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Earth Syst. Dynam., 8, 1–54, 2017
www.earth-syst-dynam.net/8/1/2017/
doi:10.5194/esd-8-1-2017
© Author(s) 2017. CC Attribution 3.0 License.
Climate engineering by mimicking natural dust climate
control: the iron salt aerosol method
Franz Dietrich Oeste1, Renaud de Richter2, Tingzhen Ming3, and Sylvain Caillol2
1gM-Ingenieurbüro, Tannenweg 2, 35274 Kirchhain, Germany
2Institut Charles Gerhardt Montpellier – UMR5253 CNRS-UM2 – ENSCM-UM1 – Ecole Nationale
Supérieure de Chimie de Montpellier, 8 rue de l’Ecole Normale, 34296 Montpellier CEDEX 5, France
3School of Civil Engineering and Architecture, Wuhan University of Technology, No. 122, Luoshi Road,
Hongshan District, Wuhan, 430070, China
Correspondence to: Franz Dietrich Oeste (oeste@gm-ingenieurbuero.com)
Received: 8 August 2016 – Published in Earth Syst. Dynam. Discuss.: 10 August 2016
Revised: 10 December 2016 – Accepted: 12 December 2016 – Published: 13 January 2017
Abstract. Power stations, ships and air traffic are among the most potent greenhouse gas emitters and are
primarily responsible for global warming. Iron salt aerosols (ISAs), composed partly of iron and chloride, exert a
cooling effect on climate in several ways. This article aims firstly to examine all direct and indirect natural climate
cooling mechanisms driven by ISA tropospheric aerosol particles, showing their cooperation and interaction
within the different environmental compartments. Secondly, it looks at a proposal to enhance the cooling effects
of ISA in order to reach the optimistic target of the Paris climate agreement to limit the global temperature
increase between 1.5 and 2 ◦C.
Mineral dust played an important role during the glacial periods; by using mineral dust as a natural analogue
tool and by mimicking the same method used in nature, the proposed ISA method might be able to reduce and
stop climate warming. The first estimations made in this article show that by doubling the current natural iron
emissions by ISA into the troposphere, i.e., by about 0.3 Tg Fe yr−1, artificial ISA would enable the prevention
or even reversal of global warming. The ISA method proposed integrates technical and economically feasible
tools.
1 Introduction
The 5th Assessment Report of the Intergovernmental Panel
on Climate Change (IPPC), released in November 2014,
states that global warming (GW) has already begun to dra-
matically change continental and marine ecosystems.
A recently noticed effect is that the vertical mixing
in oceans decreases and even reaches a stagnation point
(de Lavergne et al., 2014), thus weakening the net oceanic
cumulative intake of atmospheric CO2(Bernardello et al.,
2014a, b).
A consequence of decreasing vertical ocean mixing is a
reduced or interrupted oxygen supply to the depths of the
ocean. Currently, the formation of low-oxygen areas in the
oceans is increasing (Capone and Hutchins, 2013; Kalvelage
et al., 2013). Furthermore, climate warming entails stratifica-
tion of the water column and blocks vertical flows. Stratifica-
tion may develop by warming the upper water layer as well
as by evaporation and precipitation. Generation of a fresh-
water layer on top of the water column by precipitation, sur-
face water runoff and meltwater inflow induces stratification
(Hansen et al., 2016; van Helmond et al., 2015). Even the
opposite, brine generation by evaporation may, induce strati-
fication (Friedrich et al., 2008). Stratification blocks the oxy-
gen transfer through the water column and triggers the for-
mation of oxygen-depleted zones (Voss et al., 2013) that also
emit nitrous oxide (N2O), a potent greenhouse gas (GHG)
and a powerful ozone-depleting agent.
As iron is part of many enzymes directing the bioenergetic
transformation of nitrogen in the ocean, it has an additional
direct influence on the cycling of these elements through
Published by Copernicus Publications on behalf of the European Geosciences Union.
2 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
the oceanic environment (Klotz and Stein, 2008; Simon and
Klotz, 2013).
The severest consequence for oceanic ecosystems of such
stratification is the development of anoxic milieus within
stratified ocean basins. An example of the development of
halocline and chemocline stratification is the Black Sea (Eck-
ert et al., 2013). This ocean basin has a stable halocline which
coincides with a chemocline, dividing an oxic salt-poor sur-
face water layer from a saline anoxic sulfidic deep layer with
a black sapropel sediment rich in organic C at the basin bot-
tom (Eckert et al., 2013).
Geological past episodes with stratified ocean basins are
regularly marked by black shale or black limestone as rem-
nants of sapropel sediments. Stratified ocean basins during
the Phanerozoic epoch occurred as a consequence of elevated
CO2levels in the atmosphere. This caused high sea surface
temperatures (Meyers, 2014) and, as a global consequence, a
global increase in evaporation, precipitation and production
of brines of higher concentrations.
It has been pointed out that the increasing meltwater runoff
from past polar and subpolar ice layers may have induced the
cover of denser ocean water by a meltwater layer (Hansen
et al., 2016). According to Praetorius et al. (2015), climate
warming events during the last deglacial transition induced
subsurface oxygen minimum zones accompanied by sea floor
anoxia in the northern Pacific. This meltwater-induced strat-
ification was accompanied by meltwater iron-induced phyto-
plankton blooms. The generation of increasing precipitation
and surface water runoff accompanied by increasing brine
production plus elevated surface water temperatures during
hot high-CO2climate episodes had similar consequences in
past geological epochs (Meyers, 2014).
Ocean basin stratifications may be induced by increasing
precipitation with increased surface water runoff (van Hel-
mond et al., 2015) or by increased brine production
(Friedrich et al., 2008). Such an ocean stratification event is
characterized by regional to global ocean anoxia, black sedi-
ments with elevated organic C and a hot greenhouse climate,
as we can see from the whole Phanerozoic past (Meyers,
2014), and was often accompanied by mass extinctions.
Even the largest mass extinction of ocean biota within the
Phanerozoic epoch, during the Permian–Triassic transition,
was induced by high temperatures as a consequence of el-
evated CO2levels, which induced the change from a well-
mixed oxic to a stratified euxinic–anoxic ocean (Kaiho et al.,
2016).
What we have to face now is the extraordinary process de-
veloping from the recent situation: the combination of the
CO2-dependent temperature-rise-generated precipitation in-
crease, plus a meltwater increase. Mankind has to now find
the appropriate tool to stop this dangerous stratification pro-
cess.
Warming surface waters and a decreasing input of cold,
oxygenated surface water trigger a temperature rise in sed-
iments, transforming solid methane hydrate into gaseous
methane (CH4) emissions in seawater (Phrampus et al.,
2014). CH4oxidation consumes additional oxygen, decreas-
ing the oxygen content above those areas (Yamamoto et al.,
2014).
The same effects are expected with an anticipated increase
in spring and summer coastal upwelling intensity, associ-
ated with increases in the rate of offshore advection, decreas-
ing the nutrient supply while producing a spatial or tempo-
ral (phenological) mismatch between production and con-
sumption in the world’s most productive marine ecosystems
(Bakun et al., 2015).
These events have the threatening consequence of a
widespread lack of oxygen in the oceans. In such low-oxygen
areas (sub-oxic to anoxic) only bacterial life is possible:
higher life forms cannot exist there. Accordingly, an early
result of the progression of climate warming could lead to
a dramatic limitation of the oceanic food sources that will
be needed for the projected 9–10 billion people by 2050.
The same deleterious consequences for seafood supply can
also result from ocean surface acidification through increased
CO2dissolution in seawater and a decreased flow of surface
water currents to ocean basin bottoms, limiting reef fish and
shelled mollusk survival (Branch et al., 2013).
Any decrease in the thermohaline circulation (THC) has
severe consequences for all kinds of ecosystems as it further
triggers climate warming by different interactions. THC de-
crease induces a reduction in or eventual disappearance of
the phytoplankton fertilizers Si, P, N and Fe extracted on the
ocean surface from their sources at the bottom of the ocean
basins. Hydrothermal fluid cycling by mid-ocean ridges, off-
axis hydrothermal fluid fluxes, subduction-dependent hy-
drothermal convection fluids, hydrothermal fluxes at hot spot
sea mounts and fluid emissions from anaerobic sediments
contain said elements as dissolved or colloidal phase (Dick
et al., 2013; Hawkes et al., 2013; Holm and Neubeck, 2009;
Martin and Russell, 2007; Orcutt et al., 2011; Postec et al.,
2015; Resing et al., 2015; Sousa et al., 2013). The deeper wa-
ter of all ocean basins is enriched by these fertilizers. A THC
decrease within the ocean basins will result in a decrease in
the assimilative transformation of CO2into organic carbon.
Moreover, any THC decrease would further trigger the
acidification of the ocean surface by lowering or preventing
the neutralization of dissolved CO2and HCO−
3due to the al-
kalinity decrease from hydrothermal sources (Monnin et al.,
2014; Orcutt et al., 2011).
During the convective water flow through the huge al-
kaline ocean crust volume, estimated to be about 20–
540 ×103km3yr−1(Nielsen et al., 2006), ocean water is de-
pleted in O2but enriched in its reductant content such as CH4
(Kawagucci et al., 2011; Orcutt et al., 2011). Other elements
are enriched in this convective water flow through the Earth
crust, essential for the existence of life. The reoxygenation of
this huge water volume is retarded or even impossible with a
minimized THC.
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 3
According to model calculations (Watson et al., 2015),
the THC might have significantly changed between the last
glacial and interglacial periods. During the Cenozoic epoch,
ice-covered pole caps limited the incorporation of carbon in
the form of carbonate into the oceanic crust compared to
the warm late Mesozoic period (Coogan and Gillis, 2013).
The findings of Coogan and Gillis (2013) show that during
ice-free periods, THCs were possible with much higher ef-
fectiveness than in modern times. Even during those warm
periods with low temperature gradients between polar and
equatorial oceans, an effective production of brines lead-
ing to buoyancy differences necessary for the development
of effective THC may have ben generated (Otto-Bliesner et
al., 2002). However, increased inflow rates of high-density
brines coming from shallow shelf regions with high evapora-
tion rates induced several collapses or vertical reductions of
the strong Cretaceous THC. From here and for more than a
million years, the lower parts of ocean basins have been filled
with anoxic brines (Friedrich et al., 2008). Further aspects of
ocean stratification are discussed in Sect. 4.1.
Remnants of these anoxic events are black shale sediments
(Takashima et al., 2006). During such THC collapses, the up-
take of CO2into the oceanic crust stayed restricted to organic
carbon sediments. Additionally, the organic carbon produc-
tivity of the remaining oxic zone was decreased, as was eolic
dust input, due to phytoplankton fertilizer production being
limited to continental weathering.
These examples point out the sensitivity of the THC to dis-
turbances. Without action, the weakness of our recent THC
may worsen. Any THC collapse would not only result in se-
vere damages to ecosystems, food chains and food resources
of the oceans but would also lead to an acceleration of the in-
crease in atmospheric CO2concentration, resulting in faster
climate warming than forecasted.
The best way to prevent such worrying situations and con-
sequences is to stop GW.
A realistic chance of averting this development is the con-
trolled application of a climate cooling process, used several
times by nature throughout the last ice ages with high ef-
ficiency and based on loess dust. Loess is a windblown dust
sediment formed by progressive accumulation and composed
generally of clay, sand and silt (approximately at a ratio of
20 : 40 : 40, respectively), loosely cemented by calcium car-
bonate.
The dust concentration in the troposphere increased dur-
ing every cold period in ice ages and reached a multiple of
today’s levels (Martínez-Garcia et al., 2011). Dust deposi-
tion in the Southern Ocean during glacial periods was 3 to
10 times greater than during interglacial periods, and its ma-
jor source region was probably Australia or New Zealand
(Lamy et al., 2014). The windblown dust and its iron con-
tent effect on marine productivity in the Southern Ocean is
thought to be a key determinant of atmospheric CO2concen-
trations (Maher and Dennis, 2001). During high dust level
periods, the global average temperature fell to 10◦C (Lamy
et al., 2014; Martin, 1990; Martínez-Garcia et al., 2011),
which is 4.5 ◦C lower than the current global average temper-
ature. Loess sediments in the Northern and Southern Hemi-
sphere on continents and ocean floors originate from these
cold dusty periods.
Former geoscientists had the predominant conception that
the cold glacial temperatures had caused dustiness, and not
the reverse (Maher et al., 2010). Meanwhile, more evidence
has accumulated that mineral dust was a main factor in the
cause of the cold periods and that the iron (Fe) fraction of
windblown dust aerosol fertilized the oceans’ phytoplank-
ton, activating the assimilative conversion of CO2into or-
ganic carbon (Anderson et al., 2014; Lamy et al., 2014; Ma-
her et al., 2010; Martin, 1990; Martínez-García et al., 2014;
Ziegler et al., 2013) and carbonate, which composes the main
dry-body substance of phytoplankton, together with silica,
another component of dust (Tréguer and Pondaven, 2000).
Evidence regarding the role of iron-containing dust in trig-
gering ice ages during the late Paleozoic epoch is currently
being discussed (Sur et al., 2015).
The biogeochemical cycles of carbon, nitrogen, oxygen,
phosphorus, sulfur and water are well described in the lit-
erature, but the biogeochemical cycle of the Earth’s iron is
often overlooked. An overview of the progress made in the
understanding of the iron cycle in the ocean is given by sev-
eral authors (Breitbarth et al., 2010; Raiswell and Canfield,
2012).
The current state of knowledge of iron in the oceans is
lower than that of carbon, although numerous scientific pub-
lications deal with this topic (Archer and Johnson, 2000;
Boyd and Ellwood, 2010; Johnson et al., 2002a, b; Misumi
et al., 2014; Moore and Braucher, 2008; Moore et al., 2013;
Tagliabue et al., 2015; Turner and Hunter, 2001); however,
the iron biogeochemical cycle in the atmosphere is described
by fewer authors (Mahowald et al., 2005, 2009, 2010). This is
in contrast to the iron biogeochemical cycle in soil and land,
as almost no recent publications details the current knowl-
edge about iron in soils and over the landscape (Anderson,
1982; Lindsay and Schwab, 1982; Mengel and Geurtzen,
1986), a task we attempt in this review.
The process of iron fertilization by the injection of an iron
salt solution into the ocean surface has already been dis-
cussed as an engineering scheme proposed to mitigate global
warming (Smetacek and Naqvi, 2008). But iron fertilization
experiments with FeSO4conducted over 300km2in the sub-
antarctic Atlantic Ocean, although doubling primary produc-
tivity of chlorophyll a, did not enhance downdraft particles’
flux into the deep ocean (Martin et al., 2013). The researchers
who carried out this work attribute the lack of fertilization-
induced export into the deep ocean to the limitation of sili-
con needed for diatoms. Thus, ocean fertilization using only
iron can increase the uptake of CO2across the sea surface,
but most of this uptake is transient and will probably not
lead to long-term sequestration (Williamson et al., 2012). In
other experiments, the authors (Smetacek et al., 2012) find
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4 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
that iron-fertilized diatom blooms may sequester carbon for
centuries in ocean bottom water, and for longer in the sed-
iments, as up to half the diatom bloom biomass sank below
1 km depth and reached the sea floor. Meanwhile, dissolution
of olivine, a magnesium–iron–silicate-containing silica, with
a Mg :Fe ratio of nearly 9:1, resulted in 35% marine carbon
uptake (with the hypothesis of 1 % of the iron dissolved and
biologically available), with communities of diatoms being
one of the phytoplankton winners (Köhler et al., 2015).
The idea of climate cooling by CO2carbon conversion into
organic sediment carbon by the addition and mixture of an
iron salt solution into the ocean with a marine screw pro-
peller has been the subject of controversial debates (Boyd
and Bressac, 2016; Chisholm et al., 2002; Johnson and Karl,
2002). The eolic iron input per square meter of ocean sur-
face by natural iron salt aerosol (ISA) is on the order of tens
of milligrams of Fe per square meter per year. In compari-
son, the artificial Fe input by ship screws is several orders of
magnitude above the natural fertilization with ISA.
The small content of water-soluble iron salts (ISs) in the
dust particles triggers this fertilization effect (Duggen et al.,
2007), and the soluble-iron deposition during glaciations
had been up to 10 times the modern deposition (Conway
et al., 2015). According to Spolaor et al. (2013), most of
the bioavailable water-soluble Fe(II) has been linked, dur-
ing the last 55 000 years, to the fine dust fraction, as has
been demonstrated from ice cores from Antarctica. Glacial-
stage dust fluxes of ∼400 to 4000 times those of interglacial
times have been found from late Paleozoic epochs (Soreghan
et al., 2014), which gives an estimated carbon fixation ∼2–
20 times that of modern carbon fixation due to dust fertil-
ization. Photochemistry by sunshine is the main trigger of
the transformation of the primary insoluble-iron fraction of
dust aerosols into soluble iron salts (Johnson and Meskhidze,
2013), and the understanding of how the different iron con-
tent and speciation in aerosols affect the climate is growing
(Al-Abadleh, 2015). Currently, increased subglacial meltwa-
ter and icebergs may supply large amounts of bioavailable
iron to the Southern Ocean (Death et al., 2014). The flux of
bioavailable iron associated with glacial runoff is estimated
at 0.40–2.54 Tg yr−1in Greenland and 0.06–0.17Tgyr−1in
Antarctica (Hawkings et al., 2014), values which are com-
parable with eolian dust fluxes to the oceans surrounding
Antarctica and Greenland and will increase by enhanced
melting in a warming climate.
However, CO2uptake by the oceans is not the only effect
of iron dust. The full carbon cycle is well described in the
literature; at the same time, we know less about the iron bio-
geochemical cycle. Recently, the major role of soluble-iron
emissions from combustion sources has become more evi-
dent. Today, the anthropogenic combustion emissions play a
significant role in the atmospheric input of soluble iron to the
ocean surface (Sedwick et al., 2007). Combustion processes
currently contribute from 20 to 100 % of the soluble-iron de-
position over many ocean regions (Luo et al., 2008). Model
results suggest that human activities contribute to about half
of the soluble-Fe supply to a significant portion of the oceans
in the Northern Hemisphere (Ito and Shi, 2016) and that de-
position of soluble iron from combustion sources contributes
more than 40 % of the total soluble-iron deposition over sig-
nificant portions of the open ocean in the Southern Hemi-
sphere (Ito, 2015). Anthropogenic aerosol associated with
coal burning is maybe the major bioavailable iron source in
the surface water of the oceanic regions (Lin et al., 2015).
The Fe emission from coal combustion, higher than previ-
ously estimated, implies a larger atmospheric anthropogenic
input of soluble Fe to the northern Atlantic and northern Pa-
cific Oceans, which is expected to enhance the biological car-
bon pump in those regions (Wang et al., 2015b).
The limited knowledge about dissolved or even dispersed
iron distributions in the ocean confirms the work of Tagliabue
et al. (2015): their calculation results of the residence time of
iron in the ocean differ by up to 3 orders of magnitude from
the different published models.
The precipitation of any iron salt results from the pH and
O2content of the ocean water milieu. But the presence of
organic Fe chelators such as humic or fulvic acids (Mis-
umi et al., 2014) as well as complexing agents produced
by microbes (Boyd and Ellwood, 2010) and phytoplankton
(Shaked and Lis, 2012) life forms prevents iron from precip-
itation. In principle, this allow the transport of iron, from its
sources, to any place within the ocean across huge distances
with the ocean currents (Resing et al., 2015). But organic
material and humic acids have a limited lifetime in oxic envi-
ronments due to their depletion to CO2. But within stratified
anoxic ocean basins, their lifetime is unlimited.
The iron inputs into the ocean regions occur by atmo-
spheric dust, coastal and shallow sediments, sea ice, icebergs,
and hydrothermal fluids and deep-ocean sediments (Boyd
and Ellwood, 2010; Elrod et al., 2004; Johnson et al., 1999;
Mahowald et al., 2005, 2009; Moore and Braucher, 2008;
Raiswell et al., 2016; Wang et al., 2015b).
Microbial life within the gradient of chemoclines divid-
ing anoxic from oxic conditions generates organic carbon
from CO2or HCO−
3carbon (Borch et al., 2009; Schmidt et
al., 2008; Sylvan et al., 2012). The activity at these chemo-
clines is the source of dissolved Fe(II). Humic acid is a main
product of the food chain within any life habitat. Coastal,
shelf and ocean bottom sediments, as well as hydrothermal
vents and methane seeps, are such habitats and are known as
iron sources (Boyd and Ellwood, 2010). Insoluble Fe oxides
are part of the lithogenic particles suspended at the surface
of the Southern Ocean. Along with organic phytoplankton
substance, the suspended inorganics complete the gut pas-
sage of krill. During the gut passage of these animals, iron
is reduced and leaves the gut in a dissolved state (Schmidt
et al., 2016). There is no doubt that gut-microbial attack on
ingested organics and inorganics produces faeces containing
humic acids. This metabolic humic acid production is known
from earth worm faeces (Muscolo et al., 2009) and human
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 5
faeces (Reck et al., 2015; Wagner Mackenzie et al., 2015).
The effect of iron mobilization from lithogenic particles by
reduction during gut passage has been found in termites, too
(Vu et al., 2004). The parallel generation of Fe-chelating hu-
mic acids during gut passage guarantees that the Fe is kept
in solution after leaving the gut and entering the ocean. The
examples demonstrate that every link of the ocean food chain
may act as a source of dissolved iron.
The cogeneration of Fe(II) and Fe-chelating agents at any
Fe sources at the bottom, surface and shelves of the oceans
is the precondition of iron transport between source and
phytoplankton at the ocean surface. But the transport be-
tween sources and the phytoplankton depends on the verti-
cal and horizontal movement in the ocean basins (Misumi et
al., 2014; Moore et al., 2013). Any movement between iron
sources and the phytoplankton-rich surface in stratified ocean
basins remains restricted to the surface, near Fe input from its
sources (shelf sediments, meltwater, icebergs, rivers, surface
water runoff and dust input).
During the glacial maxima the vertical movement reached
an optimum. According to this, Fe transport from basin bot-
tom sources and dust sources to the phytoplankton was at its
maximum and produced a maximum primary productivity at
the ocean surface, but carbon burial became lowest during
that time (Lopes et al., 2015) although GHGs were at their
lowest levels during the glacial maximum. The cause of this
seeming contradiction are the changing burial ratios of or-
ganic C / carbonate C at the basin bottom(s). The burial ratio
is high during episodes with a stratified water column, and
it is very low during episodes with a vertically mixed water
column, as we demonstrate in Sect. 4 in detail.
This review aims to describe the multistage chemistry of
the iron cycle in the atmosphere, oceans, lands, sediments
and ocean crust. This article is a comprehensive review of
the evidence for connections between the carbon cycle and
the iron cycle and their direct and indirect planetary cooling
effects. Numerous factors influence the Fe cycle and the iron
dissolution: iron speciation, photochemistry, biochemistry,
redox chemistry, mineralogy and geology. In order to per-
form an accurate prediction of the impact of Fe-containing
dusts, sea salt and acidic components, atmospheric chemistry
models need to incorporate all relevant interaction compart-
ments of the Fe cycle with sun radiation, chlorine, sulfur, ni-
trogen and water. This review advocates a balanced approach
to benefit from the Fe cycle in order to fight global warming
by enhancing natural processes of GHG depletion, albedo
increase, carbon burial increase and de-stratification of the
ocean basins.
1.1 Breakdown of sections
The next four sections describe nearly a dozen different cli-
mate cooling processes induced by ISAs and their interaction
for modeling parameter development (Sects. 2–5). Then es-
timation of the requirements in terms of ISA to stop global
warming will be given in Sect. 6, followed by the description
of a suggested ISA-enhanced method to fight global warming
and induce planetary cooling in Sect. 7, and the possible risks
of reducing acids and iron emissions in the future in Sect. 8.
