Introduction
Aeolian erosion of semi-arid and arid desert surfaces contributes to an
estimated ∼ 1000 to 3000 Tg yr-1 global emission of mineral dust
aerosol (Goudie and Middleton, 2001; Engelstaedter et al., 2006; Cakmur et
al., 2006). The impact of desert dust is large (Gieré and Querol,
2010): (i) it directly affects atmospheric radiative balance due to
scattering and absorption of solar and terrestrial radiation, thereby
affecting atmospheric dynamics and climate (Carlson and Benjamin, 1980; Tegen
and Lacis, 1996: Ramanathan et al., 2001; Tegen, 2003; Balkanski et al.,
2007). Indirectly, it also affects climate via its potential for altering
atmospheric physics/microphysics (i.e., acting as nuclei for liquid and solid
cloud droplets) and the hydrological cycle (Ramanathan et al., 2001). Such
direct and indirect forcings strongly depend on mineral dust physicochemical
properties, including particle size, shape and composition/mineralogy (Tegen,
2003; Lafon et al., 2006; Formenti et al., 2011; Mahowald et al., 2014; Zhang
et al., 2015; Kok et al., 2017); (ii) it supplies sediments to downwind
marine and continental areas, affecting the surface albedo of the latter,
whereas dust entrainment in source regions is a major erosive agent, strongly
affecting soil quality (Goudie and Middleton, 2001); (iii) it also supplies
key micronutrients (e.g., iron and phosphorous) to distal ocean and inland
water environments, directly affecting the C cycle via stimulated
bioproductivity, and indirectly affecting climate via atmospheric CO2
sequestration (Jickells et al., 2005; Raiswell and Canfield, 2012); (iv) it
is involved in a range of heterogeneous reactions with manifold implications
(Usher et al., 2003). For instance, carbonates in Saharan dust increase the
pH of precipitation acting as a buffer for acid rain in Europe
(Loÿe-Pilot et al., 1986); (v) desert dust storms are a hazard with
detrimental effects on transportation (e.g., reduced visibility),
infrastructure, and (solar) energy generation (Middleton, 2017), also causing
soiling and discoloration of monuments (Comite et al., 2017); (vi) desert
mineral dust is a health hazard to humans (Karanasiou et al., 2012; Goudie,
2014). Exposure to desert dust (particulate matter with size
<10 µm, PM10 and/or with size <2.5 µm,
PM2.5) has been associated with morbidity and premature death due
to dust-related (or enhanced) cardiovascular and respiratory problems (Perez
et al., 2008), as well as several diseases related to dust-borne
microorganisms (short-term effects) (Griffin, 2007). In addition,
silicosis/pulmonary fibrosis (desert lung) and cancer-related illnesses have
been associated with desert dust exposure (long-term effects) (Giannadaki et
al., 2014).
More than half of the worldwide mineral dust aerosol comes from northern
Africa (Sahara/Sahel) (Goudie and Middleton, 2001; Prospero et al., 2002;
Engelstaedter et al., 2006), with an estimated ∼ 700–1600 Tg of
Saharan dust exported yearly across the Mediterranean Sea to Europe and the
Near East, the Red Sea to the Near East and Asia, and the North Atlantic
Ocean to the Americas (D'Almeida, 1986; Prospero, 1996; Goudie and
Middleton, 2001). Although most Saharan dust is transported across the
Atlantic Ocean (Carlson and Prospero, 1972), an estimated
80–120 Tg yr-1 is transported northward across the Mediterranean Sea
to Europe (D'Almeida, 1986). It has been pointed out that the strength of the
Saharan dust input increased since ca. the mid-20th century due to recurrent
droughts in northern Africa (Prospero and Lamb, 2003), anthropogenic-induced
desertification (Moulin and Chiapello, 2006), changes in land use, including
an increase in cultivable lands in the Sahel region (Mulitza et al., 2010),
and other larger-scale phenomena (e.g., atmospheric circulation patterns
and/or climate change) (Sala et al., 1996; Moulin et al., 1997). Nonetheless,
there is evidence of significant seasonal to decadal variability of Saharan
dust strength. Indeed, since the end of the 1980s, a trend towards decreasing
Saharan dust across the tropical North Atlantic has been reported (Ridley et
al., 2014; Evan et al., 2016), while the strength of Saharan dust affecting
southern Europe has reportedly increased in recent decades (Antoine and
Nobileau, 2006). However, for the 2001–2013 period, a decreasing strength in
Saharan dust affecting the western Mediterranean has been reported (Pey et
al., 2013; Vincent et al., 2016), which in one case (2001–2011 period) has
been related to negative summer NOA (North Atlantic Oscillation) (Pey et al.,
2013). Despite this temporal variability, dust plume intrusions in the
Mediterranean area are currently rather common (Escudero et al., 2005; Avila
et al., 2007; Cabello et al., 2012; Titos et al., 2017) and most of the
temporal variability in desert dust deposition appears to be due to very
intense but rare events (Vincent et al., 2016). Desert dust plumes lead to
both dry and wet deposition of mineral dust (Escudero et al., 2005). Wet
deposition typically occurs as “red rain”, “dust rain”, “blood rain”,
or “muddy rain” events that periodically and persistently affect southern
Europe (Prodi and Fea, 1979), and most particularly the Iberian Peninsula
(Sala et al., 1996; Avila et al., 1997; White et al., 2012), the European
region closest to northern Africa. Although known since ancient times
(Gieré and Querol, 2010), red rain events have experienced a remarkable
increase in their frequency and intensity over the last decades (Sala et al.,
1996; Escudero et al., 2005; Fiol et al., 2005; Avila et al., 2007). In some
cases they are extreme, with 10–40 g m-2 of dust deposited after a
single red rain event (Avila et al., 1997, 2007; Fiol et al., 2005),
dwarfing the average yearly Saharan dust deposition in southwestern Europe,
estimated to be about 3–14 g m-2 (Goudie and Middleton, 2001). This
was the case of the last extreme red rain event that took place in the area
of Granada (south of Spain) on 21–23 February 2017.
The global significance and impact of desert-derived mineral dust aerosol has
attracted extensive research focused on analyzing dust composition,
mineralogy, physical properties, sources, and
entrainment–transport–deposition mechanisms and patterns (see reviews by
Goudie and Middleton, 2001; Prospero et al., 2002; Scheuvens et al., 2013).
Data from these studies have contributed to a better understanding and
modeling of the dust cycle, as well as its impact on global atmospheric
dynamics, climate, and biogeochemical cycles. However, there are several
aspects of the current knowledge that are far from complete. This is the case
of the physicochemical and mineralogical features of Saharan dust, which are
highly variable and event-specific and, therefore, poorly constrained.
Although it has been acknowledged that the mineralogy and the physicochemical
features of desert dust particles are relevant factors to be considered in
modeling dust and climate, they are in general poorly constrained (Formenti
et al., 2011; Titos et al., 2017) and have been largely ignored in most
models (Krueger et al., 2004). This is likely due to the limited number of
studies dedicated to the detailed characterization of desert dust mineralogy
and physicochemical features, particularly those of individual particles
(Jeong et al., 2016).
Using a multianalytical approach, here we studied Saharan dust samples
collected immediately after the extreme red rain event that took place in
Granada (37.17806∘ N, 3.60083∘ W; ∼ 680 m a.s.l.)
on 21–23 February 2017. One specific objective of this study was to analyze
in detail the content, physical and textural properties, mineralogy, and
composition – including individual particles – of the clay fraction
(φ<2 µm). This is typically a major fraction in desert
dust, is considered responsible for most of the scattering of sun light,
includes most of the bioavailable iron, and due to its long atmospheric
residence time is the one that can affect most distant locations (Tegen and
Lacis, 1996; Sokolik and Toon, 1999; Lafon et al., 2006; Journet et al.,
2008; Formenti et al., 2014a, b; Jeong and Achterberg, 2014; Jeong et al.,
2016). Another goal of our study was to compare the mineralogy and
physicochemical properties of the clay fraction with those of the two other
relevant size fractions, i.e., sand and silt, an aspect that has been
generally ignored in previous studies. It should be noted that such an
analysis is complex and tedious, and requires a relatively large amount of
sample. The studied red rain event thus offered a unique opportunity to
collect a sufficient amount of material from a single Saharan dust event so
as to perform a detailed multianalytical study of wet-deposited desert dust.
In addition, the analysis of the synoptic-scale meteorological conditions
during this event, and the identification of transport routes and potential
dust source areas (compatible with the results of the mineralogical
analysis), were performed using satellite imagery, synoptic reanalysis, dust
forecast, and air mass backward/forward trajectory modeling. Ultimately, we
strived to shed light on the potential biogeochemical, radiative, and health
effects that such extreme Saharan dust events can have locally as well as
globally.
Results and discussion
The extreme winter Saharan dust event
On 20 and 21 February, a large Saharan dust plume coming from northern Africa
(off the coast of Algeria and Morocco) crossed the Alboran Sea (western
Mediterranean) and wiped the southern portion of the Iberian Peninsula in a
counterclockwise N-NW motion (Fig. 1a). This initial dust intrusion, which
was clearly visible in Suomi VIIRS as well as in Terra and Aqua MODIS true
color imagery on 21 February (on 20 February the southern portion of the
Iberian Peninsula was covered by clouds, thereby precluding the direct
observation of the dust plume), caused a minor red rain event in the Granada
area during the night of 21 February. Subsequently, a more massive African
dust plume penetrated in the southern part of the Iberian Peninsula during
22 February (Fig. 1a), leading to significant wet deposition in Granada (and
in almost all areas in the southern portion of the Iberian Peninsula) during
the night of 22 February and the early hours of 23 February. The volume of
precipitation was minor (∼ 1–2 mm), but the amount of dust deposited
was massive: we measured an average of 18 ± 8 g m-2 deposited in
the city of Granada. Note that the relatively high scattering in the amount
of deposited dust was due to orientation differences among the different
collection sites (i.e., shielding effects of nearby buildings and/or
vegetation). Streets, houses, and cars across the city appeared covered by
reddish-brown mud and in the nearby Sierra Nevada the snow cap displayed a
dramatic reddish-brown discoloration (compare the before and after satellite
images of the Sierra Nevada shown in Fig. 1b, c). Such an amount of deposited
Saharan dust is about 4 times higher than the average input of
5.1 g m-2 yr-1 measured by Avila et al. (1997) in northeastern
Spain for the period 1981–1992. Remarkably, Avila et al. (1997) reported
that just two red rain events accounted for nearly 62 % of the total
desert dust mass deposited in the area over 11 years: the extreme events in
November 1984 and in March 1991 delivered 16.4 and 19.4 g m-2,
respectively. These values are in very good agreement with the value reported
here.
During the last hours of 22 February and the early hours of 23 February, a
maximum ground-level PM10 concentration of
275 µg m-3 was measured in the Granada urban area (Fig. 1d),
a value more than 5 times higher than the daily limit
(50 µg m-3) established by Directive 2008/50/EU of the
European Union. Surface PM10 values rapidly decreased to values
∼ 50–75 µg m-3 following the red rain event. Note that
Saharan dust events with daily
PM10 > 100 µg m-3 are considered “extreme”
and are infrequent (1 % of all dust events) in the western Mediterranean,
but more frequent (2–5 %) in the eastern Mediterranean (Pey et al.,
2013). Nonetheless, for the case of the Iberian Peninsula, more than nine red
rain events with a dust deposition >1 g m-2 per event (i.e., extreme
events) were reported for the period 1984–1993 (Avila et al., 1997), and
four more between 1996 and 2002 (Avila et al., 2007).
NMMB-BSC dust forecast aerosol optical depth (AOD) at 550 nm for
the days before, during, and after the extreme red rain event. Images from
the Barcelona Dust Forecast Center.
NMMB-BDF modeling showed massive dust plumes initially coming from southern
Algeria and northern Mali (and northwestern Niger), with subsequent
contributions from northern Algeria, southern Tunisia, and western Libya,
entering the Iberian Peninsula on 20 February, and wiping the southern
portion of Spain in two successive waves. The second wave was more intense
than the first, with an aerosol optical depth, AOD, at 550 nm of 1.6–3.2.
Finally, the dust plume shifted course towards the central Mediterranean on
24 February (Fig. 2). It should be noted that the accuracy of the NMMB-BDF
forecasted dust outbreak and its spatio-temporal evolution is remarkable, as
demonstrated by the satellite observations (cf. Figs. 1a and 2) and the
recent lidar study of this extreme Saharan dust event by Fernández et
al. (2018).
Overall, the studied red rain event stands out as one of the most extreme of
the last decades. Modeling and satellite imagery demonstrates that the event
was associated with a massive desert dust outbreak affecting a large portion
of the central- and north-western Sahara. Entrained Saharan dust was rapidly
transported to the Iberian Peninsula where massive wet deposition of desert
dust took place.
Synoptic-scale meteorological situation during the red rain event
based on NOAA/ESRL reanalysis. Averaged geopotential high (a, c) and
wind field (b, d) for 20–23 February 2017 at both ground level
(1000 hPa) and 850 hPa (∼ 1500 m a.s.l.) are displayed.
