they could include super-coarse particles), which is the case
for wet or dry deposition samples. As discussed in Sect. 2.1.4, this
omission can affect local comparisons but overall does not introduce more
than a 3 % underestimate of DP flux over the ocean. In many cases, aerosol
samples have been collected with inlet devices that enable collection of
specific fractions of aerosols and eliminate super-coarse particles. When
bulk aerosols have been collected, then the presence of super-coarse
aerosols might introduce discrepancies between model results and
observations. Overall, the model performs better for DP dry deposition fluxes
over the oceans than over land, indicating a possible underestimate in the
continental source of P.
In Figs. S4 and S5, the results of
sensitivity simulations and the base case simulation are also presented with the aerosol
observations and dry and wet deposition fluxes, respectively. Fig. S6 also
shows the comparison of the annual cycles of the atmospheric concentrations
(TP and PO4) and deposition fluxes (dry and wet deposition), against the
TM4-ECPL monthly model results. For cruise measurements over the Atlantic
and Pacific oceans (Baker et al., 2010; Martino et al., 2014; Powell et al., 2015) and the global compilation of
deposition rates (Vet et al., 2014), the observations are also spatially averaged inside the same model grid box.
These comparisons show almost similar performance for all sensitivity
simulations but with one falling, in most cases, close to the lower edge of
observed concentrations and deposition fluxes. However, taking into account
the Wang et al. (2014) P-combustion sources,
the model performs better over the land (e.g., for TP concentrations at
Corsica, Fig. S4g; and for DP concentrations at the Finokalia monitoring
station, Fig. 6b, f, i), indicating that the base simulation underestimates
either anthropogenic combustion sources or other natural P sources.
Neglecting the P dissolution definitely degrades the comparisons of model
results with observations. On the other hand, the results show very small
sensitivity to the assumption of the soluble fraction of the primary emissions
of P. This finding supports the importance of the atmospheric processing of
dust for the atmospheric DP cycle as well as the potential underestimate of
the DP source in all sensitivity simulations. Such an underestimate could be
associated with an underestimate in the primary source or in the secondary
(atmospheric processing) of DP and deserves further studies.
Considering the scarcity of observational data and the gaps in knowledge of
P emissions and fate in the atmosphere, the simulated atmospheric P aerosol
concentrations (N = 1885) satisfactorily compare with the respective
available observations (NMB = -67 %) for TP (N = 585) and PO4
(N = 1300), and for P dry deposition fluxes (N = 819; NMB = -63 %),
indicating, however, an overall model underestimate of the observed values
(Fig. 4b). Based on these comparisons, we evaluate that an uncertainty of
about 70 % is associated with PRESENT model estimates.
Global and oceanic deposition fluxes of TP, DP and BP (in Tg-P yr-1),
as calculated by the TM4-ECPL model for PAST, PRESENT and FUTURE
simulations.
Deposition
TM4-ECPL
Global
Ocean
TP
PAST
1.262
0.270
PRESENT
1.300
0.281
FUTURE
1.270
0.272
DP
PAST
0.369
0.133
PRESENT
0.455
0.169
FUTURE
0.390
0.142
BP
PAST
0.384
0.135
PRESENT
0.470
0.172
FUTURE
0.405
0.144
Calculated annual deposition fluxes (in ng-P m-2 s-1)
for (a) TP and (b) DP for PRESENT simulation and their percentage differences
from PAST (c, d) and FUTURE (e, f) simulations, respectively. For the
PRESENT annual deposition fluxes (a, b), within brackets, the total amounts
over the globe (in parentheses only over the ocean) are also provided.
Global distribution of atmospheric phosphorus
TM4-ECPL calculates global TP and DP atmospheric burdens of 0.011 and
of 0.003 Tg-P, respectively. The calculated global annual mean TP and DP
atmospheric surface distributions for the present day are also shown in Fig. 5a and b. TP surface concentrations maximize over the major dust
regions of the world, roughly 0.1–1 µg-P m-3 (Fig. 5a), where
P-containing dust particles dominate the TP burden. Secondary maxima are
calculated over central Africa, Asia and Indonesia, where significant TP
concentrations (10–100 ng-P m-3) are associated with biomass burning
emissions and PBAPs (Fig. 5a). Over the oceans, however, TP concentrations
maximize downwind of dust source regions (roughly 10–100 ng-P m-3) and
secondary maxima of about 1–10 ng-P m-3 are calculated due to long-range
transport from continental sources, mainly over the NH.