This is in turn followed by a general discussion and conclud-
ing remarks in Sects. 9 and 10. To our knowledge, this re-
view completes, with atmosphere, solid and liquid surfaces at
the surface of the globe, oceans, sediments and oceanic crust
(Pérez-Guzmán et al., 2010), the previous ocean global iron
cycle vision of Parekh (Parekh et al., 2004), Archer and John-
son (2000), Boyd and Ellwood (2010) and of many others. It
advocates a balanced approach to make use of the iron cycle
to fight global warming by enhancing natural processes.
1.2 Components of the different natural cooling
mechanisms by ISA
The best known cooling process induced by ISA is the phyto-
plankton fertilizing stage described in the introduction. But
this process is only part of a cascade of at least 12 climate
cooling stages presented in this review. These stages are em-
bedded within the coexisting multi-component complex net-
works of different reciprocal iron-induced interactions across
the borders of atmosphere, surface ocean, sediment and ig-
neous bedrock as well as across the borders of chemistry,
biology and physics and across and along the borders of il-
luminated, dark, gaseous, liquid, solid, semi-solid, animated,
unanimated, dead and different mix phase systems. Some im-
pressions of the complexity of iron acting in the atmospheric
environment have been presented by Al-Abadleh (2015).
The ISA-induced cooling effect begins in the atmosphere.
Each of the negative forcing stages unfolds a climate-cooling
potential for itself. Process stages 1–6 occur in the tropo-
sphere (Sect. 2), stage 6 at sunlit solid surfaces, stages 7–8 in
the ocean (Sect. 3), and stages 9–12 in the oceanic sediment
and ocean crust (Sect. 4). Other possible cooling stages over
terrestrial landscapes and wetlands are described in Sect. 5.
At least 12 stages of this cooling process cascade operate as
described below.
2 Tropospheric natural cooling effects of the iron
cycle
2.1 ISA-induced cloud albedo increase
ISA consists of iron-containing particles or droplets with a
chloride content. Aerosols have significant effects on the cli-
mate (Forster et al., 2007). First, by direct scattering of radi-
ation, and second, by inducing a cloud albedo increase. The
latter effect is induced by cloud whitening and cloud lifetime
elongation. Both effects induce a climate cooling effect by
negative radiative forcing of more than −1 W m−2.
Aerosols have a climate impact through aerosol–cloud
interactions and aerosol–radiation interactions (Boucher,
2015). By reflecting sunlight radiation back to space, some
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6 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
types of aerosols increase the local albedo (which is the frac-
tion of solar energy that is reflected back to space), produc-
ing a cooling effect (Bauer and Menon, 2012). If the top
of clouds reflect back a part of the incident solar radiation
received, the base of clouds receives the long-wave radia-
tion emitted from the Earth surface and reemits part of it
downward. Usually, the higher a cloud is in the atmosphere,
the greater its effect on enhancing atmospheric greenhouse
warming, and therefore the overall effect of high-altitude
clouds, such as cirrus, is a positive forcing. Meanwhile, the
net effect of low-altitude clouds (stratocumulus) is to cool
the surface, as they are thicker and prevent more sunlight
from reaching the surface. The overall effect of other types
of clouds such as cumulonimbus is neutral: neither cooling
nor warming.
More outgoing long-wave radiation is possible when the
cirrus cover is reduced. Efficient ice nuclei (such as bismuth
triiodide) seeding of cirrus cloud might artificially reduce
their cover (Mitchell and Finnegan, 2009; Storelvmo et al.,
2013).
In order to enhance the cooling effects of low-altitude
clouds, marine cloud brightening has been proposed (Latham
et al., 2012a), for instance by injecting sea-salt aerosols over
the oceans. The effect depends on both particle size and in-
jection amount, but a warming effect is possible (Alterskjær
and Kristjánsson, 2013).
Aerosol effects on climate are complex because aerosols
both reflect solar radiation to space and absorb solar radia-
tion. In addition, atmospheric aerosols alter cloud properties
and cloud cover depending on cloud type and geographical
region (Koch and Del Genio, 2010). The overall effect of
aerosols on solar radiation and clouds is negative (a cooling
effect), which masks some of the GHG-induced warming.
But some individual feedbacks and forcing agents (black car-
bon, organic carbon and dust) have positive forcing effects (a
warming effect). For instance, brown clouds are formed over
large Asian urban areas (Ramanathan et al., 2007) and have a
warming effect. The forcing and feedback effects of aerosols
have been clarified (Bauer and Menon, 2012) by separating
direct, indirect, semi-direct and surface albedo effects due to
aerosols.
Differing from any natural dust iron-containing mineral
aerosol, the ISA aerosol does not contain any residual min-
eral components such as Fe2O3minerals, known as strong
radiation absorbers. Previous studies have shown that iron
oxides are strong absorbers at visible wavelengths and that
they can play a critical role in climate perturbation caused by
dust aerosols (Sokolik and Toon, 1999; X. L. Zhang et al.,
2015). As the primary ochre-colored aerosol particles emit-
ted by the ISA (method I, see Sect. 7) have small diameters
of <0.05 µm and are made of pure FeOOH, they become
easily and rapidly dissolved within the plume of acidic flue
gas. The ISA FeOOH aerosol is emitted with the flue gas
plumes generated in parallel and containing SO2and NOx
as sulfuric and nitric acid generators. ISA stays within the
troposphere for weeks before precipitating on the ocean or
land surfaces. Due to their small diameter and high surface
area, the aerosol particles will immediately react with HCl,
generated as a reaction product between sea-salt aerosol and
the flue-gas-borne acids. The reaction product is an orange-
colored FeCl3aerosol: ISA. During daytime the sunlight ra-
diation bleaches ISA into FeCl2and q
Cl; at night the reoxi-
dation of ISA plus HCl absorption generates ISA again. The
FeCl2aerosol particles are colorless at low humidity and pale
green during high-humidity episodes. The daytime bleaching
effect reduces the radiation absorption of ISA to much lower
levels compared to oxides such as Fe2O3.
Hygroscopic salt aerosols act as cloud condensation nu-
clei (CCN; Karydis et al., 2013; Levin et al., 2005). ISA
particles are hygroscopic. High CCN particle concentrations
have at least three different cooling effects (Rosenfeld and
Freud, 2011; Rosenfeld et al., 2008). Each effect triggers
the atmospheric cooling by a separate increase of earth re-
flectance (albedo) (Rosenfeld et al., 2014):
–cloud formation (even at low supersaturation);
–formation of very small cloud droplets, with an elevated
number of droplets per volume, which causes elevated
cloud whiteness;
–extending the lifetime of clouds, as the small cloud
droplets cannot coagulate with each other to induce pre-
cipitation fall.
Figure 1 illustrates this albedo change due to ISA CCN
particles.
Additional to climate cooling effects, CCN-active aerosols
might induce a weakening of tropical cyclones. The cooling
potential of the ocean surface in regions of hurricane gene-
sis and early development potential (Latham et al., 2012b)
may be possible by cloud brightening. Further effects such
as delayed development, weakened intensity, early dissipa-
tion and increased precipitation have been found (Y. Wang et
al., 2014; H. Zhang et al., 2009).
2.2 Oxidation of methane and other GHGs
Currently, methane (CH4) in the troposphere is destroyed
mainly by the hydroxyl radical q
OH.
Around 3 to 4 % CH4(25 Tg yr−1) (Allan et al., 2007;
Graedel and Keene, 1996) are oxidized by q
Cl in the tropo-
sphere, and larger regional effects are predicted: up to 5.4 to
11.6 % CH4(up to 75 Tg yr−1) in the Cape Verde region
(Sommariva and von Glasow, 2012) and ∼10 to >20 % of
total boundary layer CH4oxidation in other locations (Hos-
saini et al., 2016).
According to Blasing (Blasing, 2010, 2016; Forster et al.,
2007), the increase in the GHG CH4since 1750 has induced
a radiative forcing of about +0.5 W m−2. The research re-
sults of Wittmer et al. (2015a, b, 2016) and Wittmer and
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 7
CCN CCN CCN
ISA particle
FeOOH
FeCl3
H2O
HBr
HCl
H2SO4
particle
Oxidation
Combustion-induced
ISA emission
ISA-induced phytoplankton
quantity increase
Flue gas
Oxidation
H2SO4
particle
CH3Br
CH3Cl
S(CH3)2
HCl
SO2
NOX
NaCl
H2O
NaCl
Na2SO4
NaNO3
Aged sea-salt particle
HNO3
Bubble-induced sea-salt
aerosol emission
Figure 1. Process of tropospheric cooling by direct and indirect increasing of the quantity of different cloud condensation nuclei (CCN)
inducing albedo increase by cloud formation at low supersaturation, cloud whitening and cloud life elongation.
Zetzsch (2016) demonstrated the possibility of significantly
reducing the CH4lifetime by the ISA method significantly.
According to Anenberg et al. (2012) the health effects of the
combination of increased CH4and NOx-induced O3levels
with an increase in black carbon are responsible for tens of
thousands of deaths worldwide.
Any increase in the q
Cl level will significant elevate the de-
pletion rate of CH4and volatile organic compounds (VOCs)
as well as ozone (O3) and dark carbon aerosol as described
in Sects. 2.3 and 2.4.
Absorption of photons by semiconductor metal oxides can
provide the energy to produce an electron–hole pair able
to produce either a reduced or an oxidized compound. In
suitable conditions, UV and visible light can reduce a va-
riety of metal ions in different environments (Monico et al.,
2015; Oster and Oster, 1959; Thakur et al., 2015). Photore-
duced metal compounds may further act as effective chemi-
cal reductants (Ola and Maroto-Valer, 2015; Xu et al., 2015),
and the oxidized compounds such as hydroxyl radicals or
chlorine atoms can further act as effective oxidants. Za-
maraev et al. (1994) proposed the decomposition of reduc-
ing atmospheric components such as CH4by the photolyt-
ically induced oxidation power of the oxides of iron, tita-
nium and some other metal oxide containing mineral dust
components. Accordingly, Zamaraev designated the dust-
generating deserts of the globe as “kidneys of the earth” (Za-
maraev, 1997) and the atmosphere as a “giant photocatalytic
reactor” where numerous physicochemical and photochem-
ical processes occur (Zamaraev et al., 1994). Researchers
have proposed giant photocatalytic reactors to clean the at-
mosphere of several GHGs, such as N2O (de Richter et al.,
2016b), CFCs and HCFCs (de Richter et al., 2016a) and even
CO2after direct air capture (Kiesgen de Richter et al., 2013),
as almost all GHGs can be transformed or destroyed by pho-
tocatalysis (de Richter and Caillol, 2011; de Richter et al.,
2017).
Oeste (2004) suggested and Wittmer et al. (2015a, b,
2016) and Wittmer and Zetzsch (2016) confirmed the emis-
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8 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
HCl and H2O absorption
by increasing humidity
during nighttime
H2O and HCl desorption
by decreasing humidity
during daytime
HCl absorption by
decreasing HCl content
within the coat during
daytime
FeOOH
solid
Fe(III)
Fe(III)
Fe(II)
Fe(II)
Cl-
Sun radiation at daytime
O2
･Cl
HCl
CH4
･CH3
CO2 O2
Deliquescent salt coat
Figure 2. Simplified chemical reaction scheme of the generation of chlorine radicals by iron salt aerosols under sunlight radiation and the
reaction of the chlorine radicals with atmospheric methane.
sion of CH4-depleting chlorine atoms. This can be induced
in three ways: sunlight photoreduction of Fe(III) to Fe(II)
from FeCl3- or FeOOH-containing salt pans, from FeCl3- or
FeOOH-containing sea spray aerosols and from pure FeOOH
aerosol in contact with air containing parts per billion by vol-
ume amounts of HCl. Because the H abstraction from the
GHG CH4as the first oxidation step by q
Cl is at least 16 times
faster compared to the oxidation by q
OH, which is the main
CH4oxidant acting in the ISA-free atmosphere, the concen-
tration of CH4can be significantly reduced by ISA emission.
Figure 2 illustrates this climate cooling mechanism by the
ISA method with a simplified chemical reaction scheme: a
direct cooling of the troposphere by CH4oxidation induced
by ISA particles.
At droplet or particle diameters below 1 µm (between
1 and 0.1 µm), contact or coagulation actions between the
particles within aerosol clouds are retarded (Ardon-Dryer
et al., 2015; Rosenfeld and Freud, 2011; Santachiara et
al., 2012; Wang et al., 1978). Otherwise the aerosol life-
time would be too short to bridge any intercontinental dis-
tance or arrive in polar regions. This reduces the possi-
ble Cl−exchange by particle contact. But absorption of
gaseous HCl by reactive iron oxide aerosols resulting in
Fe(III) chloride formation at the particle surfaces is possi-
ble (Wittmer and Zetzsch, 2016). Gaseous HCl and other
gaseous chloro-compounds are available in the troposphere:
HCl (300 pptv above the oceans and 100pptv above the con-
tinents) (Graedel and Keene, 1996), ClNO2(up to 1500 pptv
near flue gas emitters) (Osthoff et al., 2008; Riedel et al.,
2014) and CH3Cl (550 pptv, far from urban sources) (Khalil
and Rasmussen, 1999; Yokouchi et al., 2000). By or after
sorption and reactions such as photolysis, oxidation and re-
duction, any kind of these chlorine species can induce chlo-
ride condensation at the ISA particle surface. Acid tropo-
spheric aerosols and gases such as H2SO4, HNO3, oxalic
acid and weaker organic acids further induce the formation of
gaseous HCl from sea-salt aerosol (Drozd et al., 2014; Kim
and Park, 2012; Pechtl and von Glasow, 2007). Since 2004,
evidence and proposals for possible catalyst-like sunshine-
induced cooperative heterogeneous reaction between Fe(II),
Fe(III), Cl−,q
Cl and HCl fixed on mineral dust particles and
in the gaseous phase on the CH4oxidation have been known
(Oeste, 2004; Wittmer and Zetzsch, 2016). Further evidence
of sunshine-induced catalytic cooperation of Fe and Cl came
from the discovery of q
Cl production and CH4depletion
in volcanic eruption plumes (Baker et al., 2011; Rose et
al., 2006). Wittmer and Zetzsch (2016) presented sunshine-
induced q
Cl production by iron oxide aerosols in contact with
gaseous HCl (Wittmer and Zetzsch, 2016). Further evidence
comes from q
Cl found in tropospheric air masses above the
South China Sea (Baker et al., 2015). It is known that the
troposphere above the South China Sea is often in contact
with Fe-containing mineral dust aerosols (∼18 g m−2yr−1)
(Wang et al., 2012), which is further evidence that the Fe-
oxide-containing mineral dust aerosol might be a source for
the q
Cl content within this area.
HCl, water content and pH within the surface layer of
the aerosol particles depend on relative humidity. Both liq-
uid contents, H2O and HCl, grow with increasing humid-
ity (von Glasow and Sander, 2001). In spite of growing HCl
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 9
quantity with increasing humidity, pH increases, due to de-
creasing HCl concentration within the surface layer. Hence,
since the radiation-induced q
Cl production decreases with de-
creasing pH, the q
Cl emission decreases in humid conditions
(Wittmer and Zetzsch, 2016). Under dry conditions, even sul-
fate may be fixed as solid Na-sulfate hydrates. Solubilized
sulfate slightly inhibits the iron-induced q
Cl production (Ble-
icher et al., 2014).
Night or early morning humidity similarly produces the
maximum chloride content on the liquid aerosol particle sur-
face. During daytime, the humidity decrease induces ISA
photolysis and Cl−conversion to q
Cl production by decreas-
ing water content and pH. The ISA particle surface layer
reaches Cl−minima levels during afternoon hours. In the
continental troposphere with a low sea-salt aerosol level,
these effects enable the pure ISA iron oxide aerosol parti-
cles to coat their surface with chloride solution at night and
to emit chlorine atoms in the daytime.
Freezing has different effects on the primary wet ISA par-
ticles. The frozen ice, which is changed to cloud droplets
with solubilized chloride and iron content by CCN action,
is covered by a mother liquor layer with an elevated con-
centration of both iron and chlorine when it arrives in freez-
ing conditions. Some acids such as HCl do not decrease the
mother liquor pH proportional to concentration, and the be-
havior of the ice surfaces, grown from low salt content water,
is different from high salt content water; thus, the different
kinds of ISA behave differently (Bartels-Rausch et al., 2014;
Kahan et al., 2014; Wren and Donaldson, 2012). Direct mea-
surements of molecular chlorine levels in the Arctic marine
boundary layer in Barrow, Alaska, showed up to 400 pptv
levels of molecular chlorine (Liao et al., 2014). The Cl con-
centrations fell to near-zero levels at night but peaked in the
early morning and late afternoon. Liao et al. (2014) estimated
that the Cl radicals oxidized on average more CH4than hy-
droxyl radicals and enhanced the abundance of short-lived
peroxy radicals.
Further investigations have to prove how the different
types of ISA particles behave at different temperatures in
clouds below the freezing point or in the snow layer; these
different types are the primary salt-poor Fe oxide, the poor
hydrolyzed FeCl3and the FeCl3–NaCl mixture because the
q
Cl emission depends on pH, Fe and Cl concentration.
Additional to iron photolysis, in a different and daytime-
independent chemical reaction, iron catalyzes the formation
of q
Cl or Cl2from chloride by tropospheric ozone (Sadanaga
et al., 2001). Triggering the CH4decomposition, both kinds
of iron and chlorine have a cooperative cooling effect on the
troposphere: less GHG CH4in the atmosphere reduces the
greenhouse (GH) effect and allows more outgoing infrared
(IR) heat to the outer space (Ming et al., 2014).
These reactions were active during the glacial period:
Levine et al. (2011) found elevated 13CH4/12 CH4isotope
ratios in Antarctic ice core segments representing the coldest
glacial periods. The unusual isotope ratio is explained by the
Table 1. The Henry’s law constants (Sander, 2015) and daylight
stability for different gaseous or vaporous components reacting with
or produced by ISA in the troposphere.
Substance Henry’s law Stability against
constant tropospheric day-
(mol m−3Pa−1) light (+stable;
−unstable)
CH41.4 ×10−5+
q
Cl 2.3 ×10−2+
Cl29.2 ×10−4−
HCl 1.5 ×101+
HOCl 6.5 −
q
OH 3.8 ×10−1+
H2O28.3 ×102−
much greater q
Cl preference for 12CH4oxidation than 13CH4
oxidation, which is greater than the preference of q
OH. Addi-
tional evidence gives the decreased CH4concentration dur-
ing elevated loess dust emission epochs (Skinner, 2008).
As shown in more detail in Sect. 2.3, ISA produces q
Cl
and much more hydrophilic q
OH and ferryl as other pos-
sible CH4oxidants by Fenton and photo-Fenton processes
(Al-Abadleh, 2015). To gain the optimal reaction conditions
within the heterogeneous gaseous–liquid–solid phase ISA
system in the troposphere, the CH4reductant and the oxi-
dant (Fenton and photo-Fenton oxidant) have to be directed
in such a way that oxidant and reductant can act within the
identical medium.
As seen in Table 1, according to the CH4Henry’s law con-
stant the preference of the 1.8 ppm tropospheric CH4is un-
doubtedly the gaseous phase. q
Cl has also a preference for
the gaseous phase.
Iron exists at least in part as Fe(III) during the nighttime
and at least in part as Fe(II) during daytime. The CH4oxi-
dation by q
Cl and q
OH is restricted to the daytime as during
night hours q
Cl and q
OH recombine fast to Cl2, HOCl and
H2O2in the dark (von Glasow, 2000). During daylight hours,
these recombination products photolyze again by the regen-
eration of the radicals. But even during daytime these radi-
cals and their recombination products coexist due to the cy-
cling between q
Cl, q
OH, Cl2, HOCl and H2O2. This cycling
is activated by sunlight photolysis and radical recombination
reactions (Luna et al., 2006; von Glasow, 2000).
As we learn from the Henry’s law constants in Table 1, the
oxygen species q
OH and H2O2have a much higher tendency
to stay in the liquid phase than the chlorine species q
Cl and
Cl2. Cl2has the tendency to react with water of neutral pH by
producing HOCl. But the pH values of ISA, especially if ISA
is emitted as acid flue gas plumes, are lower than 3. Within
this acidic region the tendency of HOCl generation from Cl2
decreases to very low values, and even at those humidity lev-
els at which the ISA particles become deliquescent, the ma-
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10 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
jority of the activated chlorine species will be localized in the
gaseous phase containing CH4, not in the liquid phase.
But q
OH may leave the condensed phase in favorable cir-
cumstances and enter the gaseous phase (Nie et al., 2014) and
may contribute there to the oxidation of CH4during clear dry
conditions without liquid phase at the Fe(III) surfaces.
Water-soluble ammonia (5.9 ×10−1) has a similar
Henry’s law constant to q
OH . Therefore, q
OH has the ten-
dency to stay within hydrous phases during humid condi-
tions. This tendency is 16 times lower for q
Cl. This property
is combined with the 16 times higher reactivity in compari-
son to q
OH. At an equal production of q
Cl and q
OH, the reac-
tion of q
Cl with CH4has a probability of up to 250 times
(16 ×16) that of the reaction of q
OH with CH4when the
ISA particles are wet and 16 times that of q
OH with CH4
when the ISA particles are dry. The probability of CH4oxi-
dation by ISA-derived q
Cl against ISA-derived q
OH may be
restricted by the pH increase tendency within ISA during hu-
mid episodes (decreased q
Cl generation on ISA with rising
pH) to values fluctuating between the extremes 1 and 250. In-
dependently of the kind of oxidants produced by ISA – dur-
ing dry, clear-sky and sunshine episodes – the ISA-derived
oxidants produce maximum oxidant concentrations within
the CH4-containing gaseous phase, producing optimum CH4
depletion rates.
The q
Cl reactivity on most VOCs other than CH4is at
least 1 order of magnitude higher than that of q
OH (Young et
al., 2014). Halogen organics such as dichloromethane (Pena
et al., 2014) as well as the environmentally persistent and
bioaccumulating perfluoro organics such as perfluorooctane
sulfonate may be depleted by sunlit ISA (Jin et al., 2014).
2.3 Oxidation of organic aerosol particles containing
black and brown carbon
Black carbon in soot is the dominant absorber of visible solar
radiation in the atmosphere (Ramanathan and Carmichael,
2008). Total global emission of black carbon is 7.5 Mt yr−1
(Bond et al., 2013). Direct atmospheric forcing of atmo-
spheric black carbon is +0.7 W m−2(Bond et al., 2013). In
addition to its climate relevance, black-carbon soot induces
severe health effects (Anenberg et al., 2012).
Andreae and Gelencsér (2006) defined the differences be-
tween the carbons: black carbon contains insoluble elemental
carbon; brown carbon contains at least partly soluble organic
carbon. Black carbon also contains additional extractable or-
ganics of greater or lesser volatility and/or water solubility
(Andreae and Gelencsér, 2006; Nguyen and Ball, 2006).
Black and brown carbonaceous aerosols have a positive
radiative forcing (warming effect) on clouds (Ramana et al.,
2010) as seen in Sect. 2.1 and also after deposition on snow,
glaciers, sea ice or in the polar regions, as the albedo is re-
duced and the surface is darkened (Hadley and Kirchstetter,
2012). One of the most effective methods of slowing global
warming rapidly in the short term is by reducing the emis-
sions of fossil fuel particulate black carbon, organic matter
and tropospheric ozone (Jacobson, 2002).