Synoptic situation and backward/forward trajectories
The synoptic-scale meteorological situation during the extreme dust event was
characterized by a marked northwestern African depression, nearly centered on
the leeside of the Atlas Mountains (eastern Morocco–western Algeria). Such a
depression was bounded by two anticyclones, one centered in the North
Atlantic (Azores area) and another centered in Libya (Fig. 3 and Fig. S1 in Supplement
Sect.). Air masses were advected by strong winds with westward direction in
the area of northwestern Niger and northern Mali that turned north upon
entering the southern border of Algeria, heading straight north/northeast
towards the southeastern part of the Iberian Peninsula (Fig. 3b, d). Advected
air masses turned counterclockwise upon reaching the mid portion of the
Iberian Peninsula and then moved west into the Atlantic Ocean off the coast
of Portugal, and then to the south, following the cyclonic wind field.
Supplement Fig. S1 shows the temporal (daily) and spatial evolution of the
cyclone based on NOAA/ESRL reanalysis of geopotential height and wind field
at 850 hPa. Remarkably, the center of the depression remained stationary
over the same area (northwestern Africa) for nearly 4 days (19–22 February
2017).
Escudero et al. (2005) reported that major Saharan dust episodes affecting
the Iberian Peninsula are associated with (a) a northern African
high-pressure system (located at surface or upper levels), (b) an Atlantic
depression, or (c) a northern African depression. The latter synoptic
scenario is associated with dust outbreaks which in 87 % of the cases led
to red rain. These red rain events tend to occur during winter–spring and
autumn, with dust typically coming from Algeria (mostly from northern areas,
although dust in some events reportedly came from central Algeria) (Avila et
al., 2007). Conversely, in summer, dry deposition prevails, associated with a
high thermal anticyclone system in northern Africa (Rodriguez et al., 2001;
Escudero et al., 2005). The amount of dust deposited per event is smaller in
this latter scenario. In our case, a strong northwestern Saharan depression
was responsible for the Saharan dust entrainment and transport to the Iberian
Peninsula. The synoptic scenario has strong similarities to a Sharav cyclone
(Bou Karam et al., 2010). Cyclogenesis was triggered by an upper level (i.e.,
300 hPa) N–S trough west of the Iberian Peninsula which favored the
injection of strong cold northerly winds to the warmer northern African
troposphere (Fig. S2), thereby favoring the conditions for baroclinic
instability. A very similar overall synoptic situation has been previously
reported for other major red rain events taking place in the Spanish
Mediterranean area (Sala et al., 1996), inland within the southern and
northeastern parts of the Iberian Peninsula (Avila et al., 1997, 2007;
Rodriguez et al., 2001), and in the western–central Mediterranean (Fiol et
al., 2005). Remarkably, two recent major winter dust events affecting the
western Mediterranean and associated with northern African cyclones took
place on 20–23 February 2007 (Bou Karam et al., 2010) and on 20–25 February
2016 (Titos et al., 2017). These observations suggest that there is a
recurrent pattern in the spatio-temporal synoptic conditions (e.g., northern
African cyclones) leading to such extreme (winter) Saharan dust outbreaks and
associated red rain events.
Results of NOAA's HYSPLIT modeling of (a) 72 h multiple
(25) back-trajectories for air masses arriving at Granada on 23 February 2017
(00:00 UTC) and (b–f) forward multiple (25) trajectories for
selected dust source areas. Shaded areas in (a) show potential
source areas PSA1 and PSA3.
The backward and forward trajectory analyses showed that the air masses
arriving at Granada during the studied event entrained dust over a broad area
spanning from central and southern Algeria, the northern part of Mali, and
the northwestern part of Niger to the north of Algeria, south of Tunisia, and
western Libya (Fig. 4). This broad area includes two well-known Saharan dust
potential source areas (PSAs) as defined by Formenti et al. (2011) and
Scheuvens et al. (2013): (i) PSA1, which covers the zone of chotts (ephemeral
lakes) and dry lakebeds south of the Tell Atlas in northeastern Algeria,
southern Tunisia, and northwestern Libya. This area is characterized by a
relatively high illite, palygorskite, and carbonate content, and (ii) PSA3,
which is one of the largest, most persistent, and most intense dust source
areas in the Sahara (Goudie and Middleton, 2001; Prospero et al., 2002). It
is located in the central–southern section of Algeria, and spreads along the
basins located southwest of the Ahaggar massif, the northern part of Mali,
and the frontier with Niger (e.g., Adrar des Iforas). Dust entrained from
this area is characterized by relatively high illite, kaolinite, and smectite
contents, with minor amounts of palygorskite (Scheuvens et al., 2013). The
importance of these two major source areas, particularly the second one, has
been recognized by several authors (e.g., D'Almeida, 1996; Goudie and
Middleton, 2001; Prospero et al., 2002). However, previous studies did not
recognize the contribution of such a distant area as PSA3 to extreme dust
events affecting western Europe. Avila et al. (1997) reported that three main
source areas were active during red rain events affecting the Iberian
Peninsula: (i) western Sahara (between 26–30∘ N and
14–8∘ W), (ii) Moroccan Atlas (between 30–35∘ N and
8–0∘ W), and (iii) central Algeria (between 26–30∘ N and
4∘ W–5∘ E). Avila et al. (1997) report that the latter
source area is activated by a depression over the Iberian Peninsula or over
northern Africa. Such a source area is the one most closely matching the
potential source areas identified here. However, our back- and
forward-trajectory analysis shows that dust was mobilized from Saharan
regions located much further south as well as from northern areas. We also
considered the possibility that dust reaching the Iberian Peninsula could be
entrained from PSA5, that is, the Bodélé depression (Formenti et al.,
2011; Scheuvens et al., 2013), which is considered the single most active
dust source area in the world (Goudie and Middleton, 2002; Prospero et al.,
2002). However, our forward-trajectory analysis (Fig. 4f) demonstrated that
no dust entrained from this source area during the days preceding the studied
dust event reached the Iberian Peninsula.
In summary, the synoptic scenario leading to the studied extreme red rain
event appears to be recurrent, typically occurring in late winter or early
spring, and being associated with a northern African depression. Remarkably,
this situation can mobilize and wet-deposit huge amounts of Saharan dust
entrained from distant southern areas in the Sahara, about 3000 km away from
the Iberian Peninsula, as well as from closer areas in northern Africa.
Semiquantitative XRD analysis of Saharan dust powder samples
(wt %; ±1σ).
Sample
Qtz
Kfd
Plg
Cal
Dol
Goe
Hem
Rut
Clay minerals
BulkRIRa
23 ± 1
6 ± 1
7 ± 1
12 ± 2
2 ± 1
1.7 ± 0.8
0.5 ± 0.4
0.8 ± 0.3
47 ± 3b
BulkRietveldc
24 ± 1
5 ± 1
-
24 ± 1
8 ± 1
-
-
-
44 ± 2
Sandd
73 ± 5
5 ± 2
3 ± 1
-
-
1 ± 1
-
-
18 ± 3b
Siltd
51 ± 3
5 ± 2
4 ± 2
-
-
1 ± 1
1 ± 1
-
38 ± 4b
Clayd
7 ± 2
-
-
-
-
3 ± 2
2 ± 1
-
88 ± 5b
a Values obtained using the RIR method.
b Values determined considering the intensity of the general Bragg
reflection of phyllosilicates at 4.49 Å. c Values obtained
using the Rietveld method. d This size-fraction powder sample was
subjected to carbonate elimination. Legend: Qtz: quartz; Kfd: microcline;
Plg: plagioclase; Cal: calcite; Dol: dolomite; Goe: goethite; Hem: hematite;
Rut: rutile.
XRD analysis of Saharan dust: (a) XRD pattern of the bulk
sample showing the presence of smectites (Sm), palygorskite (Pal), illite
(Il), kaolinite (Kao), goethite (Goe), quartz (Qtz), K-feldspar (Kfd),
plagioclase (Plg), calcite (Cal), dolomite (Dol), rutile (Rut), and hematite
(Hem). The general reflection of clay minerals (Clays) at 4.49 Å is also indicated. Vertical color lines show the
peak position of the different phases (only the main peaks corresponding to
each phase are labeled). (b) XRD patterns of the sand, silt, and
clay fractions (decarbonated). (c, d) XRD patterns of oriented
aggregates of the clay and silt fractions, respectively (with different
treatments). The hkl index and d-spacing (in Å) of the clay
minerals' main reflections are indicated in
(c, d). Chl: chlorite; MLC: mixed-layer clays.
XRD analysis: linking mineralogy with potential dust source
areas
The analysis of the bulk dust deposit (powder samples) showed the presence of
(in order of decreasing abundance) (Table 1) clay minerals (see below for
details on individual clay minerals and their content), quartz
(SiO2), calcite (CaCO3), plagioclase (albite,
NaAlSi3O4), K-feldspar (microcline, KAlSi3O4), dolomite
(CaMg(CO3)2), goethite (α-FeOOH), hematite (α-Fe2O3), and rutile (TiO2) (Fig. 5a). Note that phase amounts
determined by XRD (RIR or Rietveld methods) are not considered to be purely
quantitative, but rather semiquantitative due to the errors associated with
this technique (typically ± 5 wt %) (Formenti et al., 2011). Note
also that amorphous phases (e.g., amorphous silica – in diatoms – and
amorphous iron oxyhydroxides; see TEM results below) cannot be detected using
XRD, and their content was not negligible, as shown by the broad hump at
18–32∘2θ in the XRD pattern (Fig. 5a). We observed that the
Rietveld method did not allow the quantification of oxides, plagioclase, and
clays other than illite and kaolinite (Fig. S3). It also grossly
overestimated the amounts of carbonate phases, i.e., 32 ± 1 wt %,
a value more than twice the amount determined using the RIR method
(14 ± 1 wt %), or by measuring the mass differences after acid
elimination of carbonates (13 ± 1 wt %), or by TG analysis
(14.7 ± 0.2 wt %) (see below). The use of the Rietveld method for
quantifying such complex multiphase dust samples is therefore not advised. In
contrast, the RIR method yielded more reliable results, especially in the
case of carbonate phases and clay minerals.
The detected mineral phases and contents are in general typical for Saharan
dust deposited all across Europe, during both dry- and wet-deposition events
(Scheuvens et al., 2013), as well as those reported for African dust
deposited on different areas across the tropical North Atlantic (Glaccum and
Prospero, 1980; Menéndez et al., 2014; Patey et al., 2015). This is not
unexpected because these phases are the most common in northwestern African
arid and semi-arid soils (Scheuvens et al., 2013; Journet et al., 2014).
Avila et al. (1997) reported a very similar mineral composition for the case
of Saharan dust deposited after red rain events on the northeastern part of
the Iberian Peninsula. However, the authors did not report the presence of
goethite, hematite, and rutile.
The total amount of clays measured here (∼ 47 wt %) is smaller
than the average value reported by Avila et al. (1997)
(∼ 64 wt %), or the values of 60–90 wt % reported for
Saharan dust collected in Africa or after long-range transport (Formenti et
al., 2014a), although it approaches the average value of 52 wt %, and is
well within the range 34–66 wt %, reported by Patey et al. (2015) for
northern African dust deposited in the tropical North Atlantic (Cape Verde).
This is likely due to the fact that very intense Saharan dust events, such as
the one studied here, typically mobilize larger particles with a relatively
lower clay content of ∼ 50 wt % (Caquineau et al., 2002; Formenti
et al., 2014a).
The XRD pattern of the sand fraction (after elimination of carbonates) showed
intense reflections corresponding to quartz (main phase), with minor
reflections corresponding to nearly all other identified phases (Fig. 5b and
Table 1). It was unexpected to find clays and iron oxyhydroxides within this
fraction, especially considering that acid dissolution of the carbonates and
thorough washing were performed prior to size separation. Apparently, clays
and iron oxyhydroxides are intimately associated (cemented) with the coarser
quartz and feldspar grains (see SEM and TEM results below). Abundant quartz
along with small amounts of feldspars and abundant clay minerals were
detected in the silt fraction (Fig. 5b and Table 1). Conversely, very small
amounts of quartz were detected in the clay fraction (Fig. 5b), where clay
minerals were the most abundant phases, along with iron oxyhydroxides.
Figure 5c shows the XRD patterns of oriented aggregates of the clay fraction.
An intense Bragg peak at 10.4 Å which did not change position following
EG and DMSO solvation, and which collapsed to 10 Å upon heat treatment
at 550 ∘C, was observed. These are standard features of the 110
Bragg peak of palygorskite ((Mg,Al)2Si4O10(OH) 4(H2O))
(Moore and Reynolds, 1989). Note that in some publications this reflection
has been mistakenly assigned to illite
(K(Al,Mg,Fe)2(Si,Al)4O10(OH)2) (e.g., Prodi and Fea,
1979). In addition, we also observed the d200 spacing of palygorskite at
6.4 Å, unambiguously confirming its presence. The intense peak at
7.17 Å, which did not shift position upon EG and DMSO solvation, and
which disappeared upon heat treatment, corresponded to the 001 reflection of
kaolinite (Al2Si2O5(OH)4). There was a poorly defined
broad band at 12–18 Å that shifted position towards higher d-spacings
upon EG and DMSO treatments, and collapsed to 10 Å upon heat treatment.
These features are indicative of swelling clays (smectites), most likely
montmorillonite ((Na, Ca)0.3(Al, Mg,
Fe)2Si4O10(OH)2 n(H2O)) (see TEM–AEM results
below), and mixed-layer clays (MLC, interstratified illite–smectite; see
TEM–AEM results below), accounting for the broad reflection centered at
d-spacings ∼ 14–17 Å and ∼ 12 Å, respectively (AD
sample). Note, however, that the broadness of the Bragg peaks precluded an
unambiguous identification of any individual swelling clay. Finally, a
shoulder located on the right of the 110 Bragg peak of palygorskite was
observed at 10 Å, best seen in EG-solvated samples, pointing to the
presence of illite (001 reflection), confirmed by the appearance of its 002
and 003 order reflections at 5 Å and 3.33 Å, respectively.