Annual mean DP concentrations of 100 ng-P m-3 are calculated to occur
over the Sahara, the Arabian and Gobi deserts near the surface (Fig. 5b). The outflow from these source regions transports DP over the global
ocean where annual mean concentrations of about 10 ng-P m-3 are
calculated downwind of dust source regions, with the highest impact
calculated for the tropical Atlantic Ocean. The simulated concentrations of
DP over polluted regions range from 1 to 10 ng-P m-3, further
highlighting the importance of anthropogenic contributions to the DP
atmospheric burden – directly due to combustion emissions and indirectly due
to the solubilization of P when dust is mixed with atmospheric pollution
during atmospheric transport (Fig. 5b). TP emissions associated with African
dust are calculated to significantly affect the lower troposphere (Fig. 5c).
Furthermore, DP shows non-negligible concentrations in the middle
troposphere (Fig. 5d) that are attributed to transport from the source
regions and to atmospheric ageing (mainly P-solubilization processes) that
converts insoluble to soluble P, as already discussed.
Present-day phosphorus deposition flux
TM4-ECPL calculates that 1.300 Tg-P yr-1 of TP are deposited to the
Earth's surface of which about 0.281 Tg-P yr-1 over the ocean (Table 4). This oceanic deposition flux is calculated to be about half of that
estimated by Mahowald et al. (2008; 0.558 Tg-P yr-1) over oceans and at the low end of the
deposition flux range calculated by Wang et al. (2014; 0.2–1.6 Tg-P yr-1 over the ocean). The highest TP annual
deposition fluxes (up to 100 ng-P m-2 s-1) are calculated to occur
over the Sahara and the Gobi Desert while deposition fluxes up to 1 ng-P m-2 s-1 are also calculated at the outflow from dust source
regions, especially over the equatorial Atlantic and northern Pacific oceans
(Fig. 6a). The computed global DP deposition is calculated as equal to 0.455 Tg-P yr-1, of which 0.169 Tg-P yr-1 is deposited over the ocean
(Table 4), which is about 75 % higher than the estimate by
Mahowald et al. (2008; 0.096 Tg-P yr-1). The differences between the aforementioned studies can be
explained, on one hand, by the P-solubilization processes that only the
present study takes into account (and thus a greater amount of PO4 is
deposited at the Earth's surface) and, on the other hand, by the different
aerosol size representation that impacts on the lifetime of airborne
P-containing particles in the atmosphere. For this work, the highest DP
deposition fluxes are simulated to occur downwind of dust source regions,
owing to the DP content of the primary P emissions discussed in Sect. 3.1
and to P solubilization during atmospheric transport (Fig. 6b). Secondary DP
deposition flux maxima (about 0.1 ng-P m-2 s-1) are simulated
downwind of highly forested regions (i.e., Amazonia, central Africa and
Indonesia), reflecting the contribution of PBAPs to the DOP concentrations
in the atmosphere.
Figure S7 further presents the seasonal variability of DP deposition fluxes
as calculated by TM4-ECPL. The maximum seasonal DP deposition flux over the
ocean of 0.049 Tg-P is calculated to occur during June–July–August (Fig. S7c), followed by 0.048 Tg-P during March–April–May (Fig. S7b) and
by 0.038 Tg-P during September–October–November (Fig. S7d). The maximum DP
deposition flux in summer occurs when ocean stratification also maximizes,
thus leading to the highest impact of atmospheric deposition to the marine
ecosystems (Christodoulaki et al.,
2013). Furthermore, PBAP contribution maximizes in summer at regions with
important biogenic emissions (Fig. S8e–h), while dust contribution
maximizes in spring mainly over and downwind of the major deserts in the
tropical and midlatitudes of the Northern Hemisphere (Fig. S8a–d). This
is because the enhanced photochemistry during NH spring and summer increases
atmospheric oxidants and the atmospheric acidity due to NOx and
SOx oxidation. Note also that under equinox conditions, in particular
in spring, Sahara dust outbreaks are also favored (Fig. S8b). Considering
that most TP emissions occur in the NH (Fig. 2a), DP secondary formation
from IP and OP sources are simulated to maximize there (Fig. 2c,d), with
emissions from biomass burning and combustion of anthropogenic origin
further contributing to the DP deposition flux.