Both aerosol types have adverse effects on health (hu-
man, animal, livestock, vegetal), and reducing their levels
will save lives and provide many benefits (Shindell et al.,
2012). Thus, any tropospheric lifetime reduction in both dark
carbons would lead to cooling effects and further positive ef-
fects.
Both carbons are characterized by aromatic functions.
The black carbons contain graphene structures; the brown
ones have low-molecular weight humic-like aromatic sub-
stances (HULISs). HULISs derive from tarry combus-
tion smoke residues and/or from aged secondary organic
aerosol (SOA). The sources of SOA are biogenic VOCs such
as terpenes (Fry et al., 2014). HULISs contain polyphenolic
redox mediators such as catechol and nitro-catechols (Claeys
et al., 2012; Hoffer et al., 2004; Ofner et al., 2011; Pillar et
al., 2014).
The polyphenolic HULIS compounds are ligands with a
very strong bond with iron. Rainwater-dissolved HULISs
prevent Fe(II) from oxidation and precipitation when mix-
ing with seawater (Willey et al., 2008). Wood-smoke-derived
HULIS nanoparticles penetrate into living cell walls of res-
piratory epithelia cells. After arrival in the cells the HULIS
particles extract the cell iron from the mitochondria by for-
mation of HULIS–iron complexes (Ghio et al., 2015).
Beside iron, other metals such as manganese and copper
have oxygen transport properties which improve the oxida-
tion power of H2O2by Fenton reactions generating q
OH
(Chemizmu and Fentona, 2009). H2O2is a troposphere-
borne oxidant (Vione et al., 2003).
Polyphenolic and carboxylate ligands of HULIS enhance
the dissolution of iron oxides. These ligands bind to undis-
solved iron oxides (Al-Abadleh, 2015).
Iron and catechols are both reversible electron shuttles:
Fe2+↔Fe3++e;(1)
catechol ↔quinone +2e. (2)
The HULIS–iron connection enhances the oxidative degra-
dation of organic compounds such as aromatic compounds
(Al-Abadleh, 2015).
Oxidant generation by the reaction of oxidizable dissolved
or undissolved metal cations such as Fe(II), Cu(I) and Mn(II)
with H2O2was first discovered for Fe(II) in 1894 (Fenton,
1894). Since then these reactions have been known as Fenton
reactions. The mechanisms and generated oxidants of Fenton
reactions are still under discussion.
Depending on the participating metal ligand oxidants such
as q
OH, Fe(IV)O2+(=Ferryl), q
Cl, q
SO−
4, organic peroxides
and quinones may appear (Barbusi´
nski, 2009).
According to Barbusinsky (2009) the primary reac-
tion intermediate from Fe2+and H2O2is the adduct
{Fe(II)H2O2}2+, which is transformed into the fer-
ryl complex {Fe(IV)(OH)2}2+. The latter stabilizes as
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 11
Carbon soot particles
oxidized by
Fenton & photo-Fenton
oxidants become
hydrophilic
H2O2
Coagulation
ISA particle
coated by
iron salts Rainout
Hydrophobic
fresh
carbon soot
particles
ISA particle
coated by
iron salts
ISA particle
coated by
iron chloride
Smoke particles
oxidized by
Fenton & photo-Fenton
oxidants become
hydrophilic
H2O2
Hydrophobic
fresh
smoke
particles
Rainout
･Cl
Tropospheric
O₃･Cl
･CH3
CH4
HCl
O₂
Coagulation
Sunshine
Sunshine
Figure 3. Schematic representation of the cooling of the troposphere by inducing the decrease in ozone and organic aerosol particles such
as soot and smoke.
{Fe(IV)O}2++H2O. Reductants may also react di-
rectly with {Fe(IV)O}2+or after its decomposition to
Fe3++q
OH +OH−by q
OH. Fe3+reacts with H2O2
to Fe2+via q
O2H development; the latter decays into
O2+H2O.
Light enhances the Fenton reaction effectiveness. It re-
duces Fe3+to Fe2+by photolysis inducing q
OH or q
Cl gen-
eration, the latter in the case of available Cl−, which reduces
the H2O2demand (Machulek Jr. et al., 2009; Southworth and
Voelker, 2003).
This process is illustrated by Fig. 3.
The Fenton reaction mechanism is dependent on pH and
on the kinds of ligands bound to the Fenton metal. The reac-
tion mechanism with oxidants of SO2−
4, NO−
3, Cl−and 1,2-
dihydroxy benzene ligands has been studied (De Laat et al.,
2004).
In biological systems, 1,2-dihydroxy benzenes (cate-
cholamines) regulate the Fenton reaction and orient it toward
different reaction pathways (Salgado et al., 2013).
Additionally, the fractal reaction environments like
surface-rich black and brown carbons and ISA have con-
siderable influence on the Fenton reaction. By expanding
the aqueous interface, accelerations of the reaction veloc-
ity up to 3 orders of magnitude has been measured (Enami
et al., 2014). This may be one of the reasons why iron-
containing solid surfaces made of fractal iron oxides, pyrite,
activated carbon, graphite, carbon nanotubes, vermiculite,
pillared clays and zeolites have been tested as efficient Fen-
ton reagents (Pignatello et al., 2006; Pinto et al., 2012; Teix-
eira et al., 2012).
Even the oxidation power of artificial Fenton and photo-
Fenton systems is known to be high enough to hydroxy-
late aliphatic C–H bonds, inclusing CH4hydroxylation to
methanol (Gopakumar et al., 2011; Hammond et al., 2012;
Yoshizawa et al., 2000).
But the HULIS itself becomes depleted by the Fenton ox-
idation when it remains as the only reductant (Salgado et al.,
2013).
Like HULIS or humic substances, the different kinds of
black carbons act as redox mediators due to their oxygen
functionalities bound to the aromatic hexagon network such
as hydroxyl, carbonyl and ether (Klüpfel et al., 2014; Oh
and Chiu, 2009). These functionalities similarly act as hy-
droquinone, quinone, aromatic ether, pyrylium and pyrone in
the extended graphene planes as electron acceptors and donor
moieties. Soot also possesses such redox mediator groups
(Drushel and Hallum, 1958; Studebaker et al., 1956). Again
these are ligands with well-known binding activity on iron
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12 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
compounds. The difference between them and the HULIS
ligands is that the former are attached to stacks of aromatic
graphene hexagon networks instead of mono- or oligo-cyclic
aromatic hexagons of HULIS. As well as the HULIS redox
mediator ligands these hydroxyl and ketone groups transfer
electrons from oxidants to reductants and vice versa. Like the
HULIS–iron pair, the black-carbon–iron pair enhances the
redox mediation above the levels of every individual electron
shuttle (Kim et al., 2013; Lima et al., 2013; L. Wang et al.,
2014). Accordingly, any ISA doping of black carbons gen-
erates effective oxidation catalysts (Oeste, 1977; Song et al.,
2015).
Lit by sunlight the ISA-doped soot represents an oxidation
catalyst to adsorbed organics, producing its own oxidants by
the photo-Fenton reaction. In spite of the higher chemical sta-
bility of the graphene network of soot compared to HULIS
soot, by wet oxidation other oxygen groups are fixed to the
soot graphene stacks (Moreno-Castilla et al., 2000) increas-
ing soot’s hydrophilic property, which is necessary to arrange
its rainout. The hydroxyl radical attack resulting from the
photo-Fenton reaction ultimately breaks the graphene net-
work into parts (Bai et al., 2014; Zhou et al., 2012). Photo-
Fenton is much more efficient in q
OH generation than Fen-
ton because Fe(III) reduction as a regeneration step occurs
by Fe(III) photoreduction rather than consuming an organic
reductant.
The oxidized hydrophilic carbon particles are more read-
ily washed out of the atmosphere by precipitation (Zuberi
et al., 2005). ISA accelerates this oxidation process as the
iron-induced Fenton and photo-Fenton reaction cycles pro-
duce hydroxyl and chlorine radical oxidants, speeding up the
soot oxidation.
Fe(III) forms colored complexes with hydroxyl and car-
boxylic hydroxyl groups too, particularly if two of them are
in 1,2 or 1,3 position, such as in oxalic acid. The latter belong
to the group of dicarboxylic acids known to be formed as ox-
idation products from all kind of volatile, dissolved or par-
ticular organic carbons in the atmosphere (Kawamura et al.,
2003). Dicarboxylate complexes with iron are of outstand-
ing sensitivity to destruction by photolyzation (Eder, 1880,
1906; Weller et al., 2014; Zhu et al., 1993): photolysis re-
duces Fe(III) to Fe(II) by producing H2O2and oxidation of
the organic complex compounds. Then Fe(II) is reoxidized
to Fe(III) by H2O2in the Fenton reaction by the generation
of q
OH (Cunningham et al., 1988). Due to their elevated po-
larity, oxidation products containing hydroxyl and carboxyl
groups have increased wettability, are more water soluble and
are thus rapidly washed out from the atmosphere.
Due to their elevated reactivity compared to CH4, the
gas phase oxidation of airborne organic compounds by ISA-
generated q
OH or q
Cl is enhanced. By eliminating black
and brown carbon aerosols, ISA contributes to global warm-
ing reduction and to decreasing polar ice melting by sur-
face albedo reduction caused by black-carbon snow contam-
ination (Flanner et al., 2007; Ramanathan and Carmichael,
2008).
The generation of ISA by combusting fuel oil with fer-
rocene or other oil-soluble iron additives in ship engines or
heating oil burners has additional positive effects because
soot is catalytically flame-oxidized in the presence of flame-
borne ISA (detailed in Sect. 6) as a combustion product of
the iron additive (Kasper et al., 1998; Weiser et al., 2007).
2.4 Tropospheric ozone depletion by ISA
An additional GHG is the tropospheric ozone (Jacobson,
2002). Carbon dioxide is the principal cause of GW and
represents two-thirds of the global radiative forcing, but
long-lived methane and short-lived tropospheric ozone are
both GHGs and, respectively, responsible for the second and
third most important positive radiative forcing.
According to Blasing (Blasing, 2010, 2016; Forster et
al., 2007), tropospheric O3has an atmospheric forcing of
+0.4 W m−2. Any direct depleting action of tropospheric O3
by the ISA-induced q
Cl is accompanied by an indirect emis-
sion decrease of O3as the reduction of CH4and other VOC
by the ISA method decreases the O3formation (Cooper et
al., 2014).
Reactive halogen species (mainly Cl, Br) cause strato-
spheric ozone layer destruction and thus the “ozone layer
hole”. Tropospheric ozone destruction by reactive halogen
species is also a reality (Sherwen et al., 2016). For a long
time now, q
Cl and q
Br have been known as catalysts for O3
destruction in the stratosphere (Crutzen and Oppenheimer,
2008). Investigations both in the laboratory and in nature
have shown that q
Br is a much more effective catalyst of
ozone depletion within the troposphere than q
Cl (Le Bras and
Platt, 1995; Liao et al., 2014; Wayne et al., 1995).
As discussed at the end of Sect. 2.6, clear evidence exists
that the ozone-depleting “bromine explosions” known as reg-
ular phenomena developing from coastal snow layers at sun-
rise in the polar spring (Blechschmidt et al., 2016; Pratt et
al., 2013) are likely to be induced by the photolyzed precip-
itation of iron-containing dust. According to Pratt, bromide-
enriched brines covering acidified snow particles are oxi-
dized by photolyzation to q
Br.
In coastal areas of both the northern and southern po-
lar regions during springtime, inert halide salt ions (mainly
Br−) are converted by photochemistry into reactive halo-
gen species (mainly Br atoms and BrO) that deplete ozone
in the boundary layer to near-zero levels (Simpson et al.,
2007). During these episodes, called “tropospheric ozone de-
pletion events” or “polar tropospheric ozone hole events”, O3
is completely destroyed in the lowest kilometer of the atmo-
sphere in areas of several million square kilometers; this has
a negative climate feedback or cooling effect (Roscoe et al.,
2001).
In the tropics, halogen chemistry (mostly Br and I) is
also responsible for a large fraction (∼30 %) of tropo-
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 13
spheric ozone destruction (Read et al., 2008; Sommariva and
von Glasow, 2012) and up to 7 % of the global methane de-
struction is due to chlorine (Hossaini et al., 2016; Lawler
et al., 2009). It has been estimated that 25 % of the global
oxidation of CH4occurs in the tropical marine boundary
layer (Bloss et al., 2005). A one-dimensional model has been
used to simulate the chemical evolution of air masses in the
tropical Atlantic Ocean (Sommariva and von Glasow, 2012)
and to evaluate the impact of the measured halogens lev-
els. In this model, halogens (mostly Br and I) accounted for
35–40 % of total tropospheric O3destruction while the Cl
atoms accounted for 5.4–11.6 % of total CH4sinks. Sher-
wen et al. (2016) estimate the radiative forcing reduction
due to O3from the preindustrial to the present day as being
−0.066 W m−2.
The ISA-induced increase in q
Br concentration at sea-
salt-containing tropospheric conditions has been confirmed
(Wittmer et al., 2015b). This establishes ISA as part of an
ozone-depleting reaction cycle and additional cooling stage.
This depletion effect of the GHG tropospheric ozone is worth
noting.
2.5 ISA-induced phytoplankton fertilization albedo
increase (by enhancing dimethylsulfide (DMS)
emissions) and CH4oxidation efficiency (by
increasing methyl chloride and DMS emissions)
One of the largest reservoirs of gas phase chlorine is the
ca. 5 Tg of methyl chloride (MC) in the Earth’s atmo-
sphere (Khalil and Rasmussen, 1999). Methyl chloride is re-
leased from phytoplankton (Hu et al., 2013) and from coastal
forests, terrestrial plants and fungi (Khalil et al., 1999).
Dimethylsulfide (DMS) is a volatile sulfur compound that
plays an important role in the global sulfur cycle. Through
the emission of atmospheric aerosols, DMS may control cli-
mate by influencing cloud albedo (Charlson et al., 1987).
Currently, researchers (Lana et al., 2011) estimate that
28.1 (17.6–34.4) Tg of sulfur in the form of DMS are trans-
ferred annually from the oceans into the atmosphere.
Ocean acidification has the potential to exacerbate anthro-
pogenic warming through reduced DMS emissions (Six et
al., 2013). By contrast, increased emissions of DMS and MC
into the troposphere are a consequence of the ISA-induced
phytoplankton growth and DMS+MC release into the tro-
posphere. DMS is oxidized in the troposphere to sulfuric and
sulfonic acid aerosols, which are highly active CCN. This
process enhances the direct ISA cooling effect as described
in Sect. 2.1 (Charlson et al., 1987).
Upon contact of this acidic aerosol with sea spray aerosol,
sulfate and sulfonate aerosols are formed and gaseous HCl
is produced. Sulfate aerosols are known to have a negative
radiative forcing (a cooling effect) (Crutzen, 2006).
Another HCl source is the oxidation of MC. Both effects
induce the tropospheric HCl level to rise. Due to the cooling
stage described in Sect. 2.2, with the increased HCl level, ad-
ditional chlorine atoms are produced by reaction with ISA.
This effect further accelerates the CH4oxidation and its re-
moval from the atmosphere, reducing its radiative forcing.
2.6 Oxidation of CH4and other GHGs by sunlit solid
surfaces
Mineral aerosol particles adhere strongly to sunlit, dry and
solid surfaces of rocks and stones. A well-known remnant of
the dust deposition in rock or stone deserts and rocky semi-
arid regions is the orange, brown, red or black “desert var-
nish” coat covering stones and rocks. The hard desert var-
nish is the residue, which is glued together and hardened, of
the primary dust deposit. Daily sun radiation and humidity
change, as well as microbe and fungi influence, builds up
the varnish, changing the primary aerosol deposit (Perry et
al., 2005) by photolytic Fe(III) and Mn(IV) reduction during
daytime and nighttime oxidation of Fe(II) and Mn(II). The
oxidation is triggered further by Mn- and Fe-oxidizing mi-
crobes adapted to this habitat (Allen et al., 2001; Hungate et
al., 1987). Desert varnish preserves the Fe and Mn photore-
duction ability of the aerosol: lit by light the varnish can pro-
duce chlorine from chloride-containing solutions (Johnson
and Eggleston, 2013). The photo-, humidity- and microbially
induced permanent Fe and Mn valence change between night
and day (Matsunaga et al., 1995); accompanied by adequate
solubility, changes seem to trigger the physicochemical hard-
ening of every new varnish layer.
The varnish is composed of microscopic laminations of
Fe and Mn oxides. Fe plus Mn represent about one-fifth
of the varnish. Meanwhile, four-fifths of the laminations
are composed of SiO2, clay and former dust particles. The
dominant mineral is SiO2and/or clay (Dorn, 2009; Liu and
Dorn, 1996). There is little doubt that desert varnish can
build up even from pure iron oxides or iron chloride aerosol
deposits such as ISA. The optimum pH to photo-generate
the methane-oxidizing chlorine atoms from ISA is pH 2
(Wittmer et al., 2015a). Established by the gaseous HCl con-
tent of the troposphere (Graedel and Keene, 1996), a pH drop
to pH 2 at the varnish surface is possible on neutral alkaline-
free surfaces such as quartz, quartzite and sandstone. The
humidity-controlled mechanism acting between gaseous HCl
and HCl dissolved in the liquid water layer absorbed on the
solid iron oxide surface of ISA particles, as explained in Sect.
2.2, acts at the varnish surface in a similar manner to an ISA
particle surface: an FeCl3stock can pile up by Fe(II) oxida-
tion and humidity-triggered HCl absorption during the night-
time. The FeCl3stock at the varnish surface is consumed dur-
ing daytime by photolytic Fe(II) and chlorine atom genera-
tion.
ISA aerosol particles emit HCl during dry conditions. Like
oxidic ISA, desert varnish absorbs H2O and HCl from the
atmosphere, gathering it during the nighttime as a surface-
bound H2O, OH−and Cl−coating. During sunlit daytime,
chloride and water desorbs from Fe(III) as q
Cl, q
OH and
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14 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
b
d
g
e
c
a
f
j
i
h
Figure 4. Schematic representation of iron salt aerosol interactions with different solid surfaces: primary ISA precursor FeOOH particles (a)
react with gaseous HCl by generation of ISA as FeCl3coated on FeOOH particles (c). Coagulation, condensation and chemical reaction with
particles and vapors produce different kinds of liquid and/or solid ISA variants and sediments: (b) hydrolyzed FeCl3coated on soot and/or
HULIS particles; (d) hydrolyzed FeCl3coated on ice crystals; (e) hydrolyzed FeCl3coated on salt crystals; (f) hydrolyzed FeCl3coated on
ice crystals of snow layers (ISA sediment); (g) hydrolyzed FeCl3dissolved in cloud droplets; (h) FeCl3hydrolysate residue on desert varnish
(ISA sediment); (j) hydrolyzed FeCl3as dissolved residue in ocean surface water fertilizes the phytoplankton growth and ultimately triggers
the generation of sulfuric, sulfonic and dicarboxylic acids by emission of DMS, MC and other organics. This activates the tropospheric
generation of vaporous HCl by reaction of sea-salt aerosol (i) with the acids. HCl again changes the ISA precursor FeOOH aerosol (a) to
ISA (c).
H2O, leaving Fe(II) in the varnish surface. The surface Fe(II)
(and Mn(II)) is bound by oxygen bridges to the varnish bulk
of Fe(III) (and Mn(IV)), perhaps like the combination of
Fe(II) and Fe(III) within magnetite. During the nighttime the
Fe(III) (and Mn(IV)) surface coating is regenerated by mi-
crobial and/or abiotic oxidation with O2. It is worth men-
tioning that desert varnish can exist only within dry regions.
Figure 4 illustrates the interactions of ISA at the phase
borders of tropospheric aerosols, ocean surface and dry solid
surfaces.
Similar daytime-dependent microbially activated abi-
otic photoreduction and photooxidation reaction cycles are
known from aquifer environments (Gammons et al., 2007).
Thus, the CH4depletion of the former ISA deposits will per-
sist even after change into desert varnish. As explained in
Sect. 2.2, continental HCl (300 pptv above the oceans and
100 pptv above the continents) (Graedel and Keene, 1996),
ClNO2(up to 1500 pptv near flue gas emitters) (Osthoff et
al., 2008; Riedel et al., 2014) and CH3Cl (550 pptv, far from
urban sources) (Khalil and Rasmussen, 1999; Yokouchi et
al., 2000) and, in deserts, chloride salt-containing dusts are
direct and indirect sources of chloride which could provide
desert varnishes with Cl−.
Furthermore, analogously to ISA deposited on solid desert
surfaces, ISA depositions on dry snow, snow cover and ice
occurring in permanently snow-covered mountain regions or
within polar and neighboring regions preserve its CH4de-
struction activity during sunlit days in spring and summer
(Liao et al., 2014).
The global area of the desert varnish surface does not
change with changing dust precipitation rates. It only de-
pends on the precipitation frequency. It grows through deser-
tification and shrinks with increasing wet climate. Until now,
quantitative measurements about the specific amount of CH4
depletion per square meter of desert varnish are not known.
Without these data, an estimation of its influence on the CH4
depletion and climate is impossible.
The photochemical actions inducing CH4depletion of the
desert varnish surfaces resulting from dust precipitation are
concurrent with the surfaces of deserts and semideserts made
of sand or laterite soils. Their surface is colored by ochre-to-
red iron oxide pigments. Their iron components should act in
principle by the same CH4-depleting photochemistry as ISA
and desert varnish.
As mentioned in Sect. 2.4, the Cl and Br activation by iron
photolysis changes after the separation of the ingredients,
by freezing or drying of the formerly homogeneous liquid,
into solid salt-poor ice and a liquid brine coating or solid salt
and a liquid brine coating. This inhomogeneous partitioning
phenomenon of the predominant transformation of aerosol
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 15
droplets into solid states, and vice versa, applies to snow or
salt layers containing a proportion of ISA.
It has been shown that by cooling the precipitation of
buffering salts (such as carbonates, sulfates and chlorides of
bromide- and chloride-rich mother liquors) onto Arctic snow
packs or ice particles can minimize their buffering capac-
ity against pH change (Bartels-Rausch et al., 2014; Blech-
schmidt et al., 2016; Sander et al., 2006). Similar mecha-
nisms may act when liquid aerosol particles become solid
by drying.
Then, the uptake and contact over time of solid iron-
bearing particles and airborne organic and inorganic acids
and acid precursors on, or with, ice crystal surfaces may
lower the pH of the formerly alkaline particle surface, to the
reaction conditions of bromide oxidation by iron(III) pho-
toreduction.
According to Kim et al. (2010), the photogeneration of
Fe(III) oxides, proceeding slowly at pH 3.5 in bulk solution,
becomes significantly accelerated in polycrystalline Arctic
ice. This effect is accompanied by an acceleration of the
physical dissolution of the Fe(III) oxides by freezing ice
(Jeong et al., 2012; Kim et al., 2010).
The contact of Arctic snow layers with iron oxides is con-
firmed by Kim et al. (2010). Dorfman et al. (2015) found
recent loess dust sedimentation rates in the Alaskan Arctic
Burial Lake of 0.15 mm yr−1. According to research results
from artificial iron-doped salt pans (Wittmer et al., 2015b),
iron-salt-doped sea-salt aerosols (Wittmer et al., 2015a) or
sea-salt-doped iron oxide aerosols or pure iron oxide aerosols
in contact with gaseous HCl (Wittmer and Zetzsch, 2016)
chloride and bromide in sunlit surfaces are oxidized to q
Cl
and q
Br by photoreduced Fe(III) if the pH of the reaction
media is 3.5 or lower.
As known from the bromine explosions, they appear on
acidified first-year tundra and first-year sea ice snow, lit by
sunlight (Pratt et al., 2013). According to Kim et al. (2010)
and Dorfman et al. (2015) the year-old snow layers con-
tain iron(III). This confirms that sufficient reaction condi-
tions exist to produce bromine explosions by the oxidation
of iron(III) photoreduction.