Some mineralogical changes were detected in the silt fraction as compared to
the clay fraction (Fig. 5d). The 001 reflection of illite was more intense
than the 110 reflection of palygorskite, pointing to a relative increase in
the abundance of illite in this coarser fraction. Kaolinite expanded upon
DMSO treatment as shown by the shifting of the 001 peak at 7.14 to
11.18 Å. This is a standard feature of this clay mineral which enables
its distinction from chlorite, which also has a strong reflection at
∼ 7.1 Å (González-García and Sánchez-Camazano, 1968).
However, the formation of a DMSO–kaolinite intercalation complex is
dependent on kaolinite crystallinity, as shown here by the fact that
kaolinite in the clay fraction, which displayed a very broad 001 reflection
(due to poor crystallinity), did not expand. These observations suggest size
fractionation of kaolinite particles from two different sources, with the
higher crystallinity coarser phase in the silt fraction likely coming from a
closer-to-Europe source area (i.e., PSA1), and the less crystalline kaolinite
in the clay fraction likely coming from a more distant source (i.e., PSA3).
Alternatively, it could be argued that the two types of kaolinite come from
an area where mixing of dust from different source areas had occurred (Guieu
et al., 2002). Finally, it was observed that lattice expansion of swelling
clays (smectite and MLC) upon EG and DMSO treatment unveiled the presence of
a Bragg peak at 14 Å which upon heat treatment remained unaltered. The
corresponding second-order 002 reflection at 7.1 Å was also visible
after heat treatment. This behavior is characteristic of chlorite
((Mg,Fe++)5Al(Si3Al)O10(OH)8), whose presence was
masked by the broad 001 reflection of swelling clays at ∼ 14 Å and
the 001 reflection of kaolinite at 7.1 Å in the oriented AD mount or in
the bulk powder sample.
Semiquantitative analysis (RIR method) of clay minerals (wt %).
Values in parentheses show the wt % content of each type of clay mineral
in the bulk sample. Note that the contents of individual clay minerals
determined from XRD analysis of the bulk (powder) sample are
subjected to a very high uncertainty. More reliable semiquantitative results
are obtained from oriented aggregates (OA).
Sample
Ill
Pal
Sm
Kao
MLC
Chl
Bulk (powder)
47 ± 2 (22)
24 ± 3 (11)
15 ± 4 (7)a
14 ± 1 (7)b
-
-
Clay fraction (OA)
29 ± 1 (4.2)
43 ± 2 (6.3)
8 ± 1 (1.2)
14 ± 1 (2)
7 ± 2 (1)
-
Silt fraction (OA)
32 ± 1 (10.3)
33 ± 3 (10.7)
11 ± 1 (3.5)
19 ± 2 (6.2)
-
6 ± 2 (1.9)
Total silt + clay fractions (OA)
(14.5)
(17)
(4.7)
(8.2)
(1)
(1.9)
a Includes the contribution of MLC and Chl.
b Includes the contribution of Chl. Legend: Ill: illite;
Pal: Palygorskite; Sm: smectite; Kao: kaolinite; MLC: mixed-layer clay;
Chl: chlorite.
The results of the semiquantitative XRD analysis of the clay minerals in the
bulk fraction (powder samples) as well as in the clay and silt fractions
(oriented mounts) are presented in Table 2. Apparently, the most abundant
clay mineral in the bulk sample was illite. However, the analysis of oriented
mounts showed that the most abundant clay was palygorskite, illite being the
second most abundant clay mineral. This discrepancy can be explained
considering that the RIR (or the Rietveld) method for (ideally non-oriented)
illite fails to fully detract from the contribution of the main reflection of
(abundant) quartz at 3.33 Å, which overlaps the 003 main reflection of
illite. To this effect, one has to add the intensity increase related to
preferential orientation of this phyllosilicate in the bulk powder sample. As
a result, an overestimation of illite takes place. More reliable results are
thus obtained using oriented mounts and in-house experimental RIR values
referred to as the 001 reflections (or 110, in the case of palygorskite) of
the clay minerals. The amount of kaolinite was relatively high in the silt
and clay fractions, but still much lower than that of illite. The amount of
chlorite, only present in the silt fraction, was low (∼ 5 wt %),
whereas the amounts of smectite and MLC were relatively high. Note that MLC
were only identified in the clay fraction. Finally, these results show that,
contrary to common assumptions regarding the lack of abundant clay minerals
in the silt fraction (Journet et al., 2014), these phases are very abundant
in this coarser fraction. Note also that because the mass of the silt
fraction was 2.06 times that of the clay fraction (as revealed by the results
of the hydrodynamic size separation showing sand = 2 wt %,
silt = 68 wt %, and clay = 30 wt %), the amount of
smectites, which along with MLC typically concentrate in the clay fraction,
was relatively high in the silt fraction. Nonetheless, if we consider the
unit mass of silt and clay fractions, the amount of smectite in the silt
fraction (11 %) is lower than the amount of smectite plus MLC in the clay
fraction (15 %).
Most of the phases detected by XRD, with the exception of carbonates and some
of the clays, cannot be used as a reliable indicator for dust source
identification because they are common to all northern African dust source
areas (Scheuvens et al., 2013; Formenti et al., 2014a; Journet et al., 2014).
Calcite and dolomite are recognized as good indicators of a northern and
northwestern Saharan origin (Avila et al., 1997; Scheuvens et al., 2013),
although southern Algeria and the Mali–Algeria border are areas that also
contribute significant amounts of carbonates to Saharan dust (Scheuvens et
al., 2013). Interestingly, the TEM analysis (see below) showed the presence
of abundant fibrous calcite, which typically forms in playa and/or ephemeral
(and paleo) lakes in arid regions (Wanas, 2012). This points to PSA1 (chotts
in northern Algeria–Tunisia) as the most likely source area for this
mineral.
The illite / kaolinite (I / K) ratio has been used to identify Saharan
dust source regions (Caquineau et al., 2002). I / K >2 has been
associated with northwestern Saharan source areas, I / K <0.5 with
southern Saharan and Sahelian source areas, and intermediate values with
central Saharan source areas. The I / K ratio in our dust samples ranged
between 1.68 (silt fraction) and 2.07 (clay fraction) (we ignore here the
I / K ratio of the bulk powder sample, because the illite content is
overestimated for the reasons discussed above). Considering the wt % of
each size fraction (see above), the overall I / K ratio of the clay plus
silt fractions is 1.77. This points to a central Saharan source area (i.e.,
southern–central Algeria). However, and in agreement with satellite imagery,
BSA dust forecast, and backward-/forward-trajectory analyses, the existence
of two types of kaolinite with low (no expansion with DMSO) and high
crystallinity (expansion with DMSO) suggests the possibility of mixing of
dust entrained from both southern (low I / K ratio) and northern (high
I / K ratio) Saharan areas (i.e., PSA3 and PSA1). Indeed, the relatively
high carbonate content and abundant palygorskite point to an additional
north(western) Saharan source. Palygorskite has been recognized as a reliable
indicator of a Saharan provenance (Coudé-Gaussen, 1991; Scheuvens et al.,
2013), and has been detected in red rain events not only in the Iberian
Peninsula (Avila et al., 1997), but also in the western (Fiol et al., 2005),
central (Molinaroli, 1996), and eastern Mediterranean areas (Ganor et al.,
2009), the Alps (De Angelis and Gaudichet, 1991), the Netherlands (Reiff et
al., 1986), and the British Isles (Bain and Tait, 1977). Palygorskite is a
common clay in soils of the northwestern Sahara (PSA2), as well as northern
Algeria and Tunisia (PSA1) (Scheuvens et al., 2013). Nonetheless, occurrences
of palygorskite, which typically forms in saline lakes and alluvial sediments
of arid regions (Singer and Galan, 1984), have also been documented in
central and southern Algeria, as well as in northern Mali (PSA3)
(Coudé-Gaussen, 1991; Scheuvens et al., 2013). Remarkably, the amount of
palygorskite in our studied dust samples is exceptionally high (17 wt %
of the bulk sample; see Table 2). Previous studies of Saharan dust deposited
in the central Mediterranean (Mallorca and Sardinia) reported palygorskite
values ≤5 wt % (Molinaroli, 1996; Fiol et al., 2005), although Avila
et al. (1997) observed a concentration of up to 12 wt % in red rains in
northwestern Spain. It is likely that in our case, the close proximity to
northern Africa and the intensity of the dust event favored the entrainment
of palygorskite-rich dust from the northern Algeria and Tunisia areas, and
its transport with minimal segregation of the silt-sized palygorskite
particles prior to their deposition in southern Spain. Alternatively, it
might be argued that in previous studies the amount of palygorskite was
underestimated due to the difficulty in quantifying this mineral using XRD.
Such a problem is rooted in the general lack of reliable and accurate RIR
values for palygorskite: this is why we experimentally determined the RIR
value of palygorskite. The same applies for quantitative Rietveld analysis of
palygorskite due to the difficulty in obtaining accurate structural factors
for clay minerals in general, and for this clay mineral in particular. We
will show below that appropriate characterization and quantification of this
mineral in Saharan dust are of great significance due to its potential health
effects. Finally, it is worth commenting on the chlorite to kaolinite ratio
(Chl / K), which has also been proposed for discriminating Saharan dust
source regions (Scheuvens et al., 2013). Values <0.3 have been reported for
southern Algeria and northern Mali, while higher values have been observed in
soils from the northern and northwestern Sahara (see the compilation by
Scheuvens et al., 2013). We obtained Chl / K ∼ 0.2, a value which
is consistent with a southern Algerian and/or northern Mali source region.
Overall, and in agreement with previous studies that pointed out that major
Saharan dust outbreaks affecting large desert areas typically involve mixing
of dust entrained from different active dust source areas (Skonieczny et al.,
2011; Formenti et al., 2014a), our mineralogical analysis, in combination
with satellite imagery, NMMB-BSC dust forecast, and forward- and
backward-trajectory analyses, shows that this is the case here too, with PSA1
and PSA3 as the most probable dust source areas, although we cannot rule out
the possibility that dust could originate from a larger zone than PS1 and
PS3. Despite the difficulty in accurately pin-pointing the different source
areas down to a local scale during such an extreme event, our results show
that it is still possible to identify different regional dust signatures that
help constrain potential dust source areas.
Chemical composition of Saharan dust deposited in Granada during an
extreme red rain event.
Major/minor elements (wt %)
Si
Al
Fe
Mg
Ca
Na
K
Ti
P
XRF
24.05
6.62
3.69
1.24
5.72
1.66
1.53
0.45
0.04
ICP
nd
6.44
3.43
1.48
6.29
1.62
1.57
nd
0.06
Minor/trace elements (ppm)
Ba
Sr
Mn
Zn
V
Cr
Co
Ni
Pb
S
XRF
nd
162
400
nd
nd
78
nd
51
nd
nd
ICP
529
231
741
367
103
84
16
44
27
390
nd: not determined.
Chemical composition
Table 3 shows the results of XRF and ICP analysis of the dust samples
(chemical composition of bulk dust). Both techniques yielded consistent and
comparable results. Si and Al were the two most abundant elements, which
indicates that quartz and aluminosilicates (clays and feldspars) are the main
mineral phases in Saharan dust (Goudie and Middleton, 2001). The Si / Al
ratio of 3.68 (note: all ratios calculated using the average values of XRF
and ICP results) falls within the range of values reported for Saharan dust
source areas and Saharan dust deposited in western and central Mediterranean
areas (Scheuvens et al., 2013). It is in good agreement with the value (3.92)
reported for soils in southern Algeria (Guieu and Thomas, 1996). The latter
is consistent with the PSA3 source area, as discussed above. This value is,
however, larger than the Si / Al values (1.6–3.0) reported by Formenti
et al. (2014a) for Saharan dust and soils. Likely, the higher Si content in
our samples is due to the favored entrainment and transport of abundant
coarse SiO2 particles during such an extreme event. Although a
Si / Al ratio of 3.68 is also consistent with a PS5 source area (i.e.,
the Bodélé depression) (Formenti et al., 2014a), as stated above,
such a source area is ruled out here because the forward-trajectory analysis
(Fig. 4f) showed that no dust entrained from this area reached the Iberian
Peninsula. The relatively high Ca, and to a lesser extent Mg, content shows
that carbonates (calcite and dolomite) are abundant. Note, however, that such
elements are also present as octahedral or interlayer cations in clay
minerals, especially smectites and MLC (Mg and Ca), palygorskite (Mg), and
chlorite (Mg) (see AEM results below). The Ca / Al and Mg / Al ratios
of 0.21 and 0.92, respectively, agree with those reported for Saharan dust
from Algerian source areas deposited in southern Europe (Avila et al., 2007;
Scheuvens et al., 2013). Fe is the fourth most abundant element, with an
average concentration of 3.56 ± 0.18 wt %, being present both in
phyllosilicates (see TEM–AEM results below) and in iron oxyhydroxides. This
value matches that of the average Fe content in the continental crust
(3.5 wt %) (Taylor and McLennan, 1985) and falls within the range
∼ 2–11 wt % reported for Saharan dust and soils (Zhang et al.,
2015), but is slightly lower than the average value of 4.45 wt % Fe
proposed as characteristic of Saharan dust (Guieu et al., 2002), or the range
of values (4.3–6.1 wt % Fe) reported by Lafon et al. (2006) for dust
and soil samples from the Sahel, southern Morocco, the central Sahara, and
Tunisia. Nonetheless, previous studies have reported that the total iron
content of Saharan dust deposited in Europe ranges between 3.5 and
5.6 wt % Fe (Goudie and Middleton, 2001). Moreover, the Fe / Al
ratio of 0.54 in our samples falls within the range of values (0.50–0.57)
reported for red rains in northeastern Spain (Avila et al. 1998, 2007).