Annual mean percentage fractions of (a) P solubility (SP = %DP / TP) and (b) the
relative contribution of PBAPs to BP, in deposited
P-containing aerosols, as calculated by the TM4-ECPL model for the present
atmosphere (year 2008).
Phosphorus solubility
The present-day P solubility of deposited aerosols (hereafter SP = %DP / TP) is calculated to vary spatially (Fig. 7a), with minima (as low as
10 %) over dust source regions like the Sahara (where the insoluble
fraction of TP dominates aerosol content) and maxima (up to roughly 90 %)
over remote oceans such as the equatorial Pacific, southern Atlantic,
Indian and Southern oceans. Over such remote oceanic regions, high
solubility fractions are calculated due to low P-containing aerosol mass
concentrations that occur via the long-range transport of fine particles
from distant source regions, as well as the P which is associated with more aged
aerosols, and thus a greater fraction is present in the soluble mode, either
as DIP via mineral acid-solubilization processes or DOP via atmospheric
oxidation of P-containing organic aerosols and as PBAPs.
Vet et al. (2014), in their review paper of nutrient deposition, also mentioned that the P
solubility fractions of wet-only samples on coastal and inland sites have
been measured to range from 30 to 90 %, thus reflecting the effects of
combustion, biomass burning and phosphate fertilizers on airborne
phosphorus concentrations.
Anderson et al. (2010) reported that only 15–30 % of P in atmospheric aerosols at the Gulf of Aqaba was
in water-soluble phases or relatively soluble to be bioavailable to the
ecosystems. In the Mediterranean, the measured median solubilities of the
inorganic fraction of P in aerosols (ratio of PO4 to total inorganic P)
range between 20 and 45 % in the east Mediterranean with the lowest
values in dust-influenced air masses and the highest values in air masses
from the European continent (Markaki et al., 2003;
Herut et al., 1999) and have been reported to be around 38 % in the west
Mediterranean (Markaki et al., 2010). However, simultaneous observations of TP and DP deposition fluxes are
required to evaluate the solubility fraction of P (both organic and
inorganic) over remote oceans and thus to understand the atmospheric fate of
P. There are only a few aerosol data available in the literature for the
marine atmosphere (Graham and Duce, 1982; Baker et al., 2006a, b; Zamora et al.,
2013) that provide hints on the total P solubility. These data indicate P solubilities
ranging overall between 0.01 and 94 %, with the lowest values
corresponding to dust-influenced air masses and the highest to sea-salt-influenced
air masses. Over the Northern Hemisphere, Atlantic Ocean P
solubilities in aged Saharan dust aerosols have been measured to range from
0.01 to 37 % during oceanographic cruises (Baker et
al., 2006a, b). At Barbados, median solubilities of P on dust of about 19 %
and of sea-salt aerosol of about 94 % have been reported
(Zamora et al., 2013). In the southern Atlantic,
atmosphere P solubilities in aerosols of up to 67 % (median 8 % for dust
aerosol and 17 % for southern Atlantic aerosol; Baker
et al., 2006a) and of up to 87 % (median 32 %;
Baker et al., 2006b) have been reported. All these studies (but one) report P solubility as the ratio of PO4 to TP, thus neglecting the organic fraction which has been measured to
be about 28–44 % (Zamora et al., 2013). Although
these observations support high P solubilities in aged aerosols or aerosols
impacted by non-dust sources supporting the findings of our modeling
study, only the work by Zamora et al. (2013) could be
compared to the total P solubility simulated here (Fig. 7a). They indicate
that the model-simulated total P solubility is at the upper edge of observed
P solubilities.