Continents have considerable areas where the outflowing
water is drained into “endorheic” water bodies and not into
the oceans. Endorheic lakes have no outlets other than evap-
oration, and thus dissolved salts and nutrients concentrate
over time. Large surfaces of these basins are covered by salt
crusts, salt marshes, salty soils or salt lakes. Most of these
areas are situated within desert or semidesert areas (Ham-
mer, 1986). These salt environments gain iron from precipi-
tating dust or from iron-containing brines they have precip-
itated from. As soon as these environments become acidic,
they oxidize CH4by iron photolysis-induced q
Cl (Wittmer et
al., 2015b).
To summarize the climate-relevant action of ISA within
the troposphere as described in Sects. 2.1–2.6: CH4, VOC,
O3and dark carbon aerosol plus cloud albedo, in sum, have
a similar effect on the climate warming to CO2. With ISA
method significant reductions in CH4, VOC, O3are antici-
pated based on the test results from Wittmer et al. (2015a, b,
2016) and Wittmer and Zetzsch (2016), and significant re-
ductions in dark carbon aerosol and a significant increase in
cloud albedo are anticipated based on the literature cited. We
found no arguments against these statements. This allows the
conclusion that the ISA method should have significant cli-
mate cooling effects only within the troposphere.
3 Oceanic natural cooling effects of the iron cycle
3.1 Biotic CO2conversion into organic and carbonate
carbon
Vegetation uses the oxidative power of organic metal com-
pounds induced by photon absorption, oxidizing water to
oxygen and reducing CO2by organic carbon generation
(photosynthesis by chlorophyll, a green Mg–porphyrin com-
plex). This assimilation process is retarded by prevailing iron
deficiency in the oceans, which slows phytoplankton growth.
However, there is no doubt that ISA-containing dust pre-
cipitation fertilizes the phytoplankton, which in turn affects
the climate (Albani et al., 2016).
ISA triggers the phytoplankton reproduction and in-
creases the formation of organic carbon from the GHG CO2
(Martínez-García et al., 2014). The vast majority of the
oxygen thus formed, which is only slightly water soluble
(11 mg O2L−1), escapes into the atmosphere. In contrast,
the organic carbon formed remains completely in the ocean,
forming the basis of the marine food and debris chain.
Of the primary produced phytoplankton carbon, only a
small fraction arrives at the ocean bottom as organic debris
and becomes part of the sediment. Cartapanis et al. (2016)
and Jaccard et al. (2016) found direct evidence that during
the glacial maxima, the accumulation rate of organic carbon
was consistently higher (50 %) than during interglacials. This
resulted from the high dust concentrations during the glacial
maxima, fertilizing the phytoplankton with ISA.
The buildup of Ca carbonate shell and frame substances
by the calcification process at the ocean surface extracts ad-
ditional CO2C from the troposphere. The bulk of calcifi-
cation can be attributed to corals, foraminifera and coccol-
ithophores; the latter are believed to contribute up to half of
current oceanic CaCO3production (Mackinder et al., 2010).
Both carbon fixation processes increase the removal of the
GHG CO2and thus contribute to cooling the troposphere.
The Fe-fertilizing process was in progress during the ice
ages, as the evaluations of Antarctic ice cores show: the
minimum CO2concentrations and temperatures in the tropo-
sphere are connected to the high dust phases (Skinner, 2008).
It has been discussed that the alkalinity loss by phyto-
plankton calcification and CaCO3loss with phytoplankton
debris from the ocean surface is said to produce a calcium
and alkalinity deficit at the ocean surface (Meyer and Riebe-
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16 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
sell, 2015; Rost and Riebesell, 2004), producing additional
acidification at the ocean surface by CO2generation:
Ca(HCO3)2→CaCO3+H2O+CO2.(3)
At least in part, this acidification is compensated for by as-
similative generation of organic carbon by CO2consump-
tion. Both organic debris and CaCO3become part of the
ocean sediment. But if the organic debris is reoxidized during
its journey downwards, some acidification could result from
this. Acidification could result from this too if more CO2is
absorbed by the ocean than is assimilated and changed to
organic debris. Sedimentation of organic debris and CaCO3
both increase, depending on the ISA-induced phytoplankton
productivity.
The increasing amount of CaCO3sedimentation within
iron-fertilized ocean regions has been discussed by Salter et
al. (2014). In a sufficiently mixed ocean, alkalinity loss at the
surface is more than compensated for by the different sources
of alkali and earth alkali cations at the ocean bottom and
through continental weathering: in the first place these are
the mechanisms of alkalinity generated by the ocean water
reactions within the ocean sediments and their bedrock, the
oceanic crust. The latter mechanisms are described in more
detail in Sects. 4.1–4.3. The convection of the primary oxic
ocean bottom water through the ocean crust generates alka-
linity by the reduction of sulfate, nitrate and hydrogen car-
bonate, by the dissolution of silicates by reduced humic acids
and further by the serpentinization of basalt and peridotite
silicates (Alt and Shanks, 2003; Früh-Green et al., 2004). The
alkalinity extracted from the oceanic crust mainly remains
positioned in the dark water layers of the ocean basins if the
decreased THC is not able to elevate the alkaline extract into
the phytoplankton layer in sufficient quantities.
The THC activation by the ISA method is described in
Sects. 4.1–4.3.
Sudden ISA-induced phytoplankton growth generates in-
creased calcite-shell production. This lowers the Ca concen-
tration at the ocean surface. Even if the vertical cycling is
not fast enough to compensate for the Ca loss at once, or
after a small time lag, this does no harm to phytoplankton
growth because Ca is not essential to it. Just the opposite is
true: phytoplankton use calcification as a detoxification mea-
sure to get rid of calcium ions from their bodies (Müller et
al., 2015). As a consequence of this effect only the relation
between Ca carbonate sequestration and organic carbon se-
questration will decrease during the time lag.
By additional direct alkalinity production of the phyto-
plankton itself, at least parts of the acidity from lime shell
production may be compensated for: ISA-controlled phyto-
plankton growth induces an increased synthesis of organic
sulfur and of chlorine compounds (Matrai and Keller, 1994),
emitted as DMS and MC (Carpenter et al., 2012). Synthe-
sis of organic sulfur and halogen organics, as precursors of
the volatile DMS and MC emission, is realized by phyto-
plankton, by a reduction of sulfate to organic sulfides and by
the oxidation of chloride to carbon chlorine compounds. This
precursor synthesis excretes equivalent Na+and/or Ca2+al-
kalinity, as Na2SO4reduction/formation to DMS generates
Na alkalinity; NaCl oxidation/formation to MC also gener-
ates Na alkalinity: cations formerly bound to SO2−
4or Cl−
lose their anions, producing alkalinity. According to Chen
et al. (1996) and Fujita (1971) the sulfur content of phyto-
plankton exclusively exceeds the Ca2+, Mg2+and K+alka-
line load of phytoplankton lost with the phytoplankton de-
bris. Only half of the organic carbon assimilated by phyto-
plankton derives from dissolved CO2. The other half derives
from the ocean water NaHCO3anion content (Cassar et al.,
2004). Chemical reduction (the reduction of HCO−
3to or-
ganic C +O2by the assimilation of HCO−
3anions) produces
alkalinity as another compensation of the alkalinity loss by
calcification. NaHCO3reduction/formation to organic car-
bon generates Na alkalinity. The cation previously bound to
HCO−
3loses its anion and produces alkalinity.
These considerations demonstrate that any of the proposed
enhanced weathering measures to prevent ocean acidification
by increasing the alkalinity (Taylor et al., 2016) might not
be necessary if the ISA method is in action and keeps the
vertical ocean mixing sufficiently active.
During the down-dripping of the very finely shaped phy-
toplankton debris, bacterial oxidation, fish and other food
chain links minimize the organic debris by up to 1 order
of magnitude (Weber et al., 2016). Even the dissolution of
the small carbonate debris reduces the carbonate fraction un-
til it arrives at the sediment surface. In order to maximize
the effect of the ISA method, within the main ISA precip-
itation regions, the oxidation and dissolution of the organic
and carbonate phytoplankton debris during its dripping down
through the ocean water column can be reduced. To reach this
goal, we suggest farming fixed filter feeders such as mussels
and oysters within the ISA precipitation region.
Mussels and oysters produce faeces and so-called “pseudo
faeces” in the shape of fairly solid pellets. Compared to the
time of the sedimentation of the unconditioned phytoplank-
ton debris, this expands the sedimentation time difference
between excreted filter feeder faeces and the phytoplankton
faeces pellet sedimentation on the ocean floor by 1 order of
magnitude. Bivalve farming would significantly reduce the
oxidative and solution loss by phytoplankton debris attack.
Mussel and oyster farming are well-known practices which
have been employed for a long time as a measure to produce
protein-rich food. They have been proposed as an element of
climate engineering (Dimitrova et al., 2015; Lenton and Sen
Gupta, 2010).
To further optimize the CO2-C conversion to sediment-
bound C the biomass of oysters and mussels, including their
shells and fixing systems, might be periodically dumped into
the sediment.
Additional floating supports such as coral habitats,
sponges, sea lilies and sea anemones between the mussel
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 17
supports might complete and again optimize the ISA pre-
cipitation areas. The oceanic water deserts can be changed
into productive ecosystems and protein sources for an in-
creasing population by these measures, among others, for an
optimized CO2fixation induced by ISA.
A further proposal in order to maximize the CO2fixation
induced by ISA is our suggestion to integrate a solution to
the plastic waste problem on the ocean surfaces into the ISA
method. About 5 to 13 million metric tons of solid plastic
waste per year are entering the oceans (Jambeck et al., 2015).
Over the last years the plastic waste drifting on the ocean has
developed into a huge problem for oceanic ecosystems (Law
et al., 2014). Plastic keeps sunlight away from phytoplank-
ton, hampering it from effective growth. The plastic waste
drifts with the ocean currents. It then collects within accumu-
lation zones predicted by a global surface circulation model
(Cózar et al., 2014). Most plastic-covered ocean surfaces are
concentrated in central-oceanic regions with a low iron con-
tent, predestined for the application of the ISA method. Due
to the trash, there would be a reduction in the ISA efficiency
so we propose the integration of the plastic depletion prob-
lem into our ISA method: specific technologies can be in-
stalled both inside and outside of a container ship vessel, i.e.,
plastic trash collection, plastic trash sorting, plastic trash ex-
trusion, plastic trash burning, ISA production and emission.
The aforementioned processes are well known and need no
description here. Trash or waste burning has the advantages
of delivering an effective hot carrier gas with high buoyancy
for the uplift of ISA and for delivering HCl as a cocatalyst of
ISA. With the plastic extruder, most carrier parts of floating
supports on the reef coral, sponge and mussel habitats could
be produced.
In addition to the larger plastic fragments, the fine float-
ing plastic debris with particle diameters in the micrometer
range is another problem (van Sebille et al., 2015). Instead of
carrying out the micro-trash separation by technical means,
mussel and oyster farming may remove this ocean surface en-
vironmental problem. The floating micro-trash particles are
collected by the bivalves and excreted as pseudo-faeces pel-
lets and ultimately become part of the sediment layer at the
ocean bottom.
Within the iron cycle, the photolytically driven oxidant
production with iron participation may not be reduced to
q
Cl and q
OH in the troposphere and O2by assimilation:
when iron is cycled through the mantle at temperatures above
2500 K, Fe(III) is reduced to Fe(II) by the release of O2
(Bykova et al., 2016). This phenomenon may be driven by
the blackbody radiation containing a great fraction of pho-
tons with a wavelength shorter than 2µm at and above this
temperature level.
3.2 ISA activates the O2input to the deep ocean
Ocean ecosystems are based on certain balances between ox-
idizing and reducing agents. As a result of the ISA-triggered
additional input of organic carbon to the ISA emission region
(i.e., the ISA precipitation region), as described in Sect. 3.1,
oxygen consumption by increasing organic debris precipita-
tion could increase. The recent O2decline in some oceanic
regions may result, at least in part, from the deposition of sol-
uble iron deriving from flue gas pollution. Also discussed in
Sect. 3.1 is the decrease in the oxidation efficiency within the
water column by measures to increase the sinking velocity of
the organic-C-containing debris. The increase in the sinking
velocity of the organic-C-containing debris is an effect that
might completely compensate for the oxygen loss by the ox-
idation of the ISA-induced debris mass increase.
Recently, and without ISA influence, oxygen deficiency
appears to have developed in many parts of the ocean as de-
scribed in the introduction. Oxygen deficiency is usually due
to insufficient vertical water exchange owing to an increased
vertical density gradient rather than the result of increased
phytoplankton production.
Oxygen deficiency (hypoxia) is found frequently between
the oxic surface layer (the oxygenated one) and the oxic
deep-water layer (Bruland, 2006; Capone and Hutchins,
2013). Due to the climate warming, the localities with a lack
of oxygen seem to be intensifying and expanding already to-
day (Kalvelage et al., 2013).
The deepest water layer of most ocean basins results from
the Antarctic wintertime ocean surface ice generation by
fractionating seawater into salt-poor sea ice and salt-rich
dense brine. This results in the production of cold, high-
density oxic brines which sink to the bottom of the South-
ern Ocean. The cold high-density oxic brines spread as a thin
oxic bottom layer as far as the ocean basins north of the equa-
tor. The most recent severe climate warming, which induced
a disturbance of the THC, is likely to have been activated
by the increasing inflow of fresh meltwater from Greenland
into the North Atlantic. This inflow disturbs the downflow of
the Gulf Stream water (Rahmstorf et al., 2015). Due to the
increased melt of the glaciers of the Antarctic, the salt con-
tent of the ocean surface around Antarctica decreased. This
effect increased the ocean surface covered by sea ice (Bin-
tanja et al., 2013). This freezing of the salt-poor meltwater
layer decreases the production of dense brines. This again
decreases the downflow of brine, reducing again the vertical
components of the ocean currents.
Through the ISA-induced cooling, the oxygen and CO2
flux into the deep-ocean basins will be restored due to the in-
put of the cold dense oxygen and CO2-enriched polar surface
water: a reduced meltwater production of the Greenlandic
and Antarctic ice shields by falling surface layer tempera-
tures will restore and intensify the thermohaline circulation
within the northern polar regions by increasing the amount
of Gulf Stream water reaching those regions and by pro-
ducing the circum-Antarctic sea ice cover without meltwater
dilution, which induces the production of cold high-density
brines sinking to the ocean basin bottoms (Ohshima et al.,
www.earth-syst-dynam.net/8/1/2017/ Earth Syst. Dynam., 8, 1–54, 2017
18 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
2013; Rahmstorf, 2006). Figure 5 illustrates the ocean basin
vertical mixing circles.
3.3 Phytoplankton fertilizer extraction from ocean
sediments and underlying crust
The oceanic crust is composed of peridotites, basalts and ser-
pentine rock and has a layer of sediment on top. Sediments
and bedrock contain reductive and alkaline components ex-
tractable by seawater. The cause of the ocean water flow
through the sediment layer and base rock is the temperature-
difference-driven convection. Sediment compaction by grav-
ity, subduction-induced compaction and subduction-induced
hydroxyl mineral dehydration may be further reasons for wa-
ter movement through the sediment layer at the ocean bot-
tom.
Olivine is one of the main mineral components of oceanic
crust rock layers below the sediment layer. Hauck et
al. (2016) simulated the effects of the annual dissolution of
3 Gt olivine as a geoengineering climate cooling measure in
the open ocean, with a uniform distribution of bicarbonate,
silicic acid and iron produced by the olivine dissolution. An
additional aim of this work was the development of a neutral-
ization measure against the increasing acidification of sea-
water. All the components of olivine (SiO2, Fe(II) and Mg)
are phytoplankton fertilizers. They calculated that the iron-
induced CO2removal saturates at on average∼1.1 Pg C yr−1
for an iron input rate of 2.3 Tg Fe yr−1(1 % of the iron con-
tained in 3 Pg olivine), while CO2sequestered by alkaliniza-
tion is estimated to be ∼1.1 Pg C yr−1and the effect of sili-
cic acid represents a CO2removal of ∼0.18 Pg C yr−1. These
data represent the enormous potential of the ocean crust rock
as a source of phytoplankton fertilizer.
The flow of seawater through anoxic sediments and
bedrock results in the reduction in its SO2−
4content, as well
as the extraction of the soluble fraction from the sediment,
such as Mn(II), Fe(II), NH+
4and PO3−
4. The chemical and
physical extraction processes are enhanced by the action of
microbial attack at the border between oxic seawater and
anoxic sediment parts within this huge aqueous system.
At suboxic conditions soluble Fe(II) and Mn(II) have op-
timum solubility or may be fixed as solid Fe(II)3(PO4)2,
FeCO3, MnCO3, FeS2, S0and other Fe–S compounds
(Ohman et al., 1991; Roden and Edmonds, 1997; Slomp et
al., 2013; Swanson, 1988; Wallmann et al., 2008).
Silicon is mobilized too, from the dissolution of silicates
and SiO2in methanogenic conditions by complexation with
reduced humic acid (HA) (Vorhies and Gaines, 2009; Wall-
mann et al., 2008). In the reduced conditions, HA is charac-
terized by catechol and other polyphenolic functions, which
allows HA to complex with silicon (Belton et al., 2010; De-
madis et al., 2011; Jorgensen, 1976) and with other metal
cations.
Silicate dissolution mobilized Ca2+, Mg2+, Ba2+, Fe2+,
Na+, K+. Fe2+, Mn2+and PO3−
4precipitate more or
less as sulfides and carbonates within the sediment
(Fe(II)S2, CaCO3, MgCa(CO3)2, Fe(II)CO3, Mn(II)CO3,
Fe(II)3(PO4)2) and within its suboxic surface (BaSO4) and
at its oxic surface (SiO2, Fe(III)OOH, Mn(IV)O2, clay min-
erals). The authigenically formed ferromanganese nodules
(Kastner, 1999) are formed by in situ microbial precipitation
from sediment pore water, squeezed out to the seafloor on
the sediment layer (Nayak et al., 2011; Wu et al., 2013). The
main components of the nodules are the phytoplankton fer-
tilizer components SiO2, Fe oxides and Mn oxides (Nayak et
al., 2011).
Having left the border between anoxic and suboxic near-
surface sediment, the HA catechols are changed by reversible
oxidation into quinone or quinhydrone configurations by the
decay of the Si catechol complex. Like most of the chem-
ical reactions within the sediment compartment, the oxida-
tion of the HA–Si complex is directed by microorganisms.
The microorganisms involved use HA as an external redox
ferment (Benz et al., 1998; Bond and Lovley, 2002; Coates
et al., 1998; Kappler et al., 2004; Lovley and Blunt-Harris,
1999; Lovley et al., 1999; Piepenbrock et al., 2014; Straub et
al., 2005). After the arrival of the pore water originating from
the anoxic deeper sediment or from the bedrock at the sub-
oxic surface-near sediment layers, the oxidized HA releases
Si(OH)4and NO−
3produced by microbial NH+
4nitrification
(Daims et al., 2015; van Kessel et al., 2015). Depending on
the Si(OH)4concentration produced, this can trigger the pre-
cipitation of layered silicates such as smectites, glauconite
and celadonite as well as silica (Bjorlykke, 2010; Charpen-
tier et al., 2011; Gaudin et al., 2005; Polgári et al., 2013; Pu-
fahl and Hiatt, 2012; Zijlstra, 1995). Similar to HA, the clay
mineral formation within the sediment and the use of the re-
dox potential of these authigenic minerals are, at least in part,
the result of microbial action (Konhauser and Urrutia, 1999;
Kostka et al., 1996).
Due to its chelating properties, HA generate soluble to
neutral Fe complexes of high stability even in oxic and weak
alkaline ocean water conditions. As iron and HA have identi-
cal sources, especially chemoclines, even faeces HA can act
as shuttles between Fe sources and phytoplankton (Schmidt
et al., 2016). But within oxic ocean milieus they become de-
pleted, ultimately like every organic C substance, by oxida-
tion.
The deep-ocean currents take up the pore water percolates
out of the sediment, and considerable amounts of the dis-
solved, colloidal or suspended sediment originating elements
are THC-conveyed to the surface (Lam and Bishop, 2008),
where they activate the phytoplankton production again. This
also triggers the CO2conversion to organic C, resulting in a
cooling of the troposphere as described in Sect. 3.1. It also
repeatedly cools the troposphere by increasing the DMS for-
mation as described in Sects. 2.5 and 3.1.
Earth Syst. Dynam., 8, 1–54, 2017 www.earth-syst-dynam.net/8/1/2017/
F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 19
N60°30°-30°-60°S0°
Ocean basin
NADW
O2, CO2
O2, CO2O2, CO2O2, CO2
O2, CO2 O2, CO2
O2, CO2
O2, CO2
O
O2
CO2
O2, CO2 O2, CO2
O2, CO2
Thermocline
Halocline
Halocline
ISAprecip.
AABW
Sediment
O
C
2
O
O
O
O
AABW Antarctic bottom water
NADW North Atlantic Deep Water
Preferred ISA precipitation region
Figure 5. The motor of the Antarctic bottom water (AABW) current is the sea ice production of the Southern Ocean area bordering Antarc-
tica. The North Atlantic Deep Water (NADW) current is driven by decreasing Gulf Stream temperature on its way north. Climate warming
especially the faster temperature rise at higher latitudes, shifts the region of the Gulf Stream downflow and NADW further to the north, as a
result of the lower 1t between equatorial and polar surface water. This shift puts additional Greenlandic coast regions in contact with warm
Gulf Stream water and rising air temperatures, as another component of poor increasing amounts of fresh meltwater on the ocean surface.
The rising meltwater volume and the Gulf Stream, flowing further north, increase the contact region between Gulf Stream water and fresh
meltwater. This produces increasing amounts of original Gulf Stream water, which are, however, too low in density to sink and to become part
of NADW. Temperature rise at higher latitudes reduces the salt content of ocean surface water around Greenland and Antarctica, inducing
reduced NADW and AABW volumes. Due to the reduced downflow current volumes, the amounts of CO2and O2to the deep-ocean basin
are reduced as is the vertical fertilizer transport from the ocean basin bottom to the phytoplankton at the surface.
4 The main cooling effects induced by the iron cycle
on the ocean crust
4.1 Carbon storage as authigenic carbonate in the
ocean crust
The mechanism described in this section has the highest in-
fluence on the climate, due to its carbon storage capacity
which is greater than that of their sediment layer. The convec-
tive water flow through the huge alkaline ocean crust volume
is estimated to be about 20–540 ×103km3yr−1(Nielsen et
al., 2006). The oceanic crust comprises the largest aquifer
system of the Earth, with an estimated rock volume of
2.3 ×109km3and a fluid volume of 2 % of the total ocean, or
∼107km3(Orcutt et al., 2011). The system of the mid-ocean
rifts (MOR) and subduction zones and the sector between
these volcanically active regions are part of the Earth mantle
convection cycle and part of interconnected aquifer system
mentioned above. The bottom water of the ocean basins is in
close contact with this conveyor-belt-like moving rock layer
of the oceanic crust. New oceanic crust is produced at the
MOR: during its cooling, it is pulled apart from the MOR
by the moving underlying mantle, and ultimately the moving
mantle draws the crust down into the deeper mantle below
the subduction zones. The oceanic crust has a sediment layer
on top of its assemblage of multi-fractured crystalline and
volcanic rocks. Both sediment and igneous bedrock interior
are in an anoxic reduced and alkaline state; the temperature
on top of the sediment surface at the ocean bottom is round
about 0 ◦C, but the temperature increases up to >1000 ◦C
within the igneous bedrock basement. As there is no effective
sealing between cold bottom water and the high-temperature
zone, the water content of sediments and the fractured base-
ment flows through the crust in multiple thermal convection
cycles positioned between the cold surface and the hot deep.