Similarly, the (Ca + Mg) / Fe ratio is 2.06, a value that matches
that reported by Avila et al. (2007) for Saharan dust from central Algeria
deposited in the northeastern Iberian Peninsula during red rain events. The
contents of K and Na are consistent with the presence of K-feldspar,
plagioclase, and clay minerals (i.e., abundant K-containing clays such as
illite and MLC). The amount of Ti is in good agreement with the amount of
rutile detected using XRD (∼ 1 wt % TiO2). Regarding
minor/trace elements, the amounts of Ba, Mn, Zn, and Sr, associated with
carbonates, as well as Ni, Cr, P, and V, associated with silicates, are in
good agreement with those reported for dust in red rains coming from northern
Saharan sources (Avila et al., 2007). Interestingly, the Pb content (27 ppm)
is very low and nearly identical to the average values (24 ppm) in Saharan
soils with negligible anthropogenic perturbation (Guieu et al., 2002).
Typically, mixing of Saharan dust plumes with polluted air masses (i.e., with
anthropogenic perturbations) leads to a significant increase in the
Pb / Al ratio in dust wet-deposited in the central and western
Mediterranean (Guieu et al., 2002). However, in our case the Pb / Al
ratio of 3.78 × 10-4 is very similar to the Pb / Al of
3.41 × 10-4 reported by Guieu et al. (2002) for Saharan
end-members. Similarly, the amount of S, an element commonly associated with
anthropogenic pollution, is very low. This value (0.04 wt % S) is
almost 2 orders of magnitude lower than those reported by Avila et al. (2007)
for red rains in the northeastern Iberian Peninsula, and is identical to that
of western Saharan desert soil (Castillo et al., 2008). In addition, V and
Ni, which are typically enriched in anthropogenic combustion aerosols
(Sholkovitz et al., 2009), showed a V / Al ratio of
0.1 × 10-2 and a Ni / Al ratio of
0.06 × 10-2. These values are very similar to the corresponding
values of the continental crust (V / Al = 0.08 × 10-2
and Ni / Al = 0.05 × 10-2) (Taylor and McLennan, 1985;
Sholkovitz et al., 2009). These results demonstrate that the close proximity
of the southern Iberian Peninsula to northern Africa and the rapid, intense
dust advection during this Saharan dust event led to negligible contamination
from polluted (e.g., European) sources (Lyamani et al., 2005).
In summary, the content of major, minor, and trace elements in our studied
samples is fully consistent with the reported composition of Saharan dust
samples (Scheuvens et al., 2013), and closely matches the composition of dust
deposited during red rain events in the western Mediterranean, especially
those of an “eastern” provenance, as defined by Avila et al. (2007), i.e.,
those whose source areas are located in the northern and central parts of the
Sahara (Algeria, Libya, and Tunisia). These compositional results are thus
consistent with the results of previous sections pointing to PSA1 and PSA3 as
the main source regions for the dust deposited in Granada during the studied
event.
TEM photomicrographs of Saharan dust particles. (a) General
low-magnification overview of the particles. Note the abundance of fibrous
particles. The long ones are palygorskite (Pal), while the short ones
(∼ 0.5 µm) are calcite (Cal). The larger particles or
aggregates corresponds to feldspars (plagioclase, Plg), illite (Il) and
smectites (Sm). The smaller particles include kaolinite (Kao) and rutile
(Rut). (b) Detail of elongated palygorskite fiber. The [001] zone
axis SAED pattern is shown in inset. (c) Aggregate of fibrous
calcite (SAED pattern in inset). (d) Plate-like kaolinite particles
(SAED pattern in inset). (e) Illite (SAED in inset), rutile,
palygorskite and fibrous calcite. (f) Goethite (Goe) (SAED pattern
in inset) and palygorskite. (g) Quartz (Qtz) grain (SAED in inset)
surrounded by smectites (Sm). The red circled area shows clay minerals plus
iron oxyhydroxide nanoparticles forming a rim around the quartz grain.
(h) Chlorite (Chl) (SAED in inset) with adhered smectite and iron
oxyhydroxide nanoparticles. The blue and the red arrows in the SAED pattern
show the diffuse Debye rings corresponding to the reflections with
d-spacing 4.49 and 2.5 Å, respectively (see the discussion in the main
text). (i) Amorphous silica structure (diffuse haloes in SAED
pattern in inset). This is part of the skeleton of a
diatom. (j) Rhombohedral calcite crystal (SAED in inset)
surrounded by kaolinite.
TEM–AEM analyses
TEM identification of individual mineral particles was performed combining
information provided by morphology, AEM microanalysis, and selected area
electron diffraction (SAED). Figure 6 shows representative TEM images of dust
particles and aggregates. Abundant fibrous particles were observed, either
with size >1-5 µm, made up of Si, Mg, and Al (EDX results) with a
SAED pattern matching that of palygorskite (Fig. 6a, b, e, f), or with size
∼ 0.5–1 µm, identified as calcite by their high Ca content
and SAED pattern (Fig. 6a, c). Abundant non-fibrous clay mineral particles
were also identified, which based on their composition (EDX results) and SAED
pattern were kaolinite (Fig. 6a, d, j), illite (Fig. 6a, e), smectite
(Fig. 6a, g–h), and chlorite (Fig. 6h). Goethite particles up to
∼ 250 nm in size were also identified by EDX and SAED (Fig. 6f) along
with scattered iron-rich nanoparticles <100 nm in size (identified by EDX
microanalysis), typically associated with clay minerals (Fig. 6h). These
nanoparticles did not produce any diffraction spots in the SAED pattern: only
the diffraction spots of the underlying clay minerals were observed along
with diffuse Debye rings at 4.5 Å, corresponding to the general
hk0 reflection of (poorly crystalline) clay minerals, and at
∼ 2.5 Å and at ∼ 1.5 Å, which can be ascribed to
amorphous and/or poorly crystalline two-line ferrihydrite
(Fe5HO8 4H2O) (Jambor and Dutrizac, 1998), with contributions
of higher-order hk0 reflections of the clay minerals (Fig. 6h,
inset). Amorphous and/or poorly crystalline iron-rich nanoparticles, possibly
ferrihydrite, in desert dust have been previously reported (Shi et al., 2009,
2011b, 2012), typically forming coatings on clay minerals (Wagner et al.,
2012) or appearing dispersed in the clay matrices of clay-rich particles
(Jeong et al., 2016). Feldspars (plagioclase) were also identified (Fig. 6a).
Interestingly, we observed silica particles with a complex structure (septa)
which were amorphous to the electron beam (i.e., their SAED showed no
diffraction spots or Debye rings, but diffuse haloes characteristic of an
amorphous phase) (Fig. 6i). These are standard features of mineralized
skeletons of diatoms, which are abundant in dust from the Bodélé
(Formenti et al., 2011), but also occur in fluvial deposits and ephemeral
lakes, as well as paleolake basins all across the Sahara (Shi et al., 2011b),
especially in central–southern Algeria and Tunisia, and are found in Saharan
dust deposited in marine sediments and continental Europe (Gasse et al.,
1989). Large quartz grains, typically covered by phyllosilicates and iron
oxyhydroxide rims, were also identified (Fig. 6g). The presence of clay- and
iron-rich coatings on large quartz and/or feldspars, as well as on carbonate
grains, appears to be a general feature of Saharan dust particles (Jeong et
al., 2016). Finally, calcite rhombohedra (Fig. 6j), as well as very small
(∼ 100 nm) rutile and scarce ilmenite (FeTiO3) crystals
(identified by both SAED and EDX microanalysis), were also observed.
Structural formulae of clay minerals in Saharan dust from TEM–AEM
analysis.
Illite based on O10(OH)2
Si
AlIV
AlVI
Mg
Fe
Sum oct. 1
K
Ca
Na
Sum int. 2
3.46 ± 0.10
0.54 ± 0.10
1.49 ± 0.28
0.20 ± 0.01
0.37 ± 0.33
2.06 ± 0.04
0.46 ± 0.17
0.06 ± 0.08
0.06 ± 0.08
0.63 ± 0.06
Palygorskite based on O10OH
4.08 ± 0.17
0.63 ± 0.25
1.10 ± 0.42
0.24 ± 0.19
1.97 ± 0.11
0.04 ± 0.04
0.03 ± 0.03
Kaolinite based on O5(OH)4
2.01 ± 0.02
1.91 ± 0.08
0.01 ± 0.03
0.06 ± 0.05
0.01 ± 0.01
Smectite based on O10(OH)2
3.58 ± 0.15
0.42 ± 0.15
1.54 ± 0.20
0.23 ± 0.08
0.34 ± 0.14
2.11 ± 0.07
0.14 ± 0.05
0.10 ± 0.07
0.33 ± 0.12
Illite–smectite mixed layer based on O10(OH)2
3.46 ± 0.29
0.54 ± 0.29
1.67 ± 0.14
0.34 ± 0.11
0.19 ± 0.08
2.20 ± 0.11
0.22 ± 0.11
0.04 ± 0.02
0.29 ± 0.16
1 Sum of octahedral cations. 2 Sum
of interlayer charge (M+ + M2+).
Table 4 shows average values for the structural formula of the different clay
minerals identified by TEM–AEM (see Table S1 for a complete list of all
individual AEM analyses). Microanalysis results confirmed the presence of
illite, palygorskite, kaolinite, smectite, and MLC. Due to its scarcity and
mixing with other clay phases, we could not collect any clean and reliable
AEM analysis for chlorite. Illite had a composition typical for this clay
mineral, with a substantial phengitic component as shown by the relatively
high Fe and Mg content in the octahedral layer (Weaver and Pollard, 1973).
Most palygorskite analyses (Table S1) had excess Si and relatively low Al and
Mg in octahedral sites. This is due to beam damage during AEM analysis of
this beam-sensitive mineral. Nonetheless, the average structural formula of
palygorskite obtained here is standard for this fibrous clay, with a slightly
elevated Fe content not unusual for this fibrous clay (Weaver and Pollard,
1973). The AEM analysis of kaolinite revealed a slight Al deficit and very
small amounts of Fe, an element that substitutes Al in the octahedral layer,
especially in the case of poorly crystalline kaolinite (Mestdagh et al.,
1980). We identified a smectitic phase with a composition compatible with
montmorillonite. The relatively high Al content in tetrahedral positions,
resulting in a reduced amount of Si, and the abundant Al and Fe in octahedral
positions, point to beidellitic and nontronitic contributions (i.e., solid
solution between the extreme terms montmorillonite and beidellite, with a
minor nontronitic component) (Weaver and Pollard, 1973). In addition, we
identified a MLC with a relatively high K content in the interlayer, which is
consistent with an illite–smectite mixed-layer phase (with a relatively low
content of illite layers), also with a relatively high Fe content. MLC can be
distinguished from smectite by its lower Si content and corresponding higher
Al content in tetrahedral positions, as well as by its higher K content. It
is also distinguished from illite by the reduced K content and interlayer
charge.
All analyzed clay minerals contained significant amounts of Fe, an important
finding regarding its biogeochemical (e.g., iron bioavailability) and
radiative implications. It could be argued, however, that the intimate
mixing/attachment of Fe-rich nanoparticles with/on clay minerals might yield
an excess of Fe during AEM analyses. With a few exceptions (see below), this
is considered unlikely: special care was taken to collect AEM analysis only
from areas of the clay mineral particles free of Fe-rich nanoparticles. In
any case, we carefully checked our individual AEM analyses in order to
identify possible contamination with Fe. Analyses that showed an anomalously
high Fe content incompatible with previous analyses of clays in Saharan dust
(Díaz-Hernandez and Párraga, 2008; Jeong and Achterberg, 2014;
Jeong et al., 2016) were not considered for the determination of the average
structural formulae reported in Table 4. Similarly, analyses that did not
yield an appropriate sum of octahedral cations (i.e., out of the range
1.8–2.2) were not used for the calculation of the structural formulae
presented in Table 4.
The average Fe content in the clay minerals varied between 1.7 wt % in
kaolinite and 5.2 wt % in illite, with intermediate values of 2.7, 3.4,
and 4.8 wt % in MLC, palygorskite, and smectite, respectively. These
values are consistent with those reported by Jeong and Achterberg (2014) and
Jeong et al. (2016) for clay minerals in Saharan dust analyzed using
TEM–AEM. Overall, the average Fe content for all clay minerals analyzed here
is 3.53 wt %. If one considers that there is also chlorite, which has a
typical Fe content of ∼ 14 wt % in desert dust (Jeong and
Achterberg, 2014), then the average Fe content in our studied clays rises to
5.27 wt %. This latter value is in very good agreement with the average
5.4 wt % Fe content of clay minerals in Saharan dust analyzed by Jeong
and Achterberg (2014).