The soluble P originating from each source as a fraction of TP from all
sources is shown in Fig. S9 for all P source categories (within each model
grid, these fractions sum up to 100 %). Note that in our SP fraction
calculations we also include the contribution of DOP from super-coarse PBAPs.
This assumption is followed since the DOP from PBAPs is considered readily
bioavailable, compared to other super-coarse particles, such as dust, for
which the bioavailability is characterized mainly from the initial dust's
solubility rather than from atmospheric processing due to the short atmospheric
lifetime of super-coarse particles. Indeed, super-coarse particles of that
size are basically emitted and deposited in the same model grid box
(Brahney et al., 2015), and are
not therefore expected to significantly impact SP fractions over remote
oceanic areas.
The low SP values over dust source regions are mainly attributed to both the
relatively low weathering of dust aerosols (10 %) assumed in emission
fluxes and low mineral P-dissolution rate (Fig. S9a). The low water associated
with dust aerosols near dust sources and the enhanced buffering capacity of
dust carbonate leading to excess of Ca2+ concentrations (see Sect. 3.1
and Fig. 2c) thus cause low P dissolution. The model calculates high SP
values (up to 50–60 %) over regions such as the Mediterranean basin, where
the co-existence of relatively high dust concentrations and high amounts of
anthropogenic pollutants (e.g., Kanakidou et
al., 2011) tends to enhance significantly atmospheric processing of mineral
P (Nenes et al., 2011). High dust SP values
are also calculated over the open ocean of the NH, the Atlantic Ocean, the
Pacific Ocean in the outflow of the Americas, downwind of the Arabian Desert
over the Indian Ocean and over the European continent. These results are
attributed to the mineral P solubilization under polluted acidic atmospheric
conditions.
Anthropogenic combustion aerosols are calculated to contribute significantly
to SP (20–30 %) over highly populated regions of the world, mainly over
the NH as in the case of the eastern and the western coasts of the USA,
central and northern Europe and western Asia (Fig. S9b). About 5–15 % of
the calculated SP over the remote oceans is attributed to long-range
transport, where aerosols have been subjected to atmospheric ageing. Biomass
burning aerosols are calculated to contribute regionally less than 30 % to
SP, with their maximum contribution over the equatorial Atlantic and Indian
oceans due to aerosol transport and the atmospheric ageing from central
Africa and India (Fig. S9c). DP emissions and atmospheric ageing associated
with PBAPs from terrestrial sources including super-coarse P-containing
bioaerosols (i.e., pollen in the present study) are calculated to
significantly contribute to DP deposition in the tropics (about 50 % in
the outflow of Amazonia and central African and Indonesian forests on an annual
mean basis; Fig. S9d). Seasonally, this contribution is even higher during
summer; for instance, it reaches 60 % in the Mediterranean and 70 % in
the outflow over the equatorial Pacific Ocean (not shown). DP from sea spray
dominates over all the remote Southern Ocean where no other significant
primary source of DP is present (Fig. S9e), while volcanic eruptions
contribute to the SP mainly over the equatorial and northern Pacific Ocean
(Fig. S9f).