Alkalinity and alkalinity-inducing compounds of the
ocean crust rock layers extract CO2and HCO−
3from sea-
www.earth-syst-dynam.net/8/1/2017/ Earth Syst. Dynam., 8, 1–54, 2017
20 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
water by carbonate precipitation in the fissures during sea-
water percolation through the multi-fractured rock (Coggon
et al., 2012). A carbon uptake of 22 to 29 Mt C yr−1is esti-
mated during the hydrothermal alteration of the oceanic crust
(Kelemen and Manning, 2015). This is more than the car-
bon uptake by the overlying sediment layer of the oceanic
crust, which is estimated to be 13 to 23 Mt C yr−1(Kelemen
and Manning, 2015). The oceanic crust is composed of peri-
dotites, basalts and serpentine rock, with a sediment layer on
top. Said rock layers contain reductive and alkaline compo-
nents. Seawater circling through these rock layers loses its
oxygen, sulfate and nitrate and even part of its hydrogen car-
bonate content by reduction and precipitation and becomes
enriched with methane and other reductants (Evans, 2008;
Janecky and Seyfried, 1986; Kelemen et al., 2011; Müntener,
2010; Oelkers et al., 2008; Sanna et al., 2014; Schrenk et al.,
2013; Sissmann et al., 2014).
Figure 6a and b illustrate, respectively, the differences be-
tween a poorly and a sufficiently mixed ocean.
Due to the opposing chemical milieu differences between
the oxic ocean water inflow and anoxic reduced and alkaline
sediment and basement, the ocean water convection cycles
through the ocean crust act as continuous chemical reaction
systems and form habitats of intense microbial action (Ivars-
son et al., 2016). The greatest chemical reaction intensity is
found at MOR, subduction zones and volcanic sea mounts;
between MOR and subduction within the abyssal plain, con-
vection cycling occurs (Orcutt et al., 2011). Because the hy-
drogen carbonate load of the ocean water inflow reaches pre-
cipitation as carbonates of Ca, Mg, Fe and Mn within the
alkaline rock interior and by chemical reduction of sulfate,
nitrate and hydrogen carbonate, the ocean basements act as
huge CO2-carbon storages. There is no doubt that the ocean
crust carbonate depot is the most effective carbon storage,
more effective than any other organic carbon storage.
Within the huge ocean crust contact volume, seawater
changes the alkaline pyroxenes and basalts into serpentine,
diabase and carbonates by producing heat, hydrogen and
rock volume expansion and by a permanent production of
numerous fissures. The ocean water sulfates react with the
silicate components to magnetite, pyrite and barite. The sea-
water hydrogen carbonate load precipitates within the rock
fissures as magnesite, calcite, siderite and dolomite. By heat
transfer from hot rock and chemical reaction, heat circling
through the primary and newly generated multiple fissures in
the former mantle rock, the seawater inflow heats up, pro-
ducing convective flow. At fissures where the alkalized flow
of convection water containing hot CH4and H2comes out
with pH 9 to 11 and comes into contact with fresh seawater,
carbonate precipitates and builds up carbonate chimneys as
high as skyscrapers (Kelley et al., 2005).
The convective seawater flowing only through the MOR
system is estimated to be about 20 to 540 ×103km3yr−1
(Nielsen et al., 2006). This volume is more than the global
river flow of about 50km3yr−1(Rast et al., 2001).
The weathering reaction conditions and the seawater al-
kalization during the intense seawater contact with the alka-
line MOR rocks are much more aggressive, and thus more
effective, compared to reaction conditions and alkalization
during the precipitation water contact and during weathering
reactions of continental rocks. This is confirmed by the alka-
line pH of up to 11 of the “white smoker” MOR outflow in
spite of its haline salt-buffered seawater origin (Kelley et al.,
2005). Even the most alkaline runoff from limestone karst
spring freshwaters or within karst cave freshwaters does not
exceed pH levels of 8.5 (Li et al., 2010; Raeisi and Karami,
1997; Righi-Cavallaro et al., 2010). Because of the enormous
carbonate absorption capacity of the oceanic crust, it has
been proposed to use it as a storage of CO2(Kelemen and
Matter, 2008). As the igneous crust rock aquifer generates
H2during its contact with ocean water parts of the carbonate
precipitation, carbonate is reduced in part to organic and/or
graphitic C, depending on the reaction temperatures by biotic
or abiotic reduction (Galvez et al., 2013; Holm et al., 2015;
Malvoisin et al., 2012; Rumble, 2014; X. Wang et al., 2014).
There is no doubt that the efficiency of the pH-dependent
CO2absorption, carbonic acid neutralizing at the ocean sur-
faces, and the hydrogen carbonate precipitation to carbonate
processes at and within the oceanic crust are dependent on
the activity of the THC within the ocean basins. During cold
climate epochs, with unstratified water columns and undis-
turbed THC, the CO2conversion to ocean crust carbonate is
activated, as is the CO2conversion to the organic fraction of
ocean sediments. Just the opposite has been found to be true
for the burial of organic C in ocean basin bottom sediments:
according to Lopes et al. (2015) the overwhelming organic
debris fraction produced during the main glacial episodes
from the phytoplankton habitat at the surface is oxidized and
remineralized in the well-mixed ocean basin (Lopes et al.,
2015). As the CO2level in the atmosphere is at its lowest
during the main glacials, the remaining C sinks of the oceans
seem to be of much greater efficiency than the iron-induced
production of organic C by assimilation: the most promi-
nent C sink is the authigenic carbonate C burial in the alka-
line ocean crust. There seems to be no doubt that the vertical
well-mixed ocean during the main glacials works as an effi-
cient pump to transport dissolved CO2and O2to the ocean
basin bottoms: there, O2acts as a mineralizer of organic C,
and CO2C is buried as authigenic carbonate C in the oceanic
crust.
Table 2 gives an overview of some trends in C burial
depending on the climate condition change between main
glacial and interglacial.
Lopes et al. (2015) found just the opposite in ocean sedi-
ment layers produced during the warm interstadial compared
to the cold main glacial, i.e., a high burial rate of organic C in
the ocean bottom sediment. But in spite of the high organic C
burial rate, the interstadial CO2levels were kept higher than
those of the main glacial. Even in this regard, the Lopes et
al. (2015) results fit in well to our CO2sink model. Dur-
Earth Syst. Dynam., 8, 1–54, 2017 www.earth-syst-dynam.net/8/1/2017/
F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 21
Alkaline, fertilizer & Fe input
by weathering
Acidity & CO2 input by volcanic gases
& anthropogenic sources
Organic & carbonate-C precipitate
& alkalinity input by phytoplankton
O2,
CO2,
sulfate,
organic C,
salt,
Ca & MgCO3,
water,
absorbed
from crust
& sediment
Alkalinity,
fertilizers,
Fe,
reductants,
extracted
from crust
& sediment
Thermic-convective sediment
& crust extraction
Ocean
Thermo-haline Ocean
Convection
Ocean crust
Acidity & CO2 extraction
by partial melt of
the subduced crust Mantle
Mantle convection
Cold dusty glacial
atmosphere
Subduction
volcanism
Continent
A
c
idit
y &
CO
2
i
n
put
by
v
&
ant
h
ropo
g
enic
s
High Fe input by ISA
Alkaline, fertilizer & Fe input
by MOR, rivers, sediments
and icebergs
Acidity & CO2 input by volcanic gases
& anthropogenic sources
Phytoplankton = organic C +
carbonate C + alkalinity
O2,
CO2,
sulfate,
organic C,
salt,
Ca & MgCO3,
water,
absorbed
from crust
& sediment
Alkalinity,
fertilizers,
Fe,
reductants,
extracted
from crust
& sediment
Thermic-convective sediment
& crust extraction
Unstratied and
wellmixed oxic ocean
Thermohaline ocean
convection
Ocean crust
Acidity & CO2 extraction
by partial melt of
the subduced crust Mantle
Mantle convection
Subduction
volcanism
Continent
(a)
Fe input by eolic dust & ISA
Alkaline, fertilizer & Fe input
by weatering
Acidity & CO2 input by volcanic gases
& anthropogenic sources
Organic & carbonate-C precipitate
& alkalinity input by phytoplankton
O2,
CO2,
sulfate,
organic C,
salt,
Ca & MgCO3,
water,
absorbed
from crust
& sediment
Alkalinity,
fertilizers,
Fe,
reductants,
extracted
from crust
& sediment
Thermic-convective sediment
& crust extraction
Ocean
Thermo-haline Ocean
Convection
Ocean crust
Acidity & CO2 extraction
by partial melt of
the subduced crust Mantle
Mantle convection
Clean hot
greenhouse
atmosphere
Subduction
volcanism
Continent
A
c
idit
y &
CO
2
i
n
put
by
v
&
ant
h
ropo
g
enic
s
Low Fe input by ISA
Alkaline, fertilizer & Fe input
by rivers, meltwater, shelf
sediments and icebergs
Acidity & CO2 input by volcanic gases
& anthropogenic sources
Phytoplankton = organic C +
carbonate C + alkalinity
Organic C,
carbonate C,
sulde
salt,
water,
absorbed
from crust
& sediment
Alkalinity,
fertilizers,
Fe,
reductants,
extracted
from crust
& sediment
Thermic-convective sediment
& crust extraction
Oxic & acid
ocean surface
Ocean crust
Acidity & CO2 extraction
by partial melt of
the subduced crust Mantle
(b)
Mantle convection
Subduction
volcanism
Continent
Anoxic & alkaline
ocean deep
Chemocline/halocline
Figure 6. This figure presents the essential differences between an unstratified well-mixed ocean basin under a cold and dusty atmosphere
during the cold main glacial, with low atmospheric GHG concentrations (a) and a stratified ocean basin with a meltwater layer on top of a
saline ocean water layer during a warm interglacial, with a hot and dust-free greenhouse atmosphere (b). Panel (a): due to the unstratified
well-mixed water column in Basin 6A, CO2and O2absorbed at the water surface are distributed within all parts of the basin. There are high
production rates of organic carbon produced by phytoplankton in the top layer, and this carbon is oxidized on its way down to the sediment
layer, with only a minor generation of organic sediment. Carbonate carbon produced by the phytoplankton becomes dissolved to a great extent
within the deeper basin parts generating HCO−
3, CO2and HCO−
3. By cycling the basin bottom water through the alkaline bottom sediment
and ocean crust aquifer, CO2and HCO−
3become precipitated and buried as carbonate C. The recycled bottom water becomes enriched by
Fe fixed to organic chelators and is transported back to the surface. Due to the unrestricted downflow and transfer of CO2from the former
surface water into sediments and into underlying base rock as carbonate carbon, the buried carbonate C exceeds the buried organic C amount.
Panel (b): an interglacial episode with high GHG levels accompanied by elevated surface temperatures generates increased meltwater and
surface water runoff. Because the saline-poor water layer spreads to the saline ocean water and induces at least a regional stratification of the
ocean basins water column, the production of brine-induced surface water downflow is stopped, as meltwater freezing generates neither brine
nor any vertical surface water movement. This stops any down transport of absorbed CO2and O2too and generates anoxic conditions within
the underlying saline layer. The anoxic saline layer becomes anoxic and alkaline by sulfate and nitrate reduction. Any phytoplankton-induced
organic and carbonate litter trickles down through the anoxic and alkaline layer: CaCO3without dilution in the alkaline water and organic C
without oxidation in the anoxic milieu. Carbonates of Ca and Mg may precipitate in small amounts at the chemocline between light acidic
CO2saturated water and the alkaline saline layer. The carbonate precipitate then mixes with the down-falling phytoplankton-originating
litter.
www.earth-syst-dynam.net/8/1/2017/ Earth Syst. Dynam., 8, 1–54, 2017
22 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
Table 2. Interglacial climate episodes were hot, nearly dust-free and had elevated levels of GHGs. The interglacials coincided with stratified
water columns. The stratified ocean has much reduced activity due to the reduced CO2transport to the bottom of the ocean basin. As the
O2transport is reduced and the lower part of the basin is anoxic, the oxidative mineralization of the organic litter fall from phytoplankton
activity at the surface is reduced and generates sediments rich in organic substances. As sulfate, nitrate and in part CO2within the anoxic
water column are reduced to sulfide, ammonium and CH4, the pH increases to alkaline. This can induce carbonate precipitation near the
chemocline. During the glacial maxima with cold temperatures, dustiness and low greenhouse gas levels, the ocean basins had well and
vertically mixed water columns with highest carbonate C burial and lowest organic C burial.
Effect on Sediment +Sediment +
crust below crust below
well and stratified and
vertically mixed anoxic water
water column column
Mass ratio of buried sediment C : 1<1 to 1 or >1
sediment and crust oceanic crust C
carbon
Mass ratio of buried organic C : 1 up to 1 or >1
sediment and crust carbonate C
carbon
Authigenic carbonate No Yes
produced within the
water column
Tropospheric Dust High Low
parameters CO2Low High
CH4Low High
Temperature Cold Warm
ing the glacial climate warming events, enormous meltwa-
ter volumes were generated and induced stratification effects
in ocean basins by placing a meltwater blanket on the saline
ocean water surface (Praetorius et al., 2015). The transport of
CO2and O2into the basin bottoms became interrupted. The
drizzle of phytoplankton litter remained un-oxidized, and
as a further consequence the amount of carbonate C burial
within the ocean crust decreased.
The continuous availability of chemical activity, as a
chemical reaction vessel and as an alkalinity reservoir of
the oceanic crust, is maintained by the continuous genera-
tion of new crustal rock material of 21 km3yr−1(Orcutt et
al., 2011). This huge rock volume production capacity has
enough alkalinity and fertilizer reserves to maintain the ab-
sorption, neutralization and precipitation of a multiple of the
recent incoming CO2and HCO−
3.
THC is the main transport medium of carbon from the at-
mosphere into the deep on Earth. This makes THC the most
prominent climate stabilization element.
The realization of the significance of THC as a stabiliza-
tion element of our recent-climate model raises questions
about the stability of the THC. As stated in Sect. 1, the main
factor for destabilizing the THC seems to be the stratifica-
tion of the water column by the desalting of surface ocean
layers by freshwater dilution from increasing ice melting
(Hansen et al., 2016). The low-density meltwater generates
a layer on the ocean water, producing a stratified water col-
umn. The stratification hampers or prevents the transport of
CO2and O2-containing surface water into the deep-ocean
basin parts. The most severe consequence of such stratifica-
tion, to oceanic ecosystems, is the development of anoxic mi-
lieus within the stratified ocean basins.
Typical characteristics of episodes with stratified water
columns in ocean basins are the black shales and black lime-
stones representing sapropel remnants. The repeated devel-
opment of stratified ocean basins during the Phanerozoic
epoch occurred as a consequence of elevated CO2levels in
the atmosphere. This caused high sea surface temperatures
(Meyers, 2014) and as a global consequence a global in-
crease in evaporation, precipitation and production of brines
of higher concentrations.
Hansen et al. (2016) also pointed out that the increasing
meltwater runoff from polar and subpolar ice layers can in-
duce the cover of denser ocean water by a meltwater layer.
But the generation of increasing precipitation and surface
water runoff accompanied by increasing brine production
during hot high-CO2climate episodes had exactly the same
consequences in the past geological epochs, as we learn from
Meyers (2014).
We now have to fear this combination of the CO2-
dependent temperature-rise-generated precipitation increase
and the meltwater increase from glacier melt. Mankind now
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 23
has to find the appropriate tool to rise to this challenge or will
otherwise fail to meet it.
A melt increase might drive the destabilization of THC.
And at first the top layers of the ocean basins will suffer from
acidification, and the deep layers will become alkaline and
anoxic.
By starting the ISA process, the induced climate cool-
ing will decrease the Greenland glacier melt. The mini-
mized freshwater inflow to the North Atlantic Ocean re-
duces the dilution of the salty Gulf Stream and increases
the downflow quantity of oxic and CO2-containing salty sur-
face water. In parallel, the surface increase in sea ice pro-
duced on the Southern Ocean surrounding the Antarctic con-
tinent is followed by increased downflow of oxic and CO2-
containing cold brine to the bottoms of the oceanic basins.
Both effects increase the THC activation: the flow of alka-
line, phytoplankton-fertilizer-enriched and oxygen-depleted
deep-ocean water to the surface. This activates CO2absorp-
tion from the atmosphere by phytoplankton growth and by
CO2absorption
One of the proposed alternative climate engineering mea-
sures aims to absorb atmospheric CO2by reducing the sur-
face ocean acidity and by producing phytoplankton fertiliz-
ers. To transfer 1.1 ×109t yr−1CO2carbon into the ocean, it
would be necessary to crush 3 ×109t yr−1of the ocean crust
and mantle rock mineral olivine to a particle diameter of 1 µm
and suspend it at the ocean surface (Hauck et al., 2016; Köh-
ler et al., 2010, 2013). These numbers seem to be 2 orders
of magnitude too high. Keleman and Manning (2015) cal-
culate a carbon mass subduction of about 50 ×106t C yr−1
(C in oceanic crust, bedrock and sediment layer). Regardless
of which of the two calculations is wrong, the technology to
implement the Hauck et al. (2016) proposal is far from being
an economic reality.
The proposed reaction of CO2with olivine is carried out
with much better effectiveness by nature, without any costs
and within the ocean crust in sufficient quantity. To minimize
CO2emission it has been proposed to minimize power sta-
tion flue gas CO2by absorption by lime suspension (Haas et
al., 2014). This measure seems to be unnecessary when the
ISA method comes into practice.
The fertilizing elements which phytoplankton needs, such
as Si, P and Fe, are all present in the ocean crust (Lyubet-
skaya and Korenaga, 2007) and a property of the ocean crust
water extract. The intensification of the THC would also in-
crease the fertilizer concentration at the ocean surface in the
phytoplankton layer. As demonstrated, the undisturbed THC
is essential for keeping the climate stabilized (Coogan and
Gillis, 2013).
The ocean crust from the warm Mesozoic epoch, which
had no frozen polar regions, contained about 5 times
more authigenic carbonate than ocean crust younger than
60 million years (Coogan and Gillis, 2013). Coogan and
Gillis (2013) interpreted this as possible consequences of
higher bottom water temperature and/or different seawater
composition. Insua et al. (2014) found evidence that the
salinity of the ocean bottom water during the Last Glacial
Maximum had been up to 4 % greater than today. It seems
evident that the cause of the latter had been the higher vol-
ume of brine produced during sea ice freezing. This fact
demonstrates that disturbed or weakened THCs might be
the cause of the reduced carbonate C uptake of the ocean
crust. The quantity of carbonate precipitation depends on the
CO2and/or HCO−
3input with seawater. As a consequence,
the quantity of the ocean crust CO3uptake varies depend-
ing on the activities of the THCs or stratified ocean basins:
strong THCs increase the crust carbon content; weak THCs
decrease it.
Independently of the cause of stratification events (by
brine generation, by freezing or by evaporation), the ocean
basins possess a removal mechanism which extracts salt from
the brine and change the brine to seawater of a normal salt
concentration. This mechanism has kept the salt concentra-
tion of seawater fairly constant during the past geological
epochs. This effect, which achieves a constant salinity level,
depletes any brine-induced stratification and restores well-
mixed ocean basins again.
According to Hovland et al. (2006a, b, 2014), this de-
salination takes place by continuous salt removal from the
brine or seawater within the hot ocean crust. This desalina-
tion works independently of the salt concentration of brine or
seawater. The salt removal process acts within the ocean crust
aquifer at near-critical to supercritical seawater temperature
and pressure conditions. During the subduction of the salty
crust, rock chloride and carbonate change their cations with
silicate and are dissolved as HCl and CO2. Accompanied by
H2O, these gases are recycled to the atmosphere, mainly by
subduction volcanism, but this is done to a much lesser extent
by MOR and similar alkaline volcanism.
During the time lag between the onset of the ISA method
cooling and the appearance of the alkalinity and fertilizer
increase at the ocean surface, the cooling effect of ISA re-
mains reduced. But after this time lag, the ISA method in-
creases to optimal efficiency. Even from an economic view-
point, it seems better to compensate for this by increasing
the ISA emission at the beginning during the time lag than
doing the proposed suspending of olivine dust at the ocean’s
surface. Even lime-shell-bearing phytoplankton is able to ac-
cept small pH changes in CO2-induced dependent acidifica-
tion because it uses the buildup of calcium carbonate shells
as a detoxification measure to get rid of calcium ions from
their bodies (Müller et al., 2015). As a consequence of this
effect, only the relation between Ca carbonate sequestration
and organic carbon sequestration may decrease during the
time lag.
To sum up, through the huge aquifers of the alkaline and
reducing ocean crust, any transport of former surface wa-
ter enriched by CO2or HCO−
3induces carbonate C burial
within the aquifer interior. This is the situation within well-
mixed ocean basins without stratification. Any stratification
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24 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
decreases carbonate burial or even stops it. Stratification
changes the redox milieu below the stratification-induced
chemocline. The MOR and sediment-induced exhalation of
Fe and other metals by black smokers into the sulfidic strat-
ified ocean basin are prevented from contact with the plank-
tonic surface water habitat. But surface water runoff, as well
as meltwater inflow and iceberg melt during warm glacial cli-
mate intervals, may compensate for the lack of Fe from the
MOR and bottom sediment sources, as well as from the de-
creasing dust fall during the warm climate intervals (Hansen
et al., 2016; van Helmond et al., 2015).
4.2 Carbon storage as organic and inorganic marine
debris and as authigenic carbonate in the ocean
sediment
The uptake of authigenic hydrogen carbonate from the ocean
and its precipitation to the sediment also seem to play a major
role in the carbon cycle (Schrag et al., 2013). According to
Kelemen and Manning (2015), the carbon uptake by the sed-
iment layer of the oceanic crust can be estimated to be 13 to
23 Mt C yr−1. The carbon inventory consists of live and dead
organic carbon, carbonate carbon, and authigenic carbonate
produced by excess alkalinity deriving mainly from sulfate
reduction and silicate solution by reduced humic acids. Ac-
cording to Sun and Turchyn (2014), the formation of calcium
carbonate and its burial in marine sediments accounts for
about 80 % of the total carbon removed from the Earth sur-
face (Sun and Turchyn, 2014). Moreover, it seems possible to
distinguish between marine-formed sediment carbonate and
authigenic carbonate (Zhao et al., 2016).
As evidenced in Sect. 4.1, stratified ocean basins can differ
widely in the quantity and quality of the buried C depending
on the prevailing climate conditions and their direct and indi-
rect influences on ocean basin conditions. Table 2 lists some
of the most prominent results.
The cooling of the troposphere by ISA action stops melt-
water inflow, destructs the stratification and starts the vertical
mixture. During the previous stratification event, the alka-
lized deep-water layer had enormous CO2absorption capac-
ity. The alkalized anoxic sediment behaves in a similar man-
ner. This leads to much increased CO2absorption activity at
the beginning of the movement.
Accordingly, excess alkalinity is produced by the disso-
lution of silicates such as illite, kaolinite and feldspars, vol-
canic ash, pyroxene, or other silicate components of ocean
sediments and even opal by Si complexation with reduced
HA in methanogenic conditions (Meister et al., 2011; Ro-
den and Edmonds, 1997; Solomon et al., 2014; Wallmann et
al., 2008). Compensation by hydrogen carbonate induces au-
thigenic precipitation of microbial dolomite (Roberts et al.,
2004), Ca or Fe carbonate (Dewangan et al., 2013; Merinero
et al., 2008; Solomon et al., 2014; Sun and Turchyn, 2014;
Vorhies and Gaines, 2009; Wallmann et al., 2008), and other
minerals (Tribovillard et al., 2013).