STEM-HAADF photomicrographs and corresponding EDX elemental maps of
Saharan dust particles forming a micrometer-sized aggregate. Based on the
compositional analysis, particles of palygorskite (Pal), calcite (Cal), iron
oxyhydroxides (Fe-rich), rutile (Rut), illite (Il), smectite (Sm), kaolinite
(Kao), silica (SiO2), which according to SAED results is amorphous
(i.e., diatoms), plagioclase (Plg), and mixed-layer clays (MLC) were identified.
STEM/HAADF photomicrograph and corresponding EDX elemental maps of
a kaolinite (Kao) particle internally mixed with a nearly
rhombohedral-shaped goethite or hematite (Goe/Hem) crystal, a calcite (Cal)
particle, and aggregates of Fe-rich nanoparticles (nanogranular Fe
oxyhydroxides), likely ferrihydrite.
Further detailed textural and compositional insights were obtained using the
HAADF and EDX detectors of the Titan TEM operated in STEM mode. Figures 7 and
8 (as well as Figs. 4S and 5S) show dark field HAADF images (Z-contrast)
and corresponding elemental maps of representative aggregates and individual
particles. Most abundant particles appeared as micrometer-sized aggregates
made up of internally mixed silicate (quartz and amorphous silica),
aluminosilicate (feldspars), carbonate (calcite), and abundant phyllosilicate
particles (mainly kaolinite, illite, palygorskite, smectite, and MLC),
interspersed or covered with abundant iron-rich (nano)particles as well as
scarce titanium oxide (nano)particles. Note that dust aerosols can be
internally or externally mixed. Internal mixing refers to aggregates formed
by mineral particles of different composition, while external mixing involves
different mineral particles existing separately. The internal mixing of such
phases, in particular iron oxyhydroxides/clay minerals, has been previously
observed using standard TEM imaging (Jeong and Achterberg, 2014; Jeong et
al., 2016) as well as energy-loss TEM tomography (Deboudt et al., 2012). Most
of the iron-rich particles embedded and dispersed within the micrometer-sized
clay-rich aggregates, as well as those covering or attached to individual
clay particles, were <100 nm in size (Fig. 8). The tendency of iron
oxyhydroxides to concentrate in the smallest (<100–200 nm) size fraction,
typically forming nanogranular coatings or “nanoclusters” attached to clay
minerals, appears to be a general feature of Saharan dust (Lieke et al.,
2011; Wagner et al., 2012; Zhang et al., 2015). Such iron oxyhydroxide
nanoparticles were texturally different from the much larger individual
goethite crystals (compare Figs. 6f and 8). These observations further
suggest that these iron-rich nanoparticles are amorphous or poorly
crystalline ferrihydrite (Shi et al., 2009).
Interestingly, elemental maps of dust aggregates also revealed that Ca-rich
particles corresponding to calcite were commonly associated with the abundant
Fe-rich nanoparticles and clay minerals, forming internally mixed aggregates
(Figs. 7, 8, S4, S5). This contradicts the common assumption regarding
calcite mixing state in Saharan dust, which is considered to be external (Shi
et al., 2012). Moreover, no nitrogen enrichment (i.e., formation of calcium
nitrates) (Krueger et al., 2004) or sulfur-containing phases (such as gypsum)
resulting from the reaction of acid pollutant gases with carbonates were
detected (Figs. 7, 8, S4, S5). This is in agreement with the above-reported
very low Pb, V, and Ni concentrations, showing negligible mixing with air
masses including anthropogenically derived pollutants.
FESEM photomicrographs of Saharan dust particles. Large particles
(a) are made up of smaller, micrometer-sized particles
(b) which in turn are covered by nanoparticles (c).
(d) Detail of the nanogranular surface texture of dust particles.
The inset in (b) shows a representative EDX spectrum (corresponding
to the particle in the center of the image) showing a high content of Si, Al,
and Mg, with small amounts of K, Ca, and Fe. Very similar spectra are
recorded for most of the areas in the sample, demonstrating that the larger
particles (quartz, carbonates, and minor amounts of feldspars) are covered by
clays and iron oxyhydroxide (nano)particles. (e) Particle size
distribution (vol % PSD and cumulative curves) and (f) nitrogen
sorption isotherm of dust particles. The inset shows the BJH pore size
distribution.
Textural features of Saharan dust: FESEM, PSD, and N2 sorption
analyses
Figure 9a–d shows representative FESEM photomicrographs of dust mineral
particles. Abundant particles ∼ 3 to ∼ 30 µm in size
(silt fraction) were observed. Interestingly, the larger
particles (Fig. 9a) were typically ellipsoidal in shape with angular, sharp
edges, all common features of coarse Saharan dust particles (Wagner et al.,
2012; Jeong et al., 2016). They were made up of an aggregate of smaller,
well-cemented (i.e., the aggregates were intact after sonication),
micrometer-sized particles (Fig. 9b). The latter, in turn, displayed very
rough surfaces made up of nanoparticle aggregates (Fig. 9c, d). EDX
microanalyses showed that with the exception of some large quartz grains
(showing O and Si with trace amounts of Al and Fe in the EDX spectrum), or
carbonates (showing O plus Ca -calcite-, or Ca and Mg -dolomite-, with trace
amounts of Al, Mg, Si, and Fe in the EDX spectrum), the majority of the
particles did not display a clear single-phase EDX spectrum (see inset in
Fig. 9b), but included significant amounts of Ca, Mg, Al, Si, K, Na, and Fe
(and Ti). This is likely due to the contributions of clay minerals and iron
oxyhydroxides (and rutile) nanoparticles that covered nearly all dust
particles and aggregates (i.e., these phases were internally mixed). Surface
coatings of clays and oxyhydroxides (as well as other amorphous phases)
reportedly mask the identity of underlying mineral grains (Engelbrecht et
al., 2016). These results suggest that great care should be taken when using
SEM-EDX for the chemical analysis and/or mineralogical identification of
individual desert dust particles.
Figure 9e shows the volume particle size distribution (PSD) of the Saharan
dust determined by laser scattering. A polymodal PSD was observed, with a
principal mode at 17 µm, a secondary mode at 0.2 µm, and
additional less marked modes at 2 and 0.7 µm. A very similar
polymodal PSD has been previously reported for Saharan dust transported
during major outbreaks (D'Almeida and Schütz, 1983; Menéndez et al.,
2014; Titos et al., 2017), including those involving red rains (Sala et al.,
1996; Guieu et al., 2002). The samples studied had 24, 75, and 1 vol %
clay, silt, and sand content, respectively. In general, these values were in
agreement with the wt % content of the different size fractions
determined following hydrodynamic size separation. However, some differences
were observed, especially in the case of the clay fraction (i.e.,
24 vol % vs. 30 wt %). The latter can be explained by the fact
that iron oxyhydroxides, with a density ρ nearly double than that of
clay minerals (i.e., ρ of goethite, 4.3 g cm-3, and hematite,
5.3 g cm-3, vs. ρ of clay minerals,
∼ 2.6–2.8 g cm-3), were preferentially concentrated in the clay
fraction (see the XRD and TEM results, above).
Figure 9f shows the N2 sorption isotherm of Saharan dust. It was of
type IV, typical for mesoporous materials such as clay minerals, and showed a
type H3 hysteresis loop (Sing et al., 1985). The latter is due to the
presence of slit-shaped pores associated with aggregates of plate-like
(nano)particles (Sing et al., 1985), such as the clay minerals present in the
Saharan dust. The average surface area was 25 ± 1 m2 g-1.
The main contributors to such a relatively high surface area are the smallest
particles, that is, clays (especially smectites and MLC) and iron
oxyhydroxides (Elert et al., 2015). N2 sorption measurements yielded
a pore volume of 0.039 ± 0.002 cm3 g-1. The pore size
distribution determined using the BJH method (inset in Fig. 9f) was unimodal
with a maximum at 5 nm. The presence of relatively abundant (nano)pores, an
aspect of desert dust that has been typically ignored, is of relevance due to
their effect on the dust's reactivity and hygroscopicity because they would
strongly contribute to the exposed surface area of aggregates, making them
more reactive to atmospheric processing. Such pores could also affect the
dust radiative properties (Kemppinnen et al., 2015).
Combined, FESEM, PSD, and N2 sorption analyses show that although
clay-sized particles constitute a significant fraction of the Saharan dust,
and strongly contribute to its overall porosity and surface area, coarser
particles are particularly abundant. For instance, 42 vol % of the dust
particles were >10 µm in size. Although abundant coarse and even
giant particles have been reported for major Saharan dust events involving
dust entrainment and transport under strong wind conditions (Jeanicke and
Schütz, 1978; D'Almeida and Schütz, 1983; Coude-Gaussen et al., 1987;
Goudie and Middleton, 2001; Otto et al., 2007; Weinzierl et al., 2009;
Menéndez et al., 2014), most analyses of Saharan dust typically report
values <10 µm (Reid et al., 2003; Lyamani et al., 2005; Mahowald
et al., 2014). Note, however, that PSD analyses of Saharan dust are commonly
performed using cascade impactors or optical inversion techniques (e.g.,
AEronet RObotic NETwork, AERONET), which either typically exclude particles
with size >10 µm or underestimate their vol % (Raiswell,
2011). Conversely, in our case, a possible bias towards larger particle sizes
might occur due to the use of a laser scattering system for PSD analysis
(Reid et al., 2003). However, this effect appears to be minor as demonstrated
our SEM observations showing the presence of abundant particles larger than
10 µm (Fig. 9a). It could be argued that the larger particles
observed here may result from aggregation phenomena taking place after
in-cloud and/or below-cloud scavenging during the red rain event (Mahowald et
al., 2014) or during drying of the deposited red rain. Coarsening due to
aggregation of Saharan dust particles scavenged within cloud droplets
followed by drying prior to dry deposition has been claimed responsible for
the formation of the so-called “iberulites” (Cuadros et al., 2015;
Díaz-Hernández and Párraga, 2008; Díaz-Hernández and
Sánchez-Navas, 2016). Iberulites are nearly spherical, clay-rich giant
particles (∼ 100 µm in diameter) that are dry deposited in
Spain typically during summer Saharan dust events. However, we observed no
iberulite-like aggregates in our wet-deposited samples. According to Fiol et
al. (2005) such giant spherical structures rapidly disaggregate upon contact
with water. In contrast, the coarse and giant aggregates observed here show
sharp edges (Fig. 9a), typical for airborne desert dust (Weinzierl et al.,
2009), resulting from saltation–sandblasting at entrainment areas
(Coude-Gaussen et al., 1987; Mahowald et al., 2014). Such aggregates are
strongly cemented, as demonstrated by the fact that they do not disaggregate
by sonication during PSD analysis (i.e., no size reduction was observed over
the course of successive PSD analysis: -3 replicas per analysis) or by
sonication prior to TEM analysis. These results suggest that irreversible
aggregation (i.e., coarsening) due to cloud processing during the red rain
event or during drying following wet deposition is not a significant process
affecting the PSD of the Saharan dust studied here. Our results showing very
coarse particles (even giant ones) are consistent with the results of several
studies on Saharan dust analyzed during major outbreaks in different areas.
Analysis of Saharan dust both at ground stations and aloft (i.e., airborne
measurements such as those of SAMUN and FENNEC aircraft campaigns), typically
show similar PSD with a significant contribution of coarse and even giant
particles (Weinzierl et al., 2009; Ryder et al., 2013; Titos et al., 2017).
In agreement with our results, the recent lidar analysis of the extreme
Saharan dust event of February 2017 shows that the Ångström exponent
of the dust over Granada reached values close or equal to zero, meaning that
the dust particles in this event were extremely coarse (Fernández et al.,
2018).
Dissolution tests and geochemical modeling
It could be argued that some of the textural and structural features of the
most soluble phases identified here, such as the fibrous carbonates or the
amorphous/poorly crystalline ferrihydrite nanoparticles, could be the result
of dissolution/reprecipitation processes undergone by the dust particles
during the wet scavenging and the subsequent drying prior to sampling.
Similarly, it could be argued that reprecipitation of carbonates could lead
to the cementation of the aggregates, thereby playing a role in the
formation of the coarse/giant aggregates identified with SEM and laser
scattering PSD analyses. However, dissolution tests using
MilliQ® water with pH 5.6 (i.e., that of unpolluted rainwater)
and a dust / solution mass / vol ratio equal to that of the red rain event
(i.e., 18 g dust/2 L rainwater) showed that the amount of calcium dissolved
upon equilibration at pH 8 was only 0.22 mmol (Fig. S6). This value
corresponds to the dissolution of 2.1 wt % of the total amount of calcite
in the bulk dust. Such a very limited amount of dissolved calcite, which
would reprecicipitate upon drying, could not have resulted in the formation
of the abundant fibrous calcite observed here. In addition, such a dissolved
amount represents only 0.25 wt % of the total mass of the bulk dust: its
reprecipitation could not therefore lead to any significant cementation
effect that might alter the dust particle size distribution. Reprecipitation
of dissolved calcium carbonate during drying of the red rain would only lead
to a very limited regrowth of existing calcite crystals (i.e., regrowth on
an existing calcite crystal would by the most energetically favorable
situation). Such a limited amount of dissolved calcite is consistent with
our PHREEQC computer modeling results (Fig. S6 and Table S2). The amount of
dissolved Ca following equilibrium (at the final pH of 8) of the different
phases present in the dust (with their corresponding mass fractions
determined by XRD analysis) is just 0.24 mmol. This value, representing
2.3 wt % dissolved calcite, is in excellent agreement with that of the
dissolution experiment. In the case of the iron oxyhydroxides, PHREEQC
modeling showed that in simulations considering just goethite and hematites,
a negligible 3 × 10-8 and 9 × 10-9 wt %
dissolution for the former and latter phase, respectively, took place.