Sensitivity of soluble phosphorus budget to air pollutants
Atmospheric acidity strongly depends on anthropogenic SOx, NOx and
NHx emissions and impacts on dust solubility. It is thus expected to
change in response to variability in the anthropogenic emissions of air
pollutants (Weber et al., 2016). The response of
atmospheric ageing of TP, which potentially converts the insoluble TP
fraction to DP, to air pollutant emission changes is assessed here by
comparing simulations performed using anthropogenic and biomass burning past
and future emissions to the present-day simulation (see Sect. 2). In
addition to dust dissolution changes, atmospheric organic aerosol (OA) ageing is also affected
by changes in oxidants levels
(Tsigaridis and Kanakidou, 2003;
Tsigaridis et al., 2006). Furthermore, primary anthropogenic and biomass
burning emissions of P also vary, as shown in Table 1 and discussed in Sect. 2.1. In particular, PRESENT TP anthropogenic emissions are estimated to have
increased by a factor of 5 since PAST and to be reduced back to almost the
PAST levels in the FUTURE. In the simulations discussed here, meteorology
and natural emissions of dust, sea salt, PBAPs and those from volcanoes are kept
constant to those of the year 2008 (i.e., PRESENT simulation). Although for
this work, we do not account for any changes in atmospheric dust emissions for
PAST and FUTURE simulations, several studies suggest that dust may vary
strongly and perhaps be sensitive to anthropogenic climate change and land
use (Ginoux et al., 2012; Mahowald et al., 2010; Prospero and Lamb, 2003) and thus could
also be an important driver of changes in the atmospheric P cycle. Overall,
for this study, the computed changes for species that regulate the mineral-P
acid solubilization (e.g., SO42-, NO3-, NH4+)
are due to the respective combustion emission differences between PAST,
PRESENT and FUTURE simulations.
For the PAST simulation, the anthropogenic emissions (e.g., NOx,
NHx and SOx) are a factor of 5–10 lower than present-day emissions
(Lamarque et al.,
2013). Compared to the present day, the model calculates significant changes
in the aerosol pH in the PAST simulation with less acidic pH near the
surface of the NH oceans, but a more acidic pH over the USA due to extensive
coal combustion in 1850
(Myriokefalitakis
et al., 2015). The FUTURE simulation globally projects a
less acidic aerosol pH than present day
(Myriokefalitakis
et al., 2015), owing to lower NOx and SOx emissions. Indeed, for
the FUTURE simulation, anthropogenic emissions (RCP 6.0) for most of the
continental areas are projected to be lower than the present day and to
almost return to 1850 levels due to air-quality regulations
(Lamarque et al.,
2013). However, as discussed in
Myriokefalitakis
et al. (2015) for the atmospheric cycle of Fe, due to the fact that biomass
burning emissions are projected to increase in the future, the system does
not fully return to 1850 conditions. Past and future changes of the
atmospheric acidity have a significant effect on mineral-P dissolution (Fig. S10c, d) and on the ageing of atmospheric OP (Fig. S10e, f). For the PAST
simulation, the model calculates about 40 % lower acid mineral-P
dissolution (0.085 Tg-P yr-1) compared to present day (0.144 Tg-P yr-1) while for the FUTURE, the acid mineral-P
solubilization (0.100 Tg-P yr-1) is projected almost 30 % lower than nowadays (Table 3).
Past and future changes in the phosphorus deposition flux
The global annual deposition fluxes of TP and DP as computed by TM4-ECPL for
the three main simulations (PAST, PRESENT, FUTURE) are provided in Table 4.
For the PAST simulation, the model calculates a global TP deposition flux
(Table 4) that is about 3 % lower than the present-day flux. Significant
increases in TP deposition fluxes since the PAST have been calculated over
Indonesia and southeastern Asia (Fig. 6c), as a result of the present high
anthropogenic emissions over China. As for the FUTURE simulation, TP
deposition is projected to decrease globally by 2.5 % compared to
present day (Table 4) with a maximum decrease (up to 40 %) due to emission
control measures calculated for Europe, the eastern USA and China.
On a global scale, DP deposition fluxes are also calculated to be lower by
about 20 % in the PAST and by about 15 % in the FUTURE simulations
compared to the PRESENT one (Table 4). Although these reductions are
computed to be relatively low, regional reductions can be stronger (up to
60 %; Fig. 6d), especially over highly populated regions (e.g., China,
Europe) and downwind of major dust sources (e.g., India and western USA).