As mentioned in Sect. 4.1, the biological processes of
chemical sediment reduction induced by the ISA fertilization
change NO−
3, SO2−
4, Fe(III), Mn(III/IV) and HCO−
3to their
deoxygenated and reduced species, including CH4, and NH+
4
generation produces a pH increase and additional alkalinity.
Furthermore, a pH drop is induced by H2evolution from
FeS2generation from FeS and H2S (Drobner et al., 1990;
Rickard and Luther, 1997), accompanied by CO2reduction
to CH4(Conrad, 1999) as well as N2reduction to NH3(Dörr
et al., 2003). The alkalinity excess converts dissolved HCO−
3
into solid lime and dolomite (Berner et al., 1970; Krause et
al., 2012; Luff and Wallmann, 2003; Raiswell and Fisher,
2004). The solid carbonates and CH4hydrate stabilize the
sediment. Outside the polar permafrost region, methane hy-
drates are stable below 300 m b.s.l. (below sea level) and at
ocean temperatures of nearly 0 ◦C (Maekawa et al., 1995).
The carbonate precipitation sequesters additional amounts
of CO2, prevents the ocean water from acidifying and ulti-
mately improves the CO2absorption by ocean water from
the atmosphere. This again cools the troposphere.
The enhanced dissolution of silicates from the ISA in-
duced by methanogenic sedimentation additionally compen-
sates for the enhanced alkalinity loss at the ocean surface,
attributed to the calcification due to foraminifera and coccol-
ithophores phytoplankton growth by ISA fertilization.
To sum up, within a well-mixed and unstratified ocean
basin the surface layer absorbs CO2and O2and becomes
well mixed into the unstratified ocean basin by the thermo-
haline basin convection. Consequences of the good mixing
are a nearly quantitative oxidation of the food chain debris to
CO2produced by phytoplankton. Most C is buried as carbon-
ate in the ocean crust and its overlying sediment. The ratio of
organic C burial to carbonate C burial is much smaller than 1.
The results of Lopes et al. (2015) from northeast Pacific sed-
iments demonstrate that, although primary productivity was
estimated to be highest during the Last Glacial Maximum, or-
ganic C burial was lowest. This coincides with our proposed
optimum mixed O2-rich milieu throughout the whole water
column.
During situations with stratified water columns in the
ocean basins or parts of them THC convection is disturbed
or does not exist at all. A surface water layer enriched with
CO2and O2absorbed from the atmosphere cannot pene-
trate through the stratified water column to the bottom of
the basin. This induces sulfate-reducing conditions below the
surface layer. Only small amounts of surface layer CO2are
changed into carbonate C at the chemocline, with the alka-
line sulfidic and anoxic parts below the chemocline. Below
the chemocline, the water column is anoxic; the organic de-
bris sediment shows minor oxidation. It is likely that the ra-
tio of organic C burial to carbonate C burial increases greatly
during stratified conditions. Concerning the huge fraction of
organic C buried during the warm glacial intervals, according
to the results of Lopes et al. (2015) from the northeast Pa-
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 25
cific, sediments demonstrate stratification events within the
research area of Lopes et al.
Stratification events may develop by warming the up-
per water layer, as well as by evaporation and precipitation
(Friedrich et al., 2008; Hansen et al., 2016; van Helmond et
al., 2015).
4.3 Minimizing CH4emissions from sediments and
igneous bedrock
The reaction product of oceanic crust minerals containing
Fe(II) such as olivine and pyrrhotite with seawater is hydro-
gen (Bayrakci et al., 2016; Mayhew et al., 2013; Neubeck
et al., 2014). The hydrogen production rate along the MOR
alone is estimated to be ∼1012 mol H2yr−1(Worman et al.,
2016). Hydrogen is fermented by microbes with hydrogen
carbonate into methane. Methane is known as a constituent
of the springs emitted by the ocean crust rocks (Früh-Green
et al., 2004).
Such and other CH4emissions, such as anoxic sediments
outside the CH4hydrate stable pressure and temperature re-
gion, induce deoxygenation within the overlying water layer
by CH4emission (Römer et al., 2014; Yamamoto et al.,
2014). CH4emissions are induced for instance by hydrother-
mal springs (Suess et al., 1999), sediment movement (Kras-
tel et al., 2014; Paull et al., 2007), seawater warming in-
duced by climate change (Serov et al., 2015; Shakhova et
al., 2005), changing ocean circulation (Berndt et al., 2014)
and ocean sediment subduction (Elvert et al., 2000; Fischer
et al., 2013). At lower vertical sediment to ocean surface
distances, the CH4emissions reach the troposphere. As the
Arctic Ocean suffers most from the climate-change-induced
warming, the CH4release within this region rises extraor-
dinarily (Phrampus et al., 2014). The most elevated global
surface-near oceanic CH4concentrations are located within
the Arctic Ocean and the Arctic troposphere (Shakhova et al.,
2008). This might be one of the reasons for the temperature
rise in the Arctic region being higher than the average surface
Earth warming.
Within the sediment and within the suboxic ocean water
column, CH4is oxidized by sulfate. Iron is an accelerator of
this microbial fermentation reaction (Sivan et al., 2014). The
ocean water column and the underlying sediment having had
contact with ISA-originating iron are elevated in their iron
content. This has different cooling effects on the troposphere:
–First, the elevated iron content in the uppermost suboxic
sediment reduces the CH4content emitted by the sedi-
ment by the anaerobic oxidation of methane by sulfate-
reducing bacteria. Below regions with ISA precipita-
tion, not only the sediment but also the whole water
column of the ocean basin is enriched by iron. Any CH4
molecule, regardless of whether it is found in the sedi-
ment or just above in the water phase or whether it is
excreted into the water column as bubbles, is oxidized
before it arrives at the water column top. With the help
of Fe-containing enzymes, methane oxidation by sulfate
is possible. This prevents the water layers above the sul-
fate oxidation zone from oxygen loss. Sulfate oxidizers
of CH4are archaea and bacteria (Basen et al., 2011).
As these microbes use Fe-containing enzymes for their
anaerobic methane oxidation processes, they work bet-
ter in iron-rich than in iron-poor environments (Sivan et
al., 2011, 2016). The iron-containing debris fall of ISA-
fed dead phytoplankton and phytoplankton-dependent
food chain links feeds the methane-depleting sulfate re-
ducer community within or near the sediment surface.
–Second, the iron content reduces the CH4bubble-
development within the sediment layer, preventing
catastrophic CH4eruptions by sediment destabilization,
CH4bursts and sediment avalanches.
–Thirdly, an elevated iron content prevents the ocean wa-
ter column from CH4-induced oxygen deficiency by the
formation of ammonium. This oxygen deficiency pre-
vention protects against the generation of the extremely
stable and very effective GHG N2O (Naqvi et al., 2010).
The oxygen-dependent life will develop problems due to
the decreasing oxygen content within a decreased vertically
mixed ocean basin induced by climate warming. An addi-
tional input of CH4would increase the oxygen-deficit death
zones. Any CH4injection into regional oxygen-deficit zones
will immediately increase their volume. Climate models pre-
dict declines in oceanic dissolved oxygen with global warm-
ing. The climate-warming-dependent decline in the oxygen
content in many ocean regions has meanwhile become man-
ifest (Stramma et al., 2010). Braking or reversing this trend
by reducing the oxygen-depleting CH4emissions should at
least help to protect regions within the ocean basins from
methane-induced oxygen deficit.
The glacial age proved that in spite of the multiplicity of
the cooling processes induced, they caused little disturbance
to the ecosystems. This predestines ISA as a steering tool to
prevent climate fluctuations such as the recent climate warm-
ing mankind is suffering from. The present study aims to de-
scribe in Sect. 5 the technical means to realize this climate
engineering project by the ISA method.
This result is contradictious to the calculations of Duprat et
al. (2016). They found an increased phytoplankton concen-
tration within the iron-containing meltwater trail of the gi-
ant Antarctica icebergs. Duprat et al. (2016) assume that the
iceberg-induced carbon export increases by a factor of 5 to 10
within its locality of influence, and they expect an increase
in carbon export by the predicted increase in the iceberg pro-
duction (for instance, Joughin et al., 2014). We interpret the
ongoing increase in icebergs and ice melt as a further severe
warning sign that the ongoing destabilization might end soon
in an insufficiently mixed ocean.
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26 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
The only artificially realizable restoration tool to change
an insufficiently or poorly mixed ocean into a well-mixed
ocean is definitely climate cooling. The ISA method appears
to be the climate cooling method of choice because it accel-
erates the conversion of atmospheric carbons into solid and
even liquid carbons with the means of nature. Compared to
the artificial aerosol systems based on TiO2or H2SO4(Pope
et al., 2012), sea-salt aerosol has advantages, such as better
controllability and economy.
5 Iron effects on shore
5.1 Importance of iron for terrestrial landscapes
As seen in previous sections, the atmospheric deposition of
iron together with other macronutrients and micronutrients
set important controls on marine ecology and biogeochem-
istry: for terrestrial ecology and biogeochemistry the impor-
tance of iron is similar.
Iron is one of 17 essential elements for plant growth
and reproduction (Pérez-Sanz and Lucena, 1995). Iron is
an essential micronutrient (or trace element) only required
by plants in small amounts for bio-functions such as the
production of chlorophyll and photosynthesis (Hochmuth,
2011). Iron is involved in many other important physiolog-
ical processes such as nitrogen fixation and nitrate reduction
and is required for certain enzyme functions (Morrissey and
Guerinot, 2009).
Iron is the fourth most abundant element of the earth’s
crust (4.2 %) and thus iron is seldom deficient, as despite its
high abundance in soil, iron solubility is extremely low and
its availability depends on the whole soil system and chem-
istry. Chlorosis (yellowing) is associated with iron deficiency
in plants on land (Anderson, 1982; Mengel and Geurtzen,
1986), but the chemistry of iron in soils and its availability
to plants (Lindsay and Schwab, 1982) is outside the scope
of this review; thus, only a brief overview is given. How-
ever, while small amounts are necessary for growth, iron can
become toxic to plants. Iron toxicity is associated with large
concentrations of Fe2+in the soil solution (Becker and Asch,
2005) and leads to oxidative stress. As a consequence, iron-
uptake systems are carefully regulated to ensure that iron
homeostasis is maintained. Iron availability represents a sig-
nificant constraint on plant growth, and plants have devel-
oped distinct strategies to ensure Fe solubilization and uptake
(Forieri and Hell, 2014). In forests, microorganisms such as
fungi and bacteria play a role in nutrient cycling (Philpott,
2006). A particularly efficient iron acquisition system in-
volves the solubilization of iron by siderophores (Kraemer,
2004), which are biogenic chelators with a high affinity and
specificity for iron complexation.
Iron-deficiency-induced chlorosis represents the main nu-
tritional disorder in orchards and in crops grown on calcare-
ous and/or alkaline soils (Abadía et al., 2011) in many ar-
eas of the world. Iron deficiency is a worldwide problem as
calcareous soils cover over 30% of the earth’s land surface
(Basar et al., 2014), especially in arid and semiarid regions,
and it has a large economical impact because crop quality
and yield can be severely compromised (El-Jendoubi et al.,
2014; Rombolà and Tagliavini, 2006); thus, several methods
of correction have been developed. Iron canopy fertilization
(foliar fertilization) can be a cheaper, more environmentally-
friendly alternative to soil treatments with synthetic Fe(III)
chelates for the control of Fe chlorosis in fruit trees (Fer-
nández et al., 2013). But iron chelates are expensive and
have to be applied annually. Several sprays aiming to acti-
vate the Fe pools in a chlorotic leaf by foliar iron fertiliza-
tion have been tested and were generally as effective as sim-
ple spay fertilization with iron sulfate (Abadía et al., 2011),
and both are effective in regreening treated leaf areas, both in
peach trees and sugar beet plants (El-Jendoubi et al., 2014).
Iron deficiency chlorosis in soybean was solved by foliar
sprays, which significantly increased the yield of three cul-
tivars tested, and the yield responses obtained were about
300 kg ha−1(Goos and Johnson, 2000).
Although foliar Fe fertilization seems to be potentially
effective, the scientific background for this practice is still
scarce, and we did not find evidence that soluble iron con-
tained in atmospheric dust aerosols has already been proved
to be able to play this role.
The fertilizing role of African dust in the Amazon rain for-
est is well known (Yu et al., 2015b) but attributed to the P
input. On a basis of the 7-year average of transatlantic dust
transportation, Yu et al. (2015a) calculated that 182 Tg yr−1
dust leaves the coast of North Africa (15◦W), of which
43 Tg yr−1reaches America (75◦W). The dust reaching the
Caribbean and the Amazon comes mainly from northwestern
Africa (Algeria, Mali and Mauritania) (Gläser et al., 2015).
The average dust deposition into the Amazon Basin over
7 years is estimated to be 29 kg ha−1yr−1(Yu et al., 2015b),
providing about to 23 g ha−1yr−1of phosphorus to fertil-
ize the Amazon rain forest, together with Mg and Fe. Al-
though not directly related to ISA, this dust deposition allows
biomass fertilization and thus CO2removal from the atmo-
sphere.
The widespread tropical soils, mostly laterites, are defi-
cient in phosphate and nitrogen but not in autochthonous
iron. The only exception to this applies to all the epi-
phyte plants and the plants growing in soil-free localities
without any autochthonous iron. These plants might profit
from the ISA method. Such plant communities are localized
for instance on top of the well-known tepuis (table moun-
tains north of the Amazon basin near the borders of Brazil,
Venezuela and Guyana) and on the tree branches in the rain
forests without roots in the ground. From Köhler et al. (2007)
the epiphytes flora on the tree branches of the rain forests
may contain up to 16 t ha−1(Costa Rica) up to 44 t ha−1
(Colombia) of epiphyte plant +humus dry weight on the
tree branches.
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 27
The epiphytes, but much more the tepui plants, would
profit from ISA and even from undissolved iron oxides be-
cause plant roots and fungal hyphae secrete iron-solubilizing
organic acids and complexants. Microbial ferments have
enough time to turn all kinds of undissolvable Fe into dis-
solvable Fe.
Is there a climate relevance to rain forest fertilizing by
dust? Rizzolo et al. (2016) state that the iron-limited Ama-
zon rain forest profits from the seasonal deposition of iron by
Saharan dust. In particular, the deposition of iron plus other
nutrients on Amazon biota is likely to increase both epiphytic
growth and fungal and bacterial decomposition within the
canopy (Rizzolo et al., 2016). The increase in iron bioavail-
ability is also known to increase nutrient cycling within the
forest.
Large fractions of the organic biomass produced with the
help of iron and other eolic nutrients leave the Amazon re-
gion, are transported into the South Atlantic basin, and ulti-
mately become part of the shelf and basin sediments. These
are aquatic-life plants such as water hyacinth and water fern,
and plant litter such as driftwood, leaves, and, in particular,
colloidal and dissolved humic and fulvic acids. According to
Ertel et al. (1986), the flux of dissolved organic carbon frac-
tion at Óbidos, situated about 800 km upstream of the Ama-
zon mouth, is 2 ×1013 g C yr−1.
Some rain forests, such as the Amazonian one, benefit
from sporadic dust plume fertilization of Saharan origin.
Others may profit from an artificial ISA precipitation result-
ing in a significant additional epiphyte plant growth.
5.2 Importance of iron for human food and health
All organisms on Earth ride upon a “ferrous wheel” made
up of different forms of iron that are essential for life (Pérez-
Guzmán et al., 2010). Iron is an important micronutrient used
by most organisms, including higher animals and human be-
ings, and is required for important cellular processes such as
respiration and oxygen transport in the blood. Its bioavail-
ability is of concern for all the Earth’s living organisms,
especially in aquatic ecosystems, including clear-water and
oceanic ones.
In humans, iron deficiency and anemia remain the most
common nutritional disorders in the world today (Abbaspour
et al., 2014).
The World Health Organization (WHO, 2013) states that
a lack of sufficient micronutrients such as Fe and Zn rep-
resents a major threat to the health and development of the
world population. The WHO (2013) estimates that over 30%
of the world’s population are anemic, and the percentage
is even higher in developing countries (every second preg-
nant woman and about 40 % of preschool children). Iron de-
ficiency affects more people than any other condition, and
iron deficiency exacts its heaviest overall toll in terms of ill
health, premature death and lost earnings. Iron deficiency and
anemia reduce the work capacity of individuals and of entire
populations and cause maternal hemorrhage, impaired physi-
cal and cognitive development, reduced school performance,
and lowered productivity, bringing serious economic conse-
quences and obstacles to national development.
Iron deficiency in humans has been associated with heart
failure (Avni et al., 2012; Cohen-Solal et al., 2014); gas-
tric ulceration and anemia induced by Helicobacter pylori
(Beckett et al., 2016); negative impacts on skeletal integrity
(Medeiros, 2016); and cognitive disorders (Ünüsan, 2013).
Iron deficiency in infancy leads to long-term deficits in ex-
ecutive function and recognition memory (Lukowski et al.,
2013). In experiments with animals, even if the iron and
the hemoglobin levels return to normal after treatment for
early iron deficiency, there are long-lasting cognitive, phys-
iological and hematological effects (Yehuda et al., 2008).
Thus, several strategies and technologies have been elabo-
rated on to manage iron deficiency in humans (Saini et al.,
2016), such as food fortification (adding iron to food) (Allen,
2006) and biofortification (the process of enriching the nutri-
ent content of crops, vegetables or fruit as they grow). The
WHO, FAO and UNICEF issue guidelines or recommenda-
tions on food fortification with micronutrients (WHO, FAO
and UNICEF, 2009), for instance adding ferrous sulfate, fer-
rous fumarate or iron complexes to wheat and maize flour
(from 15 to 60 ppm, depending on the regional average con-
sumption ranges and on other iron food vehicles). Biofortifi-
cation can be achieved by utilizing crop and soil management
practices to increase micronutrient concentrations in the edi-
ble crop parts (Zuo and Zhang, 2011) and can provide a sus-
tainable solution to malnutrition worldwide, as other meth-
ods, such as diversifying people’s diets or providing dietary
supplements, have proved impractical, especially in develop-
ing countries. Together with dietary modification and iron di-
etary supplementation, iron fortification (suitable food vehi-
cle containing higher levels of bioavailable iron) is the main
recommendation of the WHO to increase iron intake, im-
prove nutritional status and stop iron deficiency anemia. In-
creasing available iron levels in major staple food crops is an
important strategy to reduce iron deficiency in people. The
WHO anticipates that the benefits are substantial as timely
treatment can restore personal health and raise national pro-
ductivity levels by as much as 20%.
The biofortification of bioavailable iron in staple plants
provides a sustainable and economical tool to use in order to
relieve iron deficiency in target populations globally (Jeong
and Guerinot, 2008).
In contrast with fruit trees, where foliar iron fertilization
is generally used in chlorotic leaves, canopy, Fe fertiliza-
tion is increasingly being used in cereal crops to increase the
Fe concentration in grains, in what is called biofortification.
In these crops, which are generally treated with foliar iron
sprays when there is no leaf chlorosis, applied iron has been
shown to re-translocate efficiently to other plant organs, both
in wheat (Zhang et al., 2010) and rice (Wei et al., 2012). Zuo
and Zhang (2011) have developed strategies to increase iron
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28 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
uptake by roots and transfer it to edible plant portions, allow-
ing absorption by humans from plant food sources.
5.3 Active inhibition of methane emissions from
wetlands, lakes and sediments
Lipson et al. (2010) found that in Arctic peat ecosystem, Fe
and humic reduction competes with methanogenesis as elec-
tron acceptors and inhibits some CH4production; on the ba-
sis of conservative measurements of net Fe reduction rates,
this process is comparable in magnitude to methanogenesis.
In wet sedge tundra landscapes, Miller et al. (2015) con-
ducted experiments that showed an inverse relationship be-
tween dissolved iron and CH4concentrations and found that
net CH4fluxes were significantly suppressed following the
experimental addition of iron and humic acids. Iron and hu-
mic acid amendments significantly suppressed in situ net
methane flux.
Lipson et al. (2013) conducted experiments on two dif-
ferent ecosystems: one with permafrost and naturally high
levels of soil Fe and one with no permafrost and naturally
low levels of soil Fe. The addition of Fe(III) and humic acids
(electron acceptors) significantly reduced net CH4flux for
at least several weeks post-treatment, without significantly
altering CO2fluxes. There was no significant difference be-
tween the reduction in CH4flux caused by Fe(III) and the
one caused by humic acids. The future release of GHGs from
high-latitude wetland ecosystems can significantly be altered
by this natural and widespread phenomenon. These results
also show that the suppression of CH4flux in this type of
ecosystem can be induced by the artificial addition of Fe(III),
humic acids or other electron acceptors.
L. Zhang et al. (2009) found methanogenesis and sul-
fate reduction inhibition after ferric salt dosing of anaerobic
sewer biofilms. Similar methanogenesis inhibition and even
increases in rice productivity by ferric salt addition have been
described by others (Ali et al., 2008; Liu et al., 2012, 2016;
W. Wang et al., 2014).
Amos et al. (2012) found support for the hypothesis that
Fe(III) mediates CH4oxidation in a crude contaminated
aquifer.
Although some iron oxides such as magnetite and hematite
have different properties and may facilitate methanogenesis
by some types of microorganisms (Zhou et al., 2014), it is
worth noting that the iron solubility and bioavailability prop-
erties of the ISA are similar to the ferrihydrite which inhibits
methanogenesis in the same experiments (Zhou et al., 2014),
and in general Fe(III) reduction by methanogens contributes
to Fe(III) inhibition of methanogenesis (Van Bodegom et al.,
2004).
Experiments conducted in humid tropical forest soils,
which are also an important source of atmospheric CH4and
where Fe(III)-reducing bacteria coexist with methanogens,
show that upon the addition of acetate, the production in-
crease in CH4is much greater (67 times) than that in Fe2+
(2 times), indicating that the two processes were acetate lim-
ited and suggesting that Fe(III)-reducing bacteria were sup-
pressing methanogenesis when acetate availability is limited
(Teh et al., 2008). For Roden and Wetzel (1996) a signifi-
cant suppression of CH4production in freshwater wetlands
could be mediated by Fe(III) oxide reduction within globally
extensive iron-rich tropical and subtropical soil regimes.
All these results support the hypothesis that additional to
the many photolysis-dominated CH4depletion actions by
ISA in the troposphere, even after ISA precipitation on wet-
lands, marshes, lakes, rice paddies and shelf sediments, ISA
will inhibit the emission of CH4. The degree to which Fe(III)
reduction suppresses CH4emissions under different soil con-
ditions should be considered by regional and global models
of GHG dynamics.
No published studies were found about the biogeochem-
ical cycle of iron to the continents and land in specialized
journals such as Global Biogeochemical Cycles nor in the
chapter about the biogeochemical cycles of the latest IPCC
report, and the recent Iron Model Intercomparison Project
(FeMIP) seems to focus on ocean interactions (Tagliabue et
al., 2015; Tagliabue and Dutkiewicz, 2016).
It is now well known that in large areas of the open ocean,
iron is a key limiting nutrient and that in alkaline terrestrial
landscapes iron deficiency induces plant chlorosis. The au-
thors’ hope is that bringing together in this review seemingly
disparate lines of research from diverse disciplines will result
in a more global understanding of the global biogeochemical
iron cycle, especially over terrestrial landscapes, peat bogs
and other wetlands.