Conversely, with the additional presence of ferrihydrite, 0.31 wt % of
this phase dissolved. Reprecipitation of such a very small amount of
dissolved Fe cannot account for the abundant amorphous/poorly crystalline
ferrihydrite nanoparticles observed using TEM. It follows that other
processes have to be responsible for the formation of such Fe-containing
nanoparticles, as we will discuss below.
Interestingly, PHREEQC modeling (Table S2) also shows that the SI values for
the rest of the phases present in the dust, with the exception of K-feldspar,
smectite, rutile and palygorskite, were >0. This means that the system was
supersaturated with respect to them and, therefore, they could not undergo
dissolution. In the case of K-feldspar, rutile and smectite, the calculated
dissolved amount was almost negligible. However, in the case of palygorskite,
the amount dissolved was 5.2 wt %. The release of Mg from this clay
explains why the modeled system was supersaturated with respect to dolomite
(which did not dissolve). Nonetheless, this latter result has to be taken
with care, because the solubility product used for PHREEQC modeling of the
dissolution of palygorskite was that of the very similar (compositionally and
structurally) sepiolite, which may or may not be exactly the same of
palygorskite. Note that we used the solubility product of sepiolite because
to the best of our knowledge, there is currently no reported value for the
solubility product of palygorskite. Note also that we did not include MLC and
chlorite in the PHREEQC simulations, because in the first case there is no
solubility product reported for the specific composition observed here, and
in the latter case, we do not have the actual structural formula for this
phase. In any case, their amount in the bulk dust is very minor and should
not have any significant impact in the outcome of the PHREEQC modeling.
The results discussed above refer to the dissolution of the wet-deposited
dust in the limited amount of rainwater that fell during the red rain event.
Deposition of the Saharan dust in a larger volume of water (e.g., seawater)
would therefore result in a much larger dissolved amount of the different
phases present in the dust. The latter is important when considering the
biogeochemical impact of such a red rain event on inland and sea/ocean
waters.
TG/DSC traces of Saharan dust collected in air (blue curves) and
in inert N2 (red curves) atmospheres.
TG-DSC and spectroscopic analyses
TG-DSC analyses showed a first weight loss between 100 and 700 ∘C
corresponding to the dehydration and dehydroxylation of clay minerals
(Guggenheim and van Groos, 2001) (Fig. 10). A minor contribution to such a
weight loss was due to the dehydroxylation of ferrihydrite and goethite,
which reportedly occurs at ∼ 150 and ∼ 280–400 ∘C,
respectively (Jambor and Dutrizac, 1998). These dehydroxylation processes are
endothermic, as shown by the endothermic broad band in the DSC trace. Organic
matter also decomposes in this T interval (Elert et al., 2015). The
presence of organics led to a weight loss difference of
∼ 0.6–1.1 wt % between runs carried out in air and in inert
N2 atmospheres (Fig. 10). Organic carbon undergoes oxidative
combustion in air, which was reflected here by an exothermic peak at
350 ∘C in the DSC trace. Such an exothermic peak was absent in the
run carried out in N2. Note that the above-indicated weight loss
values only represent a fraction of the organic carbon present in the Saharan
dust, because upon oxidative combustion, elemental carbon will be produced
(charring effect) which we could not quantify. In any case, the values
presented above are consistent with reported average value of organic carbon
in Saharan dust aerosol (e.g., 1.7 wt %: Gonçalves et al., 2014).
Analysis of organic matter in desert dust has shown that the elemental plus
organic carbon content is on average <2 wt %, with a large
compositional variability (Jaenicke and Schütz, 1978; Eglinton et al.,
2002; Gonçalves et al., 2014). Elemental carbon is typically associated
with biomass burning (Eglington et al., 2002), whereas organic carbon has
been associated with more or less decomposed biological residues, including
micro-organisms and microbial biofilms, as well as humic substances (Conen et
al., 2011). Micro-organisms and organic residues in desert dust aerosol,
which tend to cover mineral grains, have been suggested to be efficient
atmospheric ice nuclei, thereby having an important indirect radiative
forcing effect (Conen et al., 2011).
At T>700 ∘C carbonates (first dolomite and subsequently calcite)
decompose into CaO (or CaO + MgO in the case of dolomite), releasing
CO2 (Rodriguez-Navarro et al., 2009, 2012). This is an endothermic
process, as shown by the DSC traces. By measuring the total weight loss in
the 700–950 ∘C interval, the total amount of carbonates was
calculated to be 14.7 ± 0.2 wt %, a value in very good agreement
with the values determined by XRD (RIR) and weight loss following acid
treatment.
Spectral features of Saharan dust. (a) ATR-FTIR spectrum of dust.
The inset shows absorbance values for high wavenumbers (IR spectral region
with lower wavelength values: 2.6–4 µm). Wavenumbers of main absorbance bands
are indicated; (b) UV-Vis-NIR spectra of bulk Saharan dust,
and its three size fractions (clay, silt and sand).The inset shows the
second derivative of the absorbance spectrum of the bulk sample. Blue arrows
show the bands used for the calculation of goethite (Goe) and hematite (Hem)
relative contents in the bulk sample.
Figure 11a shows the FTIR spectrum of the collected desert dust. The broad
band centered at ∼ 3400 cm-1 corresponded to υOH
stretching of interlayer (solvation) H2O of clay minerals, and
structural water in ferrihydrite (Russell, 1979). The sharper peaks at 3693
and 3617 cm-1, and the shoulder at 3637 cm-1, corresponded to the
υOH stretching of well-ordered (crystalline) gibbsite layers of
dioctahedral phyllosilicates. The three bands are characteristic of
kaolinite. We could not resolve the broader smectite and illite bands at
3650–3550 cm-1. The small broad band at 3550 cm-1 corresponds to
the υOH stretching of the brucite layer in palygorskite (inset in
Fig. 11a). Note, however, that ferrihydrite also has a characteristic
υOH stretching band at ∼ 3615 cm-1 (Russell, 1979). We
could not resolve the υOH band of goethite at
∼ 3150 cm-1, because it was masked by the broad OH band of clay
minerals. However, the bands corresponding to the δOH in-plane
deformation at 870 cm-1 and the γOH at 780 cm-1
(Schwertmann et al., 1985) were observed, confirming the presence of this
iron oxyhydroxide. Note, however, that the band at 870 cm-1 can also
have the contribution of the υ2 out-of-plane vibration of
CO32- groups in calcite and dolomite. The small bands at 2898 and
2984 cm-1 corresponded to the C-H stretching of organic matter, thereby
confirming the TG results showing the presence of organics in the Saharan
dust. However, the lack of other well-defined bands precluded an unambiguous
identification of the specific organic compounds in the dust. Nonetheless,
the presence of a band at 1624 cm-1, which can be ascribed to both
δOH bending (which should be centered at 1644 cm-1) and
COO- symmetric stretching, points to the presence of carbohydrates
(e.g., polysaccharides) with carboxylic functional groups. The latter would
be consistent with the presence of microbial exopolymeric substances and/or
humic substances (Conen et al., 2011). The broad band at 1406 cm-1
corresponded to the υ3 anti-symmetric stretching of CO32-
groups in both calcite and dolomite. The presence of calcite was confirmed by
the very small υ4 bending band of CO32- groups at
712 cm-1. The strong band at 986 cm-1 and the shoulder at
1030 cm-1 corresponded to the Si–O stretching of phyllosilicates,
while the shoulder at 1090 cm-1 corresponded to the Si–O stretching of
quartz. The small peak at 908 cm-1 corresponded to the Al–OH
deformation of kaolinite. The doublet at 780 and 795 cm-1 corresponded
to the Si–O bending of quartz (note that these latter bands overlap with
those of goethite). The strong band at 515 cm-1 corresponded to the
Al–O–Si deformation in illite and smectites (Di Biagio et al., 2014).
Finally, the very strong band at ∼ 410 cm-1 corresponded to the
FeO6 lattice mode of iron oxyhydroxides, including goethite, hematite,
and ferrihydrite.
The results of the FTIR analysis confirm the presence of the different
mineral phases identified using XRD and TEM–AEM, as well as the presence of
organic matter. They also show that this type of dust aerosol possesses
strong longwave absorption, especially at thermal IR (6–24 µm) due
to the abundant silicate, aluminosilicate and carbonate phases, along with
minor iron oxyhydroxides. Absorption of IR radiation by desert dust has a
direct positive (warming) radiative forcing and, most importantly, it does
not only operate during the day, as in the case of solar radiation, but also
at night due to (terrestrial) thermal emission (Di Biaggio et al., 2014).
Figure 11b shows the UV-Vis-NIR absorption spectra of the bulk dust and its
three size fractions. In all cases, a strong increase in absorption at
λ<600 nm was observed, and the absorbance was systematically higher
for the clay fraction. The observed increase in the absorbance of the
OH-stretching overtone (1412 nm) and the combination band of H2O
(1920 nm) of the clay fraction (Gionis et al., 2006), compared to those of
the coarser fractions, are consistent with the higher amount of clay minerals
and iron oxyhydroxides in this finer fraction, both being responsible for the
systematically higher shortwave absorbance. Remarkably, retrieved spectra
(λ<1 µm) of the imaginary part of the complex refractive
index, k, for several Saharan dust samples (Wagner et al., 2012) show
strong similarities to the UV-Vis-NIR spectra shown here.
The UV-Vis-NIR spectra of the studied samples are standard for Saharan dust
and have been associated with the
presence of iron oxides, mainly goethite and hematite (Wagner et al., 2012;
Formenti et al., 2014b). These oxides show a remarkable increase in k
values at λ<600 nm due to strong absorption associated with
ligand-to-metal (i.e., O-to-Fe) charge transfer transitions (Sherman and
Waite, 1985). This is also the case for rutile, which has values of k for
UV radiation on the same order of magnitude of hematite (Utry et al., 2015), and
should also contribute to the absorption spectra reported here. Note,
however, that iron in the structure of clay minerals also contributes to the
absorption of shortwave solar radiation because such octahedrally coordinated
Fe3+ can also experience charge transfer transitions. UV-Vis
spectroscopic analyses of clays such as illite, kaolinite, palygorskite, and
smectites, with Fe contents very similar to those of the corresponding clays
studied here, show strong absorption at λ<600 nm due to oxo-to-iron
(III) charge transfer (Karickhoff and Bailey, 1973).
From the second derivative of the UV-Vis spectrum (inset in Fig. 11b), the
contents of goethite and hematite were calculated to be 49 ± 1 %
and 51 ± 1 %, respectively. Our XRD results showed a higher
goethite content of ∼ 77 wt %. We consider, however, the UV
spectroscopy results more reliable, due to the relatively high error
associated with the semiquantitative XRD analysis. The goethite content
calculated using the second derivative of the UV-Vis spectrum is consistent
with but slightly lower than the contents reported by Formenti et al. (2014b)
for Saharan dust. The authors found goethite contents ranging from 52 up to
78 wt %. This discrepancy might likely be due to differences in the dust
source areas. Note that an accurate evaluation of the content of these two
oxyhydroxides is of outmost importance when determining the direct radiative
effect of desert dust, because the optical properties and, particularly, k
values of hematite and goethite differ significantly, the former phase
showing stronger absorption at short wavelengths (Zhang et al., 2015).
Iron significance: bioavailability
Iron in desert dust is a key player in a range of global biogeochemical
processes. Iron is an essential micronutrient for all organisms, but is
typically depleted in some inland water bodies, such as oligotrophic lakes
(Vrede and Tranvik, 2006), and a large portion of open ocean waters (Jickells
et al., 2005). Desert dust can supply iron to such areas, especially open
ocean waters, enabling the proliferation of a range of microorganisms (e.g.,
phytoplankton) and the sequestration of CO2 as biomass, thereby
directly affecting primary production and indirectly influencing climate
(Jickells et al., 2005). However, for its bioavailability, iron must be
soluble or at least in a colloidal, poorly crystalline nanosized state, which
has been defined as filterable or soluble iron, which passes through a
0.2 or 0.4 µm filter (see the review by Raiswell and Canfield,
2012). But iron in desert dust minerals is typically insoluble (Shi et al.,
2009). Three types of iron in desert dust have been defined (Shi et al.,
2012): (i) amorphous or poorly crystalline iron oxyhydroxides (e.g.,
ferrihydrite), which form the most soluble iron fraction in desert dust; (ii)
crystalline iron oxyhydroxides (goethite, hematite, and magnetite) that are
highly insoluble. These two groups form the so-called free iron; and (iii)
structural iron, which is incorporated into the crystalline structure of
aluminosilicates, mainly in clay minerals, and which is also insoluble.