Indeed, present-day DP emission fluxes from anthropogenic combustion (0.021 Tg-P yr-1) is calculated to decrease by about 80 % on a global scale
both for the PAST and FUTURE simulations (Table 1). According to our
calculations, however, present-day DP biomass burning emissions have
increased by about 28 % from PAST and are expected to further increase by
about 22 % in the FUTURE simulation. When accounting for both DP anthropogenic
combustion and biomass burning emissions, an approximately 3-fold increase is
computed from the PAST to PRESENT but a decrease to half is expected for the
FUTURE, thus contributing significantly to the DP atmospheric deposition
changes. Hence, DP deposition fluxes are projected to decrease over the
midlatitudes of the NH where human activities dominate (Fig. 6f), with the
largest changes up to 60 % over China due to the expected air-quality
measures, while smaller changes are computed over India due to the expected
increase in its population. Note again that our simulations neglect any
changes in dust and PBAP emissions that have occurred in the past or are
expected to take place in the future. Therefore, the changes in bioavailable
P (BP) deposition fluxes shown in Table 4 are driven by changes in the
anthropogenic and biomass burning emissions and in the atmospheric oxidants
that enhance P dissolution during atmospheric ageing.
Biogeochemical implications of changes in bioavailable phosphorus
deposition
The contribution of dust to the bioavailable P deposition flux into the
ocean maximizes in the outflow from desert regions, mainly in the North
Hemisphere tropics and midlatitudes (Fig. S8a–d). However, according to our
simulations, DOP is an important fraction of bioavailable P, mainly over
continental regions. Table 4 also presents the sum of the DP and the
insoluble PBAP deposition, reported as BP, which is
considered to be readily bioavailable for marine ecosystems since it is
biological material. Figure S8e–h depict the seasonal variability of the
PBAP deposition flux computed by our model for the PRESENT simulation. It
is remarkable that our simulations suggest that bioaerosols are a major
contributor to the BP deposition fluxes; on an annual basis, PBAPs
contribute about 25 % to the global BP deposition fluxes over the oceans
(about 33 % on a global scale), but regionally more than 50 % (Fig. 7b)
in the outflow of South America over the equatorial Pacific and in the
outflow of central Africa over the southern tropical Atlantic. This finding
clearly shows that biological material is a major atmospheric carrier of
bioavailable P to the global ocean (Fig. 7b) and implies a potentially
important impact of terrestrial sources on marine ecosystems. Note that, as
mentioned in Sect. 2, PBAPs from insect fragments and plant debris are
neglected in the present study; thus, their contribution to the bioavailable
P deposition mainly over land might be underestimated here. However, large
uncertainties are associated with this innovative finding, since the
estimates of the global source of PBAPs vary by more than an order of
magnitude; their size distributions and their mass density are uncertain, and
the P content in these aerosols is also highly variable, spanning 2 orders
of magnitude (e.g., the Supplement of Kanakidou
et al., 2012 and references therein). All these
parameters have to be studied by targeted experiments to improve knowledge
of their contribution to the atmospheric P cycle. Our results also indicate
that primary anthropogenic emissions of DP, as well as
anthropogenically driven atmospheric acidity, increased the DP supply to the
global ocean since the preindustrial period, thus providing an important
external-to-the-ocean source of nutrients for the marine ecosystem. They
also show that the P solubilization from dust aerosol during atmospheric
transport and mixing with acidic pollutants is important for DP deposition
and deserves further kinetic studies to improve parameterization of the
solubilization kinetics of various P-containing minerals as a function of
acidity and temperature. These results may be particularly important for
ecosystems like the east Mediterranean where phytoplankton growth is limited
by P availability.
It is also noteworthy that the bioavailable P deposition flux from
bioaerosols maximizes in summer (Fig. S8e–h) when ocean stratification is
also the strongest, thus leading to the highest impact of atmospheric
deposition to the marine ecosystems
(Christodoulaki et al., 2013). This
flux needs to be taken into account to evaluate the atmospheric DP
deposition impact on marine ecosystems. The computed atmospheric deposition
of BP over the global ocean of 0.17 Tg-P yr-1 (Table 4) represents
about 15 % of the global riverine flux to the ocean of 0.99 Tg-P yr-1 (Meybeck, 1982). However, while riverine inputs
affect mainly the coastal regions, atmospheric deposition is a source of
nutrients for the open sea (e.g., Okin
et al., 2011).