6 Estimations of the ISA demand by the ISA method
6.1 ISA can induce a significant CH4depletion
Wittmer et al. (2015a, b, 2016) and Wittmer and Zet-
zsch (2016) reported that the ISA method is very efficient
for q
Cl generation. Hence, ISA induces the depletion of GHG
methane. This results in a prior colling effect. Therefore, ISA
appears to be a very promising cooling method with technical
and economic stakes. But the answer depends strongly on the
volume of ISA to be produced and emitted. Indeed, the ISA
plume should be released high enough in the troposphere to
get sufficient distribution and residence time in combination
with q
Cl generation quantity.
Based on results of Fe-photolysis-induced q
Cl production,
Wittmer et al. (2015a) estimated the feasibility of CH4deple-
tion by NaCl-diluted ISA. Wittmer et al. found a q
Cl emission
of 1.9 ×105q
Cl cm−3at a Cl−/ Fe(III) molar ratio of 101
within the pH range of 2.1–2.3. The same q
Cl generation was
found at the suboptimal pH of 3.3–3.5 and at a Cl−/ Fe(III)
molar ratio of 51. This Cl generation is 4 times higher than
the reference which corresponds to a significant CH4life-
time reduction in the troposphere (Wittmer et al., 2015a). A
pH range of around 2 corresponds to the natural aerosol pH
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F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 29
within the oceanic boundary layer. The optimum efficiency
of q
Cl production by the photolysis of ISA corresponds to
pH 2, whatever the source of Cl−, NaCl or gaseous HCl and
if ISA is an iron(III) oxide or an iron(III) chloride aerosol
(Wittmer et al., 2015a).
According to Lim et al. (2006) and to Meyer-Oeste (2010)
the optimum q
Cl production by sunlight photolysis of FeCl3
solutions or ISA is generated in the acidic pH range. The
efficient q
Cl generation is necessary for an efficient CH4de-
pletion by ISA. FeCl3has an acidic pH from the beginning
because it hydrolyzes in accordance with Eq. (4), except if it
has come about by the condensation and hydrolysis of FeCl3
vapor, by the nebulization of pure FeCl3solution, or by com-
bustion to pyrogenic FeOOH and reaction and hydrolysis
with HCl and H2O to FeCl3solution.
FeCl3+2H2O→FeCl2OH +H3O++Cl−(4)
6.2 ISA demand calculation
Current CH4depletion by q
Cl is estimated to be 3.3 % (Platt
et al., 2004) to 4.3 % (Allan et al., 2007). According to the
results of Wittmer et al. (2015a), at a Cl−/ Fe(III) molar ratio
of 101, this amount would rise 4-fold from 13 to 17 %.
Wittmer et al. (2015a) used their results obtained
at a Cl−/ Fe(III) ratio of 51 at the pH of 3.3–3.5:
1.9 ×105q
Cl cm−3. We consider that this pH is suboptimal.
Instead the results obtained at a Cl−/ Fe(III) ratio of 101 at
the pH of 2.1–2.3: 1.9 ×105q
Cl cm−3should be used.
Moreover, Wittmer et al. (2015a) made two limitative es-
timations:
–They only focused on the Cl delivery in the condensed
state by coagulation as a Cl−transfer option between
ISA particles and the Cl source sea-salt aerosol, ignor-
ing other Cl sources, Cl aggregate states and Cl transfer
mechanisms.
–According to this model, the ISA particles should con-
tinuously lose their Cl−load by q
Cl emission in the day-
light, and as a consequence they could only gain back Cl
by coagulation with sea-salt aerosol particles. As further
consequences of this model, the Cl−/ Fe(III) ratio of
ISA particles would decrease, their diameter would in-
crease and their residence time in the troposphere would
decrease.
But according to Graedel and Keene (1996) and Keene et
al. (1999), the next prominent source of inorganic Cl in the
troposphere beside sea-salt aerosol is vaporous HCl. This is
the main source from which the ISA particles can refill the
chloride lost by photolysis. The main Cl uptake mechanism
from this Cl source is sorption from the gaseous phase.
The main HCl sources are the sea-salt reaction with acids,
CH4and other hydrocarbon reactions with q
Cl (Keene et al.,
1999), flue gases of coal, and biomass and garbage combus-
tion (McCulloch et al., 1999), as shown in the global Re-
active Chlorine Emissions Inventory (Keene et al., 1999),
HCl from chlorocarbons being a significant part (Sanhueza,
2001), in particular from CH3Cl, which is the largest, natural
contributor to organic chlorine in the atmosphere (Lobert et
al., 1999). Wittmer et al. (2015a) estimate that the global pro-
duction rate of 1785 Tgyr−1of sea-salt aerosol Cl−has to be
doped with iron at a Cl−Fe(III) molar ratio of 51; however,
we consider that it has to be estimated at a molar ratio of 101
(according to the first point above).
The calculations made with these limitative assumptions
resulted in an iron demand of 56 Tg yr−1Fe(III) to obtain
the desired CH4depletion effect (Wittmer et al., 2015a).
However, with the limitative assumption that there is no Cl−
source other than sea salt, the calculations with a Cl−/ Fe(III)
ratio of 101 result in an Fe(III) demand of only 18 Tg yr−1.
ISA can be produced from pyrogenic iron oxides accord-
ing to method I (see Sect. 7). Pyrogenic oxides have parti-
cle sizes lower than 0.1 µm. Diameters of the NaCl-diluted
ISA particles of the Wittmer tests (Wittmer et al., 2015a) are
round about 0.5 µm. This confirms the test results of Wittmer
et al. (2015a) as a calculation basis.
But Wittmer et al. (2015a) made two other limitative as-
sumptions:
–ISA has the same particle size and corresponding sur-
face range as sea salt.
–ISA has the same residence time as sea-salt aerosol in
the troposphere.
Due to their coarse aerosol particle range, the residence
time of sea-salt particles in the troposphere is inferior to
1 day (Jaenicke, 1980), while the artificial ISA particles with
diameters lower than 0.5 µm have residence times in the tro-
posphere from at least 10 days to several weeks (Kumar et
al., 2010; Penner et al., 2001).
Known salt aerosol generation methods by vapor con-
densation or nebulization (Biskos et al., 2006; Gupta et
al., 2015) allow not only the flame-descending ISA type 1
(Oeste, 2004) but also the condensation- and nebulization-
descending ISA variants 2 and 3 (see Sect. 7) to be produced
with aerosol particle diameters between 0.1 and 0.01 µm. Di-
ameters of salt aerosol particles from these physical aerosol
generation methods are up to, or more, than 1 order of mag-
nitude smaller than of those used in the experiments by
Wittmer et al. (2015a).
Analogously to CCN behavior in cloud processing
(Rosenfeld et al., 2008), most of the small-sized ISA parti-
cles are protected by their small sizes from coagulation or co-
alescence with sea-salt aerosol particles. This effect prevents
ISA from leaving the optimum active atomic chlorine emis-
sion conditions: a low pH and low particle diameter range.
The residence time difference of more than 1 order of mag-
nitude in comparison to sea-salt aerosol further reduces the
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30 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
Fe demand for ISA production from 18 Tg yr−1to less than
1.8 Tg yr−1.
The properties of the ISA particles produced by the most
preferred ISA method variant are explained in Sect. 4. The
difference between them and the NaCl-diluted ISA tested
by Wittmer et al. (2015a) is that ISA particles are made
of FeCl3×nH2O, undiluted by NaCl, or FeOOH coated
by FeCl3×nH2O, undiluted by NaCl (Meyer-Oeste, 2010;
Oeste, 2009). The Cl−/ Fe(III) molar ratios of FeCl3×nH2O
are at 3 or even lower. The Cl−/ Fe(III) molar ratio of typi-
cal ISA particles is at least 30 times smaller than the mo-
lar Cl−/ Fe(III) ratio of 101 of the ISA tested by Wittmer et
al. (2015a). This reduces the Fe demand for ISA production
again at least by 1 order of magnitude from <1.8 Tg yr−1to
about <0.2 Tg yr−1.
Wittmer et al. (2015a) considered only sea-salt aerosol
particles as transport vehicles for ISA and as the only pos-
sible contact medium to gain chloride ions as q
Cl source. It is
well known that coal combustion is a major source of active
chlorine (Keene et al., 1999; McCulloch et al., 1999; San-
hueza, 2001) as well as iron (Ito et al., 2016; Luo et al., 2008;
Sedwick et al., 2007; Wang et al., 2015b); thus, both iron and
chlorine are jointly issued by other mechanisms and sources.
As stated in our Sect. 6.2, sea-salt aerosol has residence
times in the troposphere of less than 1 day depending on its
coarse-particle diameters; it is not possible to for aerosol to
cross intercontinental distances in this time.
In reality the chloride transfer between sea-salt aerosol
particles and ISA particles may take place without any con-
tact or coagulation because the troposphere is an acidic en-
vironment. The troposphere is a source of organic and in-
organic acids, which are in permanent contact with sea-salt
aerosol. The acid ingredients in contact with sea spray pro-
duce HCl. Furthermore, ISA is produced by combustion and
is elevated by flue gas plumes: acid precursors such as SO2
or NOxhave higher concentrations within the flue gas plume
compared to the tropospheric environment. The acids gen-
erated by flue gas plume produce additional HCl by reaction
with the sea-salt aerosol (von Glasow, 2000). As a result, ISA
and ISA precursors may absorb any chloride requirement via
HCl vapor from the sea spray source itself (Wittmer and Zet-
zsch, 2016).
Additionally to the q
Cl emission increase with increas-
ing iron concentration in the aerosols tested, the results of
Wittmer et al. verify an increase in q
Cl emission with de-
creasing pH (Wittmer et al., 2015a). According to Wittmer
and Zetzsch (2016), Meyer-Oeste (2010) and Oeste (2009),
oxidic ISA aerosol particles may be generated free of any
pH-buffering alkaline components. This hampers their pH
decrease by air-borne HCl to the optimum pH of around 2.
Sea-salt buffering of the absorbed HCl (Sullivan et al., 2007)
by the alkali and earthen alkali content of sea-salt aerosol
can occur only by coagulation, most probable in a minor ISA
particle fraction but not in the bulk. From the beginning of
its action in the troposphere, ISA remains in the optimum q
Cl
emission mode: low pH and high iron concentration levels.
Preferred ISA is produced by the ISA method variant 1
or variant 3 as described in Sect. 7. Hence, ISAs are com-
posed of particles made by flame pyrolysis or iron salt vapor
condensation. The ISA particles mentioned have diameters
of 1/10 of the particle diameters of the Wittmer tests. These
ISA particles have an optimum chlorine activation efficiency
–in an appropriate chloride content or chloride delivering
environment;
–at a pH <2;
–if they are emitted above the tropospheric boundary
layer.
In such cases, the Fe demand may even fall short of the cal-
culated 0.2 Tg Fe yr−1due to the particles’ greatly extended
surface area and greatly extended residence time in the atmo-
sphere.
It has to be noted that this ISA demand calculation re-
sult refers only to the ISA cooling property depending on
CH4depletion; other cooling properties depending on cloud
albedo, the depletion of CO2, black and brown aerosol, ozone
decrease, and other causes are still unconsidered.
Other oxidation activity on GHGs and aerosols is induced
by the q
OH generation activity of ISA: volcanic eruption
plumes contain high concentrations of q
Cl plus q
OH (Baker et
al., 2011) and are characterized by decreased CH4concentra-
tions (Rose et al., 2006). The co-absorption of H2O and HCl
is the main reason for the generation of volcanic ash particle
coats containing soluble-Fe salts originating from insoluble-
Fe oxides and Fe silicates (Hoshyaripour et al., 2015; Martin
et al., 2012). Gaseous HCl from the eruption plume entails
Fe chlorides covering the surfaces of volcanic ash particles
(Ayris et al., 2014). Therefore, it is reasonable that the pho-
tolysis of those chlorides is the origin of both: q
Cl and q
OH
generation in volcanic plumes.
Hydroxide radical q
OH can change from the liquid aerosol
phase into the gaseous phase (Nie et al., 2014) but not nearly
as easily as q
Cl can. Indeed, the Henry’s law solubility con-
stant of q
OH is about 1 order of magnitude higher than that
of q
Cl and is in the same range as that of NH3(Sander, 2015).
But when their hygroscopic water layer shrinks in dry air or
by freezing, ISA particles might act as q
OH emitters. These
additional q
OH emissions might further increase the CH4ox-
idation potential of volcanic ash or artificial ISA and thus re-
duce even more the Fe demand for ISA; though this has not
been tested yet, it cannot be ruled out.
In order to take care not to increase the cooling effect too
much, a reasonable goal might be to start the ISA method
with a global ISA emission of 0.1 Tg Fe yr−1. This quantity
corresponds to the magnitude of the actual Fe input from the
atmosphere into the oceans in the form of soluble salt, which
is estimated to be from 0.1 up to 0.26 Tg yr−1(Ito and Shi,
Earth Syst. Dynam., 8, 1–54, 2017 www.earth-syst-dynam.net/8/1/2017/
F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 31
2016; Johnson and Meskhidze, 2013; Myriokefalitakis et al.,
2014). Doubling or even tripling this input quantity by the
ISA method is technically easy and economically feasible as
will been seen in Sect. 7.
7 The ISA method: how to increase artificial iron
emissions
The preceding calculation shows that the ISA method has
the potential to cut back on the rise in CH4and CO2and the
small decline in atmospheric oxygen content (Keeling and
Shertz, 1992; Manning and Keeling, 2006) because it acts
through a bundle of chemical and physical means. The ISA
method might retard, stop or even help to restore these GHG
contents to preindustrial levels. By the ISA method, doubling
or tripling the ISA level in the troposphere seems to be pos-
sible by feasible technical and economical means.
Since 2004 proposals have been published (Meyer-Oeste,
2010; Oeste, 2004, 2009, 2015; Oeste and Ries, 2011) to
modify combustion processes and flue gas emissions from
traffic and power-generating combustions and their warm up-
lifting flue gases in order to use them as ISA plume emission
sources in the troposphere. Any hot flue gas plumes emitted
by ship and air traffic and fossil and solar power are predes-
tined for the ISA method.
At least three variants of ISA production are proposed.
–Variant 1: the emission of flame pyrolytic FeOOH
aerosol with particle diameters smaller than 100 nm
(Buyukhatipoglu and Clyne, 2010; Kammler et al.,
2001) as an ISA precursor by co-combustion of organic
iron or carbonyl iron additives with liquid or gaseous fu-
els or heating oils combusted in ship or and jet engines
or by oil or gas combustors. The co-combustion of iron
compounds is a possible measure in coal power stations
by mixing the ISA-precursor-containing “oil” combus-
tion flue gas to the coal combustion flue gas after the
dry flue gas cleaning stage. Useful side effects of iron
additives are fuel efficiency optimization and soot emis-
sion minimization (Kasper et al., 1998; Kim et al., 2008;
Madhu et al., 2015; Weiser et al., 2007). The emitted
FeOOH aerosol plumes convert into the ISA plume im-
mediately after leaving the emission sources due to the
high reactivity of flame pyrolytic Fe oxides. The time it
takes to cover the flame-pyrolytic FeOOH particle sur-
face with Fe(III) chlorides through HCl absorption from
the gaseous phase is several times shorter that than nec-
essary for the generation of iron chlorides from natu-
ral iron oxide minerals in loess dust particles (Rubas-
inghege et al., 2010; Sullivan et al., 2007).
–Variant 2: the injection of vaporous ISA precursor iron
compounds such as FeCl3into a carrier gas. By coming
into contact with the carrier gas and/or the atmosphere,
the vaporous iron compounds condense and/or convert
by physical and/or chemical means directly into ISA.
Contrary to all other ISA precursors, the sunlit FeCl3
vapor is photoreduced by concomitant generation of q
Cl
(Rustad and Gregory, 1980). Thus, methane-depleting
q
Cl emission can start even before this ISA precursor
has changed into hydrated FeCl3.
–Variant 3: the injection of an ultrasonic nebulized aque-
ous FeCl3solution as an ISA precursor into a carrier
gas. ISA is generated by water evaporation from the
aerosol droplets.
The preferred heights of ISA plume generation in the tro-
posphere are 1000 m above ground or at higher altitudes in
order to pass the boundary layer. There, the ISA plumes
have optimum conditions to spread over a large area due
to sufficient lifetimes. The necessary buoyancy to lift up the
ISA plumes can be regulated by controlling their carrier gas
temperatures. Uplift towers (Ming, 2016), vortex generators
(Michaud and Renno, 2011) or tethered balloons (Davidson
et al., 2012; Kuo et al., 2012) are preferential means to direct
ISA by carrier gas uplift to the heights mentioned.
The primary ochre-colored FeOOH aerosol particles emit-
ted by ISA method I have diameters of <0.05 µm. According
to previous studies, iron oxides are strong absorbers at visi-
ble wavelengths and might play a critical role in climate per-
turbation caused by dust aerosols (Sokolik and Toon, 1999;
X. L. Zhang et al., 2015). But this effect is not applicable
to the ISA methods FeOOH aerosol because it is emitted by
flue gas plumes generated in parallel and containing SO2and
NOxas sulfuric and nitric acid generators. Due to their small-
diameter-dependent high surface area the aerosol particles
immediately react with HCl. HCl is generated by the reaction
between sea-salt aerosol and flue-gas-borne acids.The pri-
mary reaction product is the orange-colored FeCl3aerosol:
ISA. But daytime sun radiation bleaches ISA by FeCl2and
q
Cl generation; nighttime reoxidation of ISA plus HCl ab-
sorption regenerates FeCl3again. FeCl2is colorless at low
humidity and pale green at high humidity.
The provision of phytoplankton to optimize its growth
with other nutrients such as Mn, Zn, Co, Cu, Mo, B, Si and
P by the ISA method is possible by at least variants 1–3 of
the ISA method, i.e., by co-combustion, co-condensation or
co-nebulizing.
Global fixing regulations of GHG emission certificate
prices, values and ISA emission certificate credit values
would be simple but effective measures for the quickest
worldwide implementation of the ISA flue gas conditioning
method.
Anderson (2016) mentioned that “the complete set of 400
IPCC scenarios for a 50 % or better chance of 2 ◦C assume
either an ability to travel back in time or the successful and
large-scale uptake of speculative negative emission technolo-
gies”. He refers to going back in time because some scenar-
ios assume “the successful implementation of a stringent and
global mitigation regime in 2010.”
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32 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
A large part of the research devoted to climate engineer-
ing methods concerns SRM (sunlight reduction methods),
such as mimicking the effects of large volcanic emissions by
adding sulfate aerosols to the stratosphere as suggested for
instance by Crutzen (2006). Numerous other types of parti-
cles have been suggested for these aerosols, for instance tita-
nia by Jones et al. (2015). But SRM only buys time and has
numerous drawbacks.
On the one hand, SRM does not address the main cause of
global warming (GHG emissions) nor does it prevent ocean
acidification. On the other hand, several carbon dioxide re-
moval (CDR) technologies do, but their costs are much larger
than SRM and the scale requested poses many technological
challenges; see, for instance, “Scaling up carbon dioxide cap-
ture and storage: from megatons to gigatons” (Herzog, 2011).
Very few CDR methods without the emission of disadvan-
tageous pollution are known. One of those is the terra preta
method: it is characterized by the mixing of ground biochar
into agricultural soils. The climate relevancies of this method
are a sustained fixation of former CO2carbon, minimization
of fertilizer consumption and N2O emission reduction from
the fertilized terra preta soils. Char has similar properties
within the soil environment as humic substances, but in the
environment, char is resistant against oxidation.
Comparing the terra preta method to other CDR methods,
such as fertilizing the ocean by micronutrients, shows that
lower specific material expenses are needed by CDR meth-
ods per unit of CO2removed from the atmosphere (Betz et
al., 2011). The ISA method we propose is a member of this
CDR group; thus, this result is also valid. In addition, the
further climate effects of the ISA method (such as the deple-
tion of CH4, tropospheric ozone and soot, plus cloud whiten-
ing) reduce the specific material expense level. Furthermore,
the ISA method mimics a natural phenomenon (mineral iron-
dust transport and deposition) and only proposes to improve
the efficiency of an already existing anthropogenic pollution.
Myriokefalitakis et al. (2015) estimates that “The present
level of atmospheric deposition of DFe [dissolved iron] over
the global ocean is calculated to be about 3 times higher than
for 1850 emissions, and about a 30 % decrease is projected
for 2100 emissions. These changes are expected to impact
most on the high-nutrient–low-chlorophyll oceanic regions.”
Their model “results show a 5-fold decrease in Fe emis-
sions from anthropogenic combustion sources (∼0.013 Tg-
Fe yr−1in the year 2100 against ∼0.070 Tg-Fe yr−1in the
present day), and about 45 % reduction in mineral-Fe dis-
solution (∼0.078 Tg-Fe yr−1) compared to the present day
(∼0.175 Tg-Fe yr−1).”. Meanwhile the model used by Mis-
umi et al. (2014) predicts an iron supply increase to high-
nutrient, low-chlorophyll (HNLC) surface waters by 2090,
especially in the eastern equatorial Pacific, attributed by the
authors to changes in the meridional overturning and gyre-
scale circulations that might intensify the advective supply
of iron to surface waters. Furthermore, several authors (An-
nett et al., 2015; Bhatia et al., 2013; Hawkings et al., 2014;
Raiswell et al., 2008, 2016) point out that both glacial and
deep-water Fe sources may increase with continued climate
warming due to Fe input from other sources, such as shelf
sediments, meltwater, icebergs, rivers, surface water runoff
and dust input.
Recently, Boyd and Bressac (2016) suggested rapidly
starting tests to determine the efficiency and side effects of
CDR ocean iron fertilizing methods and analyzed possible
geopolitical conflicts together with some other geoengineer-
ing methods (Boyd, 2016).
Several experts, for instance Hansen et al. (2016), recently
expressed the urgent warning that mankind only has a short
time left to address and control climate warming. As a con-
sequence, it is necessary for mankind to find climate control-
ling substances which might generate the most effective and
reversible climate cooling effects within the shortest period
as soon as possible. Lifetimes of ISA emissions in the tropo-
sphere are much shorter than those of sulfates in the strato-
sphere. Of course, such tools and agents have to be rapidly
evaluated against side effects to ecosystems, human health
and, last but not least, their economic burdens.
8 Interaction of the ISA method with other measures
to protect the environment
According to Wittmer and Zetzsch (2016), an elevated HCl
content in the atmosphere triggers the methane-depleting
coating of oxidic ISA precursors by photolytic active Fe(III)
chlorides. Any measure triggering the reduction in the HCl
content of the atmosphere would impair the effectiveness of
the ISA method based on this kind of method.
In this sense all kinds of measures to reduce the sulfur and
NOxcontent of the flue gas content of gaseous, liquid and
gaseous fuels would decrease the effectiveness of oxidic ISA
precursors, as the S and NOxoxidation products sulfuric acid
aerosol and gaseous nitric acid are the main producers of HCl
by changing sea-salt aerosol into sulfate and nitrate aerosol.
Even the measures of reducing the energy production from
fuel burning by changing to wind and photovoltaic energy
would reduce this HCl source.
Sea-salt aerosols produce HCl after contact with organic
aerosol and organic volatile matter as the organic compounds
generate acid oxidation products such as oxalic acid (Drozd
et al., 2014; Laskin et al., 2012; B. Wang et al., 2014). A large
fraction of organic aerosols and secondary organic aerosols
originates from anthropogenic sources such as combustion.
The change to wind and photovoltaic energy would reduce
this HCl source.
The proposed climate engineering measure of producing
sulfuric acid aerosol within the stratosphere by inducing an
albedo increase would increase the HCl content, through
contact of the precipitating acid aerosol with tropospheric
sea-salt aerosol. Even the proposed climate engineering mea-
sure of increasing the sea-salt aerosol content of the tropo-
Earth Syst. Dynam., 8, 1–54, 2017 www.earth-syst-dynam.net/8/1/2017/
F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 33
sphere by artificial sea-salt aerosol as a cloud whitening mea-
sure could be used as an ISA method trigger if flue gas is used
to elevate the sea-salt aerosol.