Remarkably, the amount of soluble Fe in desert dust, expressed as the
fractional iron solubility, FFS (i.e., fraction of soluble iron vs. total
iron), typically increases during long-range transport from values
∼ 0.1 % to up to ∼ 80 % (Journet et al., 2008; Mahowald
et al., 2009; Shi et al., 2012). It has been proposed that this is due to
(i) mixing with anthropogenic (Sholkovitz et al., 2009, 2013) and/or biomass
(Guieu et al., 2005; Paris et al., 2010) combustion aerosol with higher
solubility than mineral dust (Mahowald et al., 2009); (ii) physical size
sorting due to preferential deposition of the larger particles during
transport, thereby increasing the relative amount of smaller particles with a
higher surface-to-volume ratio and enhanced Fe solubility (Baker and
Jickells, 2006). Note, however, that recent research has shown that size
sorting has a very minor effect (if any) on Fe solubility (Shi et al.,
2011b);
and (iii) various atmospheric processes including photoreduction (which
according to Zhu et al. (1999) has only a minor effect), organic-mediated
complexation (Paris and Desboeufs, 2013), and chemical (acid) in-cloud and/or
aerosol processing (Shi et al., 2009, 2012). Such an atmospheric processing,
as well as the ultimate fate of iron once desert dust is deposited in distant
locations, is strongly linked to where and how iron is incorporated into
different desert dust minerals with different structure, solubility,
crystallinity, and particle size. However, as pointed out by Raiswell and
Canfield (2012), few studies of aeolian dust contain any detailed
characterization of the iron mineralogy.
Our XRD results showed the presence of 1.41 wt % free Fe in crystalline
iron oxyhydroxides (goethite and hematite). Our AEM analysis of individual
clay minerals showed that structural Fe was also abundant in the Saharan dust
particles. Considering the average Fe content in the different clay minerals
identified here (Table 4) and their fractional content in the bulk sample
(Table 2), the structural iron in the clay minerals amounts to 1.98 wt %
Fe, that is,, 55.6 % of the average total Fe (see Table S3 for details
regarding the calculation of structural Fe). This yields a total
(free + structural) iron content of 3.39 wt %. This value is lower
than the average total iron content of 3.56 wt % determined by ICP-OES
(3.43 wt %) and XRF (3.69 wt %) analyses (see Table 3). It could be
argued that the missing Fe (0.17 wt %) is incorporated in other
silicates such as feldspars. However, the amount of Fe in such phases is
either very low or negligible. If we consider the values of
0.13–0.54 wt % Fe in feldspars reported by Journet et al. (2008) and
the fractional content of feldspars (0.13) in our dust samples, this yields
0.01–0.07 wt % structural Fe in such tectosilicates. These values
cannot account for the missing iron. It could also be argued that the missing
iron is in the carbonates. However our AEM analyses of carbonate phases
showed no Fe (Fig. S7). It follows that the missing (free) iron has to be
incorporated in the amorphous and/or poorly crystalline ferrihydrite detected
using TEM–SEAD, which cannot be quantified as free iron by XRD because this
latter technique does not identify amorphous phases. The amorphous and/or
poorly crystalline ferrihydrite would thus amount to an average of 11 %
of the free iron (i.e., 4.9 % of the total iron). If we consider the
lowest and highest values of total iron determined by ICP-OES and FRX
analyses, then we obtain that the amount of amorphous and/or poorly
crystalline ferrihydrite can range between 2.9 and 18.5 % of the free
iron (i.e., 1.3–8.2 % of the total iron). It should be noted that these
values, calculated based on the amount of iron in each Fe-containing phase
determined by XRD and the Fe content in clay minerals determined from
TEM–AEM analyses, are subjected to significant uncertainty. Therefore, the
amount of structural and free iron, including the amount of amorphous and/or
poorly crystalline Fe-oxyhydroxides (ferrihydrite) presented above are not
intended to be purely quantitative results, but a (rough) estimate of the Fe
speciation in the studied dust. Nonetheless, these results confirm TEM
observations showing that the amount of nanosized amorphous and/or poorly
crystalline ferrihydrite is not negligible.
Interestingly, leaching experiments using a small dust / solution mass / vol
ratio, as it is commonly done in studies on dust aerosol Fe solubility
(e.g., Journet et al., 2008; Buck et al., 2010; Shi et al., 2011a, 2012)
showed that in the case of MilliQ ® water (pH 5.6) and an
ammonium acetate buffer solution (pH 4.7) the values of FFSs were
22.7 ± 1.1 % and 17.9 ± 0.9 %, respectively (i.e., average
value of ∼ 20 %). These values are a factor of two larger
than the highest value for amorphous and/or poorly crystalline ferrihydrite
presented above. This could either means that our calculation of the amount
of amourphous and/or poorly crystalline ferrihydrite was too conservative,
or that there is another source for soluble Fe of, at least, a similar
magnitude than that of nanosized ferrihydrite. The latter is the most likely
case, being clays the most likely source of such a soluble Fe (Journet et
al., 2008).
The amount of structural Fe in our studied dust is consistent with, although
slightly higher than the values of ∼ 40–50 % commonly reported for
Saharan dust (Formenti et al., 2011, 2014a, b). Structural iron in clays has
been proposed as an important potential source of bioavailabable Fe (Journet
et al., 2008; Formenti et al., 2014a). In dioctahedral clays, such as illite
and kaolinite, as well as montmorillonite and MLC, iron is present in the
octahedral layer as Fe3+, alone or in association with minor amounts of
Fe2+ (Weaver and Pollard, 1973; Mestdagh et al., 1980; Johnston and
Cardile, 1987). In the case of palygorskite, which has a mixed di- and
trioctahedral character, Mg2+ in the octahedral layer can be substituted
by both Fe2+ and Fe3+ (Gionis et al., 2006). In trioctahedral
chlorite iron is incorporated as Fe2+ substituting Mg2+ in
octahedral positions both in the 2 : 1 structural unit and in the
interlayer brucite layers (Weaver and Pollard, 1973). Interestingly, Cwiertny
et al. (2008) observed a correlation between Fe2+ content in
aluminosilicates and increased FFS after acid processing of desert dust, and
Schroth et al. (2009) concluded that Fe2+-bearing silicates are more
soluble than iron oxyhydroxides, being the former an important source of
bioavailable soluble iron. In addition to the effect that Fe2+ may have
on the potential solubility of structural iron, a significant amount of our
studied clay particles with relatively high Fe content were <100 nm in
size, which makes them significantly more soluble than larger particles
(Raiswell and Canfield, 2012). These compositional and size effects may favor
the release of soluble structural iron upon atmospheric processing, and/or
post-atmospheric processing once the dust particles are deposited in the
ocean or in inland water bodies. Overall, it is very likely that the
relatively high FFS (i.e., ∼ 20 %) of the Saharan dust determined
by our leaching tests is not only related to the presence of soluble
nanosized Fe-oxyhydroxides (i.e., amorphous and/or poorly crystalline
ferrihydrite), but also to the presence of Fe-containing clays, which are
considered as an important contributor to the pool of soluble Fe (Journet et
al., 2008; Cwiertny et al., 2008).
In agreement with previous studies, the free iron in our samples is mainly
present as hematite and goethite (Lázaro et al., 2008; Shi et al., 2012;
Formenti et al., 2014b). However, we detected a significant amount of
amorphous/poorly crystalline ferrihydrite with a calculated upper bound of
8.2 % of the total iron, which is in very good agreement with the upper
bound value of 7.4 % reported by Shi et al. (2011b) for Saharan soils.
These values are, however, markedly smaller than the ferrihydrite content of
∼ 71 % reported by Schroth et al. (2009) for northern African dust
deposited on a buoy in the northeastern Atlantic Ocean. The existence of such
a huge concentration of ferrihydrite in Saharan dust has been, however,
questioned by Shi et al. (2011b) and Raiswell and Canfield (2012) on the
basis that the quantification of such an amorphous or poorly crystalline
phase is difficult. The authors also state that the very low FFS (<1 %)
reported by Schroth et al. (2009) is not consistent with such a huge amount
of ferrihydrite, because experimental evidence shows that ferrihydrite
nanoparticles are significantly more soluble than clay minerals and/or
crystalline goethite or hematite. Shi et al. (2012) conclude that
ferrihydrite nanoparticles, such as those identified here, are the most
likely source of soluble and bioavailable Fe in desert dust. How, where, and
when such Fe-rich nanoparticles are formed has been a matter of intensive
research and discussion. Two not mutually exclusive possibilities have been
considered: (i) iron-rich nanoparticles are already present in the dust
source region, and are formed due to (limited) weathering of crystalline
iron-containing phases such as clays (see above) and/or iron oxyhydroxides
(goethite and hematite) (Shi et al., 2011b; Raiswell, 2011), and/or (ii) they
are formed during dust transport via atmospheric processing (Shi et al.,
2012).
With a few exceptions (e.g., Shi et al., 2011b), the presence of amorphous
or poorly crystalline iron oxyhydroxides such as ferrihydrite in Saharan
dust source regions has been overlooked (e.g., Lafon et al., 2006; Debout et
al., 2012; Formenti et al., 2014b). It is considered unlikely that
metastable ferrihydrite could survive in the source region without rapidly
transforming into more stable goethite or hematite (Shi et al., 2012).
Nonetheless, ferrihydrite has been shown to remain untransformed for a few
hundred days at STP conditions (Raiswell and Canfield, 2012), and silica,
clay minerals, and a range of organic substances (all present in Saharan
soils and dust) reportedly contribute to its stabilization (Jambor and
Dutrizac, 1998). It is thus very likely that the amorphous iron oxyhydroxide
and/or poorly crystalline ferrihydrite nanoparticles in our dust samples
were already present in the entrained soil. Indeed, dissolution of
iron-containing clay minerals following intermittent exposure to aqueous
solutions during, for instance, rain events, fluvial transport and/or
flooding in (ephemeral) lake waters, is a plausible mechanism for the
formation of iron-rich nanoparticles (ferrihydrite) in the dust source areas
(Shi et al., 2011b; Canfield and Raiswell, 2012). This is consistent with
results by Poulton and Raiswell (2005) showing that clay minerals in natural
riverine environments commonly are associated with iron-rich nanoparticles.
Regarding the second hypothesis for the formation of iron-rich nanoparticles
in Saharan dust, numerous field and laboratory studies have shown that
atmospheric processing strongly contributes to the formation of soluble iron
(Shi et al., 2012). Shi et al. (2009) showed that Saharan dust wet deposited
in the western Mediterranean included highly soluble, bioavailable, and
poorly crystalline nanosized two-line ferrihydrite, which contained trace
concentrations of Al, Cr, Si and Ca. The latter might indicate formation by
atmospheric processing of clay minerals. The authors found no such nanosized
ferrihydrite in Saharan dust dry deposited in the eastern Mediterranean
(collected during a different event and presumably coming from a different
source area). Parallel experiments using Saharan soils and pure synthetic
goethite confirmed that precipitation of nanosized ferrihydrite occurs after
acid leaching, favored during (partial) drying and formation of wet mineral
aerosols, and subsequent pH increase during simulated in-cloud processing
(Shi et al., 2009, 2012). These observations might suggest that acid cloud
processing during wet deposition led to the formation of iron rich
nanoparticles. However, the lack of spatial and temporal relationship between
the wet and dry deposition events studied by Shi et al. (2009) precludes
drawing any final conclusion regarding whether the iron nanoparticles were
already present in the entrained desert dust prior to wet deposition or were
the result of atmospheric processing. In any case, a prerequisite for
in-cloud and aerosol acid processing of iron-containing phases is the
interaction with acid pollutant gases (e.g., SO2 and
NOx), and the absence of carbonates or their external
mixing. Ito and Feng (2010) underlined that internally mixed carbonates will
buffer atmospheric acid-processing of iron-containing phases, thereby
strongly limiting the formation of potentially bioavailable, poorly
crystalline or amorphous iron phases. Our HAADF analyses (Figs. 7, 8) clearly
show that carbonates were internally mixed with iron oxyhydroxide
nanoparticles and were not affected by acid dissolution. Moreover, we
detected no sulfate by-products such as gypsum or (calcium) nitrates. It
follows that the iron-rich nanoparticles in the Saharan dust deposited during
the studied red rain event must have already been present in the source
areas. This is an important result when considering the bioavailability of Fe
in Saharan dust. It shows that a significant fraction of Fe in Saharan dust
is already present in the source region as potentially bioavailable nanosized
amorphous and/or poorly crystalline iron oxyhydroxides (Shi et al., 2011b).
It also shows that the presence of acid gases and their associated
atmospheric acid-processing of iron-phases is not absolutely necessary to
have nanosized, more soluble and potentially more bioavailable Fe-rich phases
in desert dust. Our results suggest that the importance of such anthropogenic
acid gases in enabling the delivery of soluble and bioavailable Fe in desert
dust to open oceans, and their subsequent impact on CO2 drawdown (Li
et al., 2012) might be overestimated.
Another important aspect to consider regarding the bioavailability of
iron-containing phases is their interaction with organic compounds such as
carboxylic acids, or more complex molecules having different functional
groups. They can complex Fe, facilitating the dissolution of iron-containing
phases (especially clay minerals) and its bioavailablity (see review by Shi
et al., 2012), as demonstrated for the case of several organic acids, such
as oxalic and humic acids (Paris and Desboeufs, 2013). Our TG/DSC and FTIR
results showed the presence of organic carbon. It is however not clear what
role such organic carbon played in the possible processing of iron
oxyhydroxides during Saharan dust transport and scavenging. We can only
hypothesize that the presence of abundant functional groups (e.g.,
carboxylic groups) in such organics could enable the complexation of Fe and
facilitate its bioavailability.