9 Discussion
In order to fight global warming, this review proposes to en-
hance the natural actions of Cl atoms in the troposphere, to-
gether with the synergistic action of iron in the atmosphere,
ocean, oceanic sediment and land compartments, as a climate
engineering method. The main results expected are a diminu-
tion of long-lived well-mixed atmospheric methane and car-
bon dioxide, but the diminution of local short-lived tropo-
spheric ozone is also possible, as well as effects on the Earth
albedo, the restoration of the oxygen flux into the deep-ocean
basins, organic carbon storage, etc.
The most important actor in the process of CO2-C trans-
fer from the atmosphere to the Earth interior is the carbon-
ate C precipitation in the crust rocks and sediments below
the ocean. The ocean crust acts like a conveyor belt be-
tween crust evolution at MOR and its subduction zones into
the mantle. The media transported are carbonate C, small
amounts of organic C, ocean salt, ocean water and sediments.
This process is part of the homeostasis of the planet. Distur-
bances of this system are induced by stratification processes
within the ocean basins caused by density differences be-
tween different layers of the water column. Most stratifica-
tion events are induced by climate warmings. Any of these
homeostasis disturbances are removed by the system on ge-
ological timescales. Signs of such disturbances are more or
less prominent events of extinction and of elevated organic
C content in the ocean sediments. Because the recent cli-
mate warming will induce a new ocean stratification event,
mankind must stop it. Like several interglacial stratification
events in the glacial periods, the actual stratification is also
induced by increasing meltwater discharge. The past inter-
ruptions of the interglacial climate warmings teach us that
the interruption events were accompanied, as a rule, by dust
events. As demonstrated, the climate cooling effects of these
dust events are induced by the chemical and physical actions
of ISA.
In high-nutrient, low-chlorophyll oceanic areas, where the
contribution of atmospheric deposition of iron to the surface
ocean could account for about 50 % of C fixation, as well
as in oceanic nitrogen-limited areas, where atmospheric iron
relieves the iron limitation of diazotrophic organisms (thus
contributing to the rate of N fixation), atmospheric deposi-
tion of iron has the potential to augment atmospherically sup-
ported rates of C fixation (Okin et al., 2011) and thus “cool
the Earth” by removing CO2from the atmosphere.
Maybe the iron atmospheric deposition over terrestrial
landscapes and wetlands has similar effects? Are there pos-
sible benefits of atmospheric deposition of soluble iron over
the continents in which iron deficiency in plants occurs over
30 % of their area and which have high pH calcareous soils
that make soil Fe unavailable for plants (Abadía et al., 2011)?
Iron-deficiency-induced chlorosis in plants can be solved by
the addition of soluble-iron complexes to the soil or by fo-
liar application of sprays containing mineral iron (for in-
stance FeSO4) (Basar et al., 2014) or iron chelates (Fe-EDTA
among others) (Fernández et al., 2013). Iron, sulfate and sev-
eral organic iron complexes such as iron oxalate are known
constituents of atmospheric dust (Johnson and Meskhidze,
2013), but unfortunately no published work was found about
possible effects on plant chlorosis by foliar deposition of sol-
uble iron from atmospheric dust.
We did not find studies about the impacts of atmospheric
iron nutrient deposition on terrestrial ecosystem productiv-
ity. More research is needed to continue to enhance our un-
derstanding of the possible benefits of iron cycling in fresh-
water and terrestrial landscape environments, as well as in
atmospheric and sediment environments, in particular of its
numerous potential capacities to fight global warming. The
cooling effects of ISA and iron reviewed in this article al-
ready provide an insight into the progress made on under-
standing the iron cycles from a range of perspectives.
There is abundant literature on the many geoengineering
methods that have been proposed to cool the Earth (Lack-
ner, 2015; Z. Zhang et al., 2015). In particular, the injection
of sulfate aerosols into the stratosphere is the most studied
method, as it mimics the episodic action of natural volcanoes
(Ming et al., 2014; Pope et al., 2012). Injected particles into
the stratosphere reduce the radiative balance of Earth by scat-
tering solar radiation back to space, so several types of par-
ticles are envisioned with a wide range of side effects (Jones
et al., 2015).
The literature also describes many options to deliver sul-
fates, their precursors (or other particles), to the stratosphere
(Davidson et al., 2012). For instance, airplane delivery of the
sulfate aerosols by the kerosene combustion process requires
military jets due to commercial aircrafts’ limited altitude of
10 km (30 000 feet) where 20 km are required (Davidson et
al., 2012).
In the case of ISA, the altitude needed to cool the Earth is
much lower (it is in the troposphere), and the total quantities
to deliver are 1 order of magnitude smaller. So air travel is
a possible means for ISA delivery. But the global jet fuel
consumption is only about 240 000 t yr−1. Even by assuming
the very high emission rate of 1 kg ISA precursor iron per
ton of jet fuel, only 24 t yr−1might be emitted. This seems
far from the order of magnitude of the target ISA emissions.
From the many other possible delivery strategies envi-
sioned for SRM by stratospheric aerosols, many are not
suited for ISA, such as artillery, missiles and rockets (David-
son et al., 2012); it would be cheaper and cause less pollution
to use the flue gas of a reduced number of thermal power
plants. This might be efficient enough to deliver the artifi-
cial iron aerosol needed over the boundary layer in order for
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34 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
aerosols to stay in the troposphere for several days or weeks
and become widely distributed (Williams et al., 2002).
According to Luo et al. (2008), the deposition of soluble
iron from combustion already contributes 20 to 100% of the
soluble-iron deposition over many ocean regions.
As an example, we calculated the possible production and
emission of the ISA precursor FeOOH aerosol using the flue
gas of the German power station Niederaußem; with an input
of 25 million t yr−1of lignite (brown coal), this power station
produces 3600 MW.
According to ISA production variant 1 (Sect. 6), the ISA
precursor FeOOH aerosol may be produced by the burning
of a ferrocene (Fe(C5H5)2) oil solution containing 1 % fer-
rocene in a separate simple oil burner. The hot oil-burner
flue gas containing the ISA precursor FeOOH aerosol is in-
jected and mixed into the cleaned power station flue gas.
The power station flue gas emission rate is calculated to
be 9000 m3flue gas per ton of lignite. As the ISA precur-
sor containing flue gas will be elevated to heights of more
than 1000 m above ground, dust levels of the ISA precur-
sor FeOOH aerosol of 20 mg m−3flue gas seem to be ac-
ceptable. This allows a quantity of 180 g of FeOOH per
ton of combusted lignite (9000 m3t−1×0.02 g m−3). At
a lignite quantity of 25 million t yr−1, this corresponds to
4500 t FeOOH yr−1. FeOOH has an iron content of 63 %.
This corresponds to a possible iron emission of 2831 t yr−1
and a possible ferrocene consumption of 9438 t yr−1.
Corresponding to this calculation, about 100 of such huge
power stations should have the ability to produce a sufficient
ISA quantity of an equivalent of 200000 to 300000tFeyr−1.
Further optimization of the cooling capacity of the produced
ISA is possible by a co-emission of HCl, for instance by co-
burning an organic HCl precursor.
This example illustrates that ISA emission at only
100 power stations, or any similar ISA emission measures,
is quite feasible compared to the alternative of carbon cap-
ture and storage (CCS) through CO2capture of 40 Gt yr−1
from the flue gas and compression of the CO2until the liq-
uid state is reached, followed by the transportation and CO2
storage by injection into underground rock aquifers or into
old and depleted fossil fuel reservoirs.
In order to increase the effectiveness of the buoyancy ca-
pacity of the power stations the usual wet cooling tower
might be replaced by a dry cooling tower to mix the dry and
warm air emission from the cooling tower with the hot flue
gas as additional buoyancy and dew point reduction mean.
Furthermore, the flue gas buoyancy may increase by increas-
ing the flue gas temperature. This or other simple techniques
to realize ISA plumes may be used within the troposphere.
One alternative delivery method that seems promising and
can easily be adapted to the ISA method is the use of teth-
ered balloons (Davidson, 2012), which will cost much less
as 1 or 2 km altitude will be sufficient for ISA emissions,
requiring much lower pressures in the pipes than for SO2de-
livery at 20 km forthe geoengineering method. Technical and
economic feasibility have already been studied for the SPICE
project (Kuo et al., 2012), which planned to release seawater
spray at 1 km altitude.
Furthermore, iron emissions only stay in the troposphere
for weeks compared to SRM sulfates in the stratosphere,
which stay for 1 or 2 years. In the case of any uninten-
tional side effect or problem occurring, stopping the emis-
sions quickly is possible and the reversibility of their effects
are much shorter than for solar radiation management by sul-
fate aerosols.
Other geoengineering strategies to cool the Earth, such as
carbon dioxide removal by iron fertilization (Williamson et
al., 2012) have several pros and cons, such as localized re-
lease and less dispersion, in a form that is not readily bio-
available, resulting in restricted cooling effects and high ex-
penses.
The idea of ocean fertilization by iron to enhance the CO2
conversion by phytoplankton assimilation arose within the
last 2 decades. The mixing of an iron salt solution by ships
into the ocean surface was proposed. This idea was debated
controversially. An example of this debate is the discussion
between Johnson and Karl (2002) and Chisholm et al. (2002).
Deeper insight into this debate is given by Boyd and Bres-
sac (2016).
The iron fertilization procedure tests done so far have been
restricted to relatively small ocean regions (Boyd et al., 2007;
Johnson et al., 2002a, b). These tests produced iron concen-
trations some orders of magnitude above those produced by
natural ISA processing, which are on the order of tens of
milligrams of additional dissolved iron input per square me-
ter per year. In this sense the ISA method is quite differ-
ent from iron fertilization. As known from satellite images,
phytoplankton blooms are induced by natural dust emission
events from the Sahara, Gobi and other dust sources; there
is no doubt about the fertilizing effect of iron. Moreover,
this kind of natural iron fertilization enhancing the transfer
of CO2C into organic sediment carbon via the oceanic food
chain seems to be uncontradicted and accepted (Hansen et
al., 2016).
The ISA method allows the use of the same atom of iron
several times by catalytic and photocatalytic processes in the
atmosphere, with different cooling effects (such as albedo
modification and the enhancement of the methane destruc-
tion), and then through gravity or by rain precipitation, it
sinks and reaches the ocean surface, with further cooling ef-
fects such as the enhancement of CO2carbon fixation.
Harrison (2013) estimates that a single ship-based fertil-
ization of the Southern Ocean will result only in a net seques-
tration of 0.01 t C km−2for 100 years at a cost of USD 457
per ton of CO2, as the economic challenge of distributing
low concentrations of iron over large ocean surface areas has
been underestimated (Aumont and Bopp, 2006); in addition
to this, there are the numerous loss processes (i.e., soluble-
iron loss and organic carbon that does not sink to the bottom
Earth Syst. Dynam., 8, 1–54, 2017 www.earth-syst-dynam.net/8/1/2017/
F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 35
of the ocean), resulting in reduced net storage of carbon per
square kilometer of ocean fertilized.
Figure 7 summarizes many of the cooling effects of the
ISA method.
The organic C / carbonate C burial ratio in sediments and
bedrock increase after the ISA method starts, until it reaches
a maximum. Then this ratio begins to decrease as soon as the
vertical current components in the ocean basin begin to act.
Then the ratio reaches a very low permanent level, while the
total of buried C reaches a permanent maximum level when
the maximum vertical mixing conditions have been achieved
by the ISA method.
Why does ISA appear to be more effective than ocean
iron fertilization? For ocean iron fertilization several tons
of Fe(II) are dispersed in a short time (hours) over only
some square kilometers of ocean with several drawbacks;
a massive algae bloom can change the local biotopes. By
contrast, ISA releases iron continuously, reaching the en-
tire 510 million km2of Earth’s surface. The current iron in-
puts (in the form of soluble salts) into the oceans are es-
timated to be between 0.1 and 0.26 Tg yr−1(Ito and Shi,
2016; Johnson and Meskhidze, 2013; Myriokefalitakis et
al., 2014). As water covers nearly 72% of Earth’s surface
(362 million km2), if ISA evenly distributes 1TgFeyr−1(in
addition to current natural and anthropogenic emissions),
which is 4 times more than the expected needs (Sect. 5.2),
on average every square kilometer of ocean will receive
5.4 g Fe km−2day−1(1/510 t Fe km−2yr−1).
10 Conclusion
In ideal circumstances the ocean acts as an optimum trans-
port medium for CO2carbon from the atmosphere into the
ocean crust. Such circumstances are present when the verti-
cal cycling components between the ocean surface and ocean
bottom are undisturbed.
Any stratification event disturbs this cycling and inter-
rupts CO2transport. Climate warming can induce stratifica-
tion events by producing huge amounts of meltwater. Recent
research found signs of at least a regional development of a
beginning stratification.
The numerous climate cooling effects of natural dust
shown in this review, according to its soluble-iron content,
demonstrate that dust is of a central significance as a steering
element of this carbon transport from the atmosphere into the
ocean crust.
This review article demonstrates the enormous effects of
atmospheric iron dusts and focuses first on the tropospheric
aerosol particles composed partly of iron and chloride (iron
salt aerosols, ISAs), showing their cooperation and inter-
actions with several components of the atmosphere, for in-
stance with CH4, as the chlorine atom is responsible for the
removal of a significant part of this GHG (3 to 4% of CH4)
in the troposphere (Allan et al., 2007; Graedel and Keene,
1996). This article summarizes a dozen of other possible di-
rect and indirect natural climate cooling mechanisms induced
by iron biogeochemistry in all the Earth compartments: the
atmosphere, oceans, land (surface, soil), the sediment and the
crust.
These dozen possible climate cooling effects due to the
multistage chemistry of iron within the atmosphere, hydro-
sphere, geosphere and lithosphere are described together for
the first time and are summarized in Table 3, which shows
the most probable climate cooling effects of ISA. They in-
clude the ocean fertilization effect, which allows enhanced
algal and phytoplankton growth, which in turn removes min-
eral CO2from the atmosphere and transforms it in organic
carbon, part of which can sink to the bottom of the oceans
and be stored for long periods of time by the different mech-
anisms that are described.
In order to explicitly address the interaction of climate and
biogeochemistry, the complex interactions between climate
and the cycles of C, N, P, H2O and micronutrients call for
models that integrate global biogeochemical cycles of terres-
trial, oceanic and atmospheric components of the biosphere.
While the iron biogeochemical cycle between the atmo-
sphere and the ocean is considered in numerous publications,
the treatment of key processes and feedbacks within the ter-
restrial compartment has been rather limited, and further de-
velopment is urgently needed.
Mineral dust aerosols containing iron and other impor-
tant nutrients or micronutrients are well-studied components
of the iron biogeochemical cycle in the atmosphere and the
oceans, but the absence of recent publications about the full
iron biogeochemical cycle over terrestrial landscapes, soils,
wetlands and all clear-water compartments (glaciers, ice,
snow, lakes and groundwater) points to a lack of an up-to-
date overview. In our opinion, atmospheric chemistry mod-
els need to incorporate all relevant interaction compartments
of the Fe cycle with sun radiation, chlorine, sulfur, nitrogen,
oxygen, carbon and water in order to model several planetary
cooling effects of the iron cycle.
Acid rain sulfate (SO2−
4) deposition on peatlands and wet-
lands from natural sources (volcanoes), or anthropogenic
sources (fossil fuel combustion) is a known suppressant of
CH4production (Nedwell and Watson, 1995; Watson and
Nedwell, 1998) and emissions (Gauci et al., 2002, 2004,
2005) and may be an important process in terms of global
climate. The importance of the Fe input associated with an-
thropogenic aerosol deposition in terrestrial biogeochemistry
deserves further investigation as do the possible impacts of a
drastic diminution of anthropogenic iron and sulfate emis-
sions from combustion processes expected by 2050 to satisfy
the Paris climate agreement.
This review completes the previous global iron cycle pre-
dictions (Archer and Johnson, 2000; Johnson et al., 2002b;
Johnson and Meskhidze, 2013; Krishnamurthy et al., 2009;
Martin et al., 1994; Moore et al., 2004; Parekh et al., 2004;
Pérez-Guzmán et al., 2010) and advocates a balanced ap-
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36 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
Figure 7. Summary of the principal cooling effects of the proposed iron salt aerosol method.
proach to profit from the iron cycle to fight global warming
by enhancing natural processes.
Climate cooling by natural ISA involves the troposphere,
dry solid surfaces, ocean waters, ocean sediment, the ocean
crust and land. Several GHG factors are controlled by ISA:
CO2, CH4, tropospheric O3, black carbon, dust, cloud albedo
and vertical ocean mixing.
Using mineral dust as a natural analogue tool, this arti-
cle proposes to enhance the natural ISA in order to raise and
heighten the cooling impacts of at least two of the dozen nat-
ural effects found: i.e., CH4removal by tropospheric q
Cl and
CO2removal by soluble-Fe ocean fertilization.
The ISA method proposed is feasible, probably with few
to no environmental side effects, as it relates to chemical
and/or physical combustion processes occurring currently.
Actual iron production and coal combustion together with
other combustion sources already release a very significant
part of the global bioavailable iron in the northern oceans
into the atmosphere: from 15 % (Ito and Shi, 2016) to 80 %
(Lin et al., 2015; Wang et al., 2015b), depending on the iron
solubility parameters taken into account.
The present level of atmospheric deposition of soluble Fe
over the global ocean is considered to be about 3 times higher
than for 1850 emissions (Myriokefalitakis et al., 2015), as in-
creases in anthropogenic and biomass burning emissions re-
sulted in both enhanced Fe combustion emissions and a more
acidic environment and thus more than twice the soluble-Fe
deposition (nearly 0.5 Tg-Fe yr−1today versus nearly 0.2 Tg-
Fe yr−1in 1850).
The inevitable reduction in aerosol emissions to improve
air quality in the future might accelerate the decline in
oceanic productivity per unit warming and accelerate a de-
cline in oceanic net primary productivity (NPP; Wang et
al., 2015a). The Myriokefalitakis model-projected results
for 2100 indicate about a quarter decrease in the atmospheric
deposition of soluble Fe, with a 5-fold decrease in Fe emis-
sions from anthropogenic combustion sources (∼0.070 Tg-
Fe yr−1today against ∼0.013 Tg-Fe yr−1in 2100). These
changes are expected to impact most on the high-nutrient–
low-chlorophyll oceanic regions. According to Myriokefali-
takis et al. (2015), in view of the importance of Fe as a mi-
cronutrient for marine ecosystems, the calculated projected
changes in soluble-iron emissions require the implemen-
tation of comprehensive mineral-Fe dissolution processes
as well as Fe combustion emissions in coupled climate–
biogeochemistry models to account for feedbacks between
climate and biogeochemical cycles. This review shows that
the effects on CH4of ISA and of anthropogenic Fe emissions
in the troposphere also deserve to be taken into account.
According to Wang et al. (2015b), taking into considera-
tion the relatively high solubility of anthropogenic iron, com-
bustion source contributions to soluble-Fe supply for north-
ern Pacific and northern Atlantic oceanic ecosystems could
be amplified by 1–2 orders of magnitude. To stop global
warming, we estimated the requirements in terms of ISA by
Earth Syst. Dynam., 8, 1–54, 2017 www.earth-syst-dynam.net/8/1/2017/
F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control 37
Table 3. Principal effects of the ISA method proposed – or its natural equivalent – and their probable consequences for the different biosphere
compartments.
Compartment Locality Effect Most Time delay
and/or action probable between
cooling cooling onset
efficiency or offset after
ISA method
start or stop
Troposphere Boundary layer Cloud albedo increase + + + <1 yr
and lower Methane and VOC depletion + + + <1 yr
troposphere Black and brown carbon precipitation ++ <1 yr
Ozone depletion ++ <1 yr
Continent Forests and Organic C burial increase by +<5 yr
other primary assimilation increase
producers
Wetlands, Methane emission decrease + + + <5 yr
marshes, peat by methanogenesis inhibition
bogs, lake
sediments
Desert surfaces Methane and VOC depletion ±<1 yr
Ocean and ocean Phytoplankton Organic and carbonate C + + ++1<1 yr
sediment aquifer and other burial increase by +2<1 yr
at the ocean food chain links assimilation increase
bottom
Ocean crust Activation of the Carbonate C burial increase +++++3>10 yr
aquifer ocean basin in the ocean crust rock +/+ + +4>10 yr
vertical cycling
1The euxinic and alkaline bottom water of the stratified ocean have no oxidation and calcite solution capacity, thus producing a high burial rate of
organic sediment C and carbonate C. 2The oxic, hydrogen carbonate and CO2-containing bottom water of the well-mixed ocean have a high
oxidation capacity and high calcite dissolving capacity, thus producing a low burial rate of organic and inorganic sediment C. 3The high inorganic
C load of the oxic, hydrogen carbonate and CO2-containing bottom water of the well-mixed ocean reaches total precipitation within the alkaline
and reducing crust aquifer, thus producing a very high burial rate of inorganic C and small amounts of organic C precipitation. 4The euxinic and
alkaline bottom water of the stratified ocean has a low content of dissolved inorganic C and contains methane C up to saturation, thus producing a
low to medium C burial rate during cycling through the crust aquifer.
the extrapolation of the experiments on iron-catalyzed acti-
vation by artificial sea-salt aerosols (Wittmer et al., 2015a;
Wittmer and Zetzsch, 2016). Our first estimations show that
by doubling the current natural Fe emissions by ISA emis-
sions into the troposphere, i.e., by about 0.3 Tg Fe yr−1, ar-
tificial ISA would enable the prevention or even the reversal
of GW.
The adjustable flue gas temperatures for different types of
combustion are a means to lift the ISA plumes to optimal
heights within the troposphere. Thus, we believe that the ISA
method proposed integrates technical and economically fea-
sible tools that can help to stop GW.
In accordance with our remarks in Sect. 2, the reactions
of ISA in the troposphere provide the most prominent results
for a surface temperature decrease (Meyer-Oeste, 2010). This
stops further ice melting, which activates the different verti-
cal ocean water movements. As a result, the dissolved CO2
is then buried as carbonate C within the ocean bottom sedi-
ments and crust.
Author contributions. Franz Dietrich Oeste suggested the review
idea and performed an initial bibliographical search completed by
Renaud de Richter. Franz Dietrich Oeste and Renaud de Richter
prepared the paper and the figures with contributions from all coau-
thors. Tingzhen Ming and Sylvain Caillol also contributed to struc-
turing the paper, ideas, bibliographical entries and English correc-
tions.
Competing interests. The authors declare that they have no con-
flict of interest.
Acknowledgement. This research was supported by the Scien-
tific Research Foundation of Wuhan University of Technology
(No. 40120237) and the ESI Discipline Promotion Foundation of
WUT (No. 35400664).
The coauthors would like to thank both reviewers (S. M. Elliott
and an anonymous reviewer) for their constructive and thoughtful
www.earth-syst-dynam.net/8/1/2017/ Earth Syst. Dynam., 8, 1–54, 2017
38 F. D. Oeste et al.: Climate engineering by mimicking natural dust climate control
reviews, which greatly improved this paper, in particular Sects. 5,
9 and 10. We also thank Rolf Sander and Cornelius Zetzsch for
their constructive comments and Louise Phillips for grammatical
corrections and rereading. The authors would like to thank the
Copernicus production office for language editing and corrections,
which greatly improved this article.
Edited by: B. Kravitz
Reviewed by: S. M. Elliott and one anonymous referee
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