Finally, it should be pointed out that the FFS of the studied Saharan dust
(∼ 20 %) is in the same range as (or even higher than) the FFS
reported for aerosol deposited in the Mediterranean (10–11 %), tropical
North Atlantic (3–35 %), and Barbados (6 %) (data compiled in Fan et
al., 2006). Note, however, that in most cases reported in the literature,
anthropogenic inputs and/or atmospheric processing likely contributed to the
measured FFS. Aerosol samples collected in the tropical North Atlantic
directly associated with air masses arriving from northern Africa and with no
anthropogenic mixing have FFS values of only 0.4–2 % (Sholkovitz et al.
2009), although higher values of 3–17 % have also been reported (Buck et
al., 2010). In our case, mixing with anthropogenic combustion aerosols or
atmospheric acid processing is ruled out. The high FFS of our samples, very
similar to the upper bound reported by Buck et al. (2010), is mainly due to
the structural/textural, compositional, and mineralogical features of the
dust present in the source areas activated during this extreme event. Due to
the magnitude of the event, and the relatively high FFS of the Saharan dust,
the biogeochemical impact of its deposition in inland waters as well as in
the Mediterranean and North Atlantic waters affected by the event could have
been very significant.
Effects of mineralogy, mixing state, and PSD on dust direct radiative
forcing
Scattering and absorption of incoming solar (shortwave) and outgoing thermal
(longwave) radiation by desert dust aerosol have a cooling effect at land
surface and a warming effect at tropospheric levels (Carlson and Benjamin,
1980; Alpert et al., 1998). However, the magnitude and even the sign of the
direct radiative forcing are not well constrained. While some researchers
consider that the net radiative forcing of mineral dust on the climate
system is negative (Gieré and Querol 2010; Allen et al., 2016), others
report that under specific scenarios the direct forcing can be positive
(Carlson and Benjamin, 1980), leading to regional (Overpeck et al., 1996) or
even global warming (Kok et al., 2017). The uncertainties regarding the sign
and magnitude of the direct radiative forcing of desert dust are rooted in
the fact that they depend on many poorly constrained factors such as: (i) the
characteristics of mineral dust (concentration, vertical distribution,
PSD, shape, internal/external mixing, and composition/mineralogy) and (ii) external
variables such as surface albedo below dust plumes, temperature at
ground level, and presence/absence of clouds (Balkanski et al., 2007;
Kemppinen et al., 2015). Composition/mineralogy and particle size, as well
as mixing state, appear to be the most critical factors controlling dust
direct radiative forcing (Zhang et al., 2015).
Regarding composition/mineralogy, iron-containing phases play a key role in
the absorption and scattering of solar and terrestrial radiation (Tegen and
Lacis, 1996; Sokolik and Toon, 1999; Zhang et al., 2015). However, most
models for the direct radiative forcing of desert dust typically only
consider the presence of hematite (e.g., Balkanski et al., 2007; Wagner et
al., 2012). Our results, as well as those of others (e.g., Formenti et al.,
2014b; Zhang et al., 2015), show that this is an oversimplification that may
have an important impact on the outcome of such models, because in addition
to hematite, goethite and iron oxyhydroxide nanoparticles (ferrihydrite) are
also present in significant amounts in Saharan dust. In addition, their
actual mixing state is another important aspect of desert dust that has been
generally overlooked in climate models. While an external mixing of
(alumino)silicates and iron oxyhydroxide particles in proportions typically
found in desert dust has a net negative radiative forcing, their internal
mixing can lead to a net positive radiative forcing due to a change in the
effective refractive index resulting in enhanced absorption (Sokolik and
Toon, 1999). The latter situation is the one observed here: iron
oxyhydroxides, generally forming nanosized aggregates and concentrated in the
clay fraction, are internally mixed with carbonate, silicate, and
aluminosilicate particles (SEM and TEM results). This appears to be a general
feature of Saharan (Deboudt et al., 2012; Wagner et al., 2012; Jeong et al.,
2016) and Asian (Jeong and Achterberg, 2014) desert dust. We observed that
iron oxyhydroxides are the main contributors to the strong absorption of
UV-Vis radiation regardless of size fraction (Fig. 11), which is consistent
with our electron microscopy observations showing aggregates of
iron-containing nanoparticles closely cemented with clay minerals, and
covering (i.e., internal mixing) larger silt- and sand-sized particles
(Figs. 6–8). Note, however, that clay minerals containing structural iron
can also contribute to the absorption of shortwave solar radiation as
indicated above. A few models consider the light absorption behavior of
illite alone or internally mixed with iron oxyhydroxides, showing that illite
actually displays relatively strong shortwave absorption (Zhang et al.,
2015). However, no model has ever considered the effect that structural iron
has on the absorption properties of other typical clay minerals present in
desert dust, such as smectites, kaolinite, palygorskite, chlorite, and MLC.
It should be noted that the combined effect of clay minerals and iron
oxyhydroxides is actually responsible for the shortwave radiative effects of
the studied Saharan dust. The longwave radiative effect, basically absorption
of thermal radiation, is mainly associated with silicate and aluminosilicate
phases (including clay minerals) as well as carbonate phases. It could be
argued that the combined shortwave and longwave scattering/absorption of
internally mixed clay minerals/iron oxyhydroxides/(alumino)silicates or
carbonates will be the relevant and overall radiative effect of the Saharan
dust. According to Sokolic and Toon (1999), such an internal mixing would
likely have a net positive direct radiative forcing under specific
circumstances (i.e., surface albedo, dust load, vertical distribution, and
surface T).
In addition to these compositional/mineralogical and mixing effects on dust
radiative forcing, another critical aspect to be considered is the PDS of
desert dust. Smaller particles (i.e., clay fraction) are more effective in
scattering solar radiation than larger ones. The latter, in turn, are more
effective in absorbing energy (solar and thermal) (Tegen, 2003; Otto et al.,
2007). Most models of dust radiative forcing typically consider the smallest
particles only (geometric diameter <10 µm) (Tegen and Lacis,
1996; Tegen, 2003). Such small sizes, however, do not represent the actual
size of dust particles in major dust events, which in turn are the ones that
most significantly contribute to entrain and transport desert mineral dust
(Skonieczny et al., 2011; Mahowald et al., 2014; Kok et al., 2017). Indeed,
one of the causes of uncertainty in climate models is that the size
distribution of dust particles is poorly constrained and typically the amount
of smaller particles is overestimated (Kok, 2011). Although large
(>10 µm) particles have been considered to settle by
gravitational forces within hours after entrainment (Tegen, 2003),
experimental evidence shows that large, and even giant particles with size
>50 µm, can be transported for days over distances of several
thousands of km (Pitty, 1968; Franzén, 1989; Betzer et al., 1988).
Experimental observations have also shown that during intense dust events a
significant amount of large particles (>30 µm) are transported
for more than 12 h (Ryder et al., 2013), and fast gravitational settling
seems to be prevented by atmospheric processes involving upward air movement
due to solar heating of the dust and/or intermittent turbulence (Maring et
al., 2003). Underestimating the contribution of large, long-range transported
desert dust particles (geometric diameter >10 µm), has a direct
impact in the outcome of radiative forcing models as well as in the
estimation of global dust emissions (Kok, 2011). Models would overestimate
cooling by fine particles due to their scattering of solar radiation,
neglecting the fact that coarser particles can induce a net warming by
absorbing both solar and thermal radiation (Otto et al., 2007; Kok et al.,
2017). Also, most global circulation models are tuned to match radiative
measurements, so that an overestimation of the radiative cooling induced by
clay-sized particles will be compensated by a reduction in the modeled
quantity of emitted dust (Kok, 2011). Recent modeling results indicate that a
positive direct radiative forcing at the top-of-the-atmosphere (TOA), leading
to global warming, can be achieved when realistically considering the amount
of very coarse (geometric diameter >10 µm) mineral dust particles
(Kok et al., 2017), which are quite abundant close to source areas and, as
shown by our results, can reach far downwind regions during extreme dust
events. It is thus very likely that coarse desert dust would have a net
positive radiative forcing close to source areas and over continental land.
Although the residence time of such larger particles is only a few tens of
hours, their persistent emission and transport may have a significant
radiative effect not only locally or regionally, but also globally. This is
particularly relevant under a global warming scenario resulting in increased
desert dust strength as predicted by recent modeling results (Kok et al.,
2018).
It could be argued that the studied extreme Saharan dust event is a rare one,
whose time span is limited and, therefore, its impact on the direct radiative
forcing should be also limited. However, these extreme events are recurrent,
and typically take place in southern Europe and the Mediterranean area on a
yearly basis (e.g., Avila et el., 2007; Cabello et al., 2012; Titos et al.,
2017). This suggests that extreme dust events can have an impact on the
atmospheric radiative budget regionally and, in the long term, even globally.
Health hazard
Most studies on the impact of Saharan mineral dust on human health have
focused on its short-term effects. In particular, a positive relationship
between hospital incidences and mortality, and desert dust outbreaks have
been established, especially for the case of PM10 (Perez et al.,
2008). Nevertheless, both a negative and a positive correlation between desert
dust PM10 and mortality has been established (Karanasiou et al.,
2012; Zhang et al., 2016). However, although there is a current lack of
detailed knowledge, recent research shows that there is a clear positive
correlation between exposure to natural mineral dust PM2.5 and
human mortality associated with respiratory and cardiovascular health issues
(Zhang et al., 2016). Modeling of the impact of dust PM2.5
estimates that the global fraction of cardiopulmonary deaths caused by desert
dust aerosols is ∼ 1.8 %, this value being ∼ 15–50 % in
countries of the so-called “dust belt” (Giannadaki et al., 2014). These
studies suggest that the clay fraction of desert dust can be the one that has
the most deleterious effects in the short term.
Little is known, however, about the long-term health effects of desert dust
exposure and inhalation. This could be the case of the potential fibrogenic
and carcinogenic risk of such inhalable mineral particles. Such a potential
long-term health effect of desert dust has been, however, generally ignored.
One exception is the study by Giannadaki et al. (2014) who found a link
between cardiovascular and lung cancer death with desert dust
PM2.5. Nonetheless, the authors ignored the role of larger
particles and the actual mineralogy of dust was not taken into account for
health-risk evaluation. Overall, the role of specific minerals in desert
dust, particularly the abundant clay minerals, and among them, the fibrous
clays, was not considered. Moreover, the possible link between such
potentially carcinogenic fibrous minerals and Fe, an element which has been
associated with increased risk for cancer development following exposure to
fibrous minerals due to its capacity to generate free radicals via the Fenton
reaction (Nolan et al., 1991; Ghio et al., 2004), was not evaluated.
Several studies have focused on the evaluation of the potential health risk
of clay minerals and associated phases present in dust (Plumlee et al.,
2006). For instance, kaolinite has been reported to be a potential
respiratory hazard. However, quartz, typically associated with clay minerals
such as kaolinite, seems to play an overruling role in the respiratory
illnesses associated with clay dust inhalation (Carretero et al., 2006).
Indeed, silica dust inhalation has been shown to be fibrogenic and
carcinogenic (Ding et al., 2002). Palygorskite, which is the most abundant
clay mineral in the studied Saharan dust samples, is a non-regulated fibrous
mineral reported to be carcinogenic and cytotoxic especially if fiber length
is over 5 µm (Rödelsperger et al., 1987), even if its content
is less than 1 % in mineral dust (Nolan et al., 1991). There are,
however, conflicting results regarding the health effects of palygorskite.
While a few studies have shown no toxicity to human embryonic intestinal
cells or low toxicity to rat pleural mesothelial cells, palygorskite has been
reported to induce hemolysis and cytotoxicity in mouse, rat, and rabbit
macrophages, and bovine and human endothelial cells (see Larson et al., 2016,
and references therein). Moreover, inhalation tests in rats have shown that
large palygorskite fibers can induce bronchoalveolar hyperplasia, alveolar
tumors, and mesothelioma (Donaldson and Borm, 2006). Iron presence either in
the mineral structure or elsewhere (adsorbed colloidal particles) seems to
enhance its carcinogenic potential (Nolan et al., 1991). This is highly
relevant for the palygorskite fibers in Saharan dust deposited in the Iberian
Peninsula and studied here, as well as for all the southern and eastern
European areas where abundant palygorkite in desert dust has been detected.
Importantly, in addition to their abundance in Saharan dust, as shown here,
and their high crystallinity and fiber length over several µm, they
include structural iron and are associated with abundant poorly crystalline
(colloidal) iron oxyhydroxides, typically attached to the clay mineral
surfaces (Fig. S8). Note, however, that not all palygorskite fibers are
equally hazardous. Poorly crystalline, small fibers (less than 1 µm
long) have been shown to be non-carcinogenic (Larson et al., 2016).
It could be argued that such an extreme event as the one studied here is rare
and of a short time span, and that therefore its health impact should not be
significant. Despite its short time span, intense dust events have been
demonstrated to have a direct impact on patient hospitalization and death
rates (Perez et al. 2008; Karanasiou et al., 2012). We thus want to stress
that recurrent extreme events such as the one studied here and those taking
place in the Mediterranean area in an almost yearly basis over the last
years/decades, plus the continuous high dust loads in Northern Africa, can
indeed have a significant short- and long-term health impact. Due to its
relevance, we would like to conclude pointing out that the possible link
between palygorskite fibers/iron-rich nanoparticles in desert dust and their
long-term health effects should be the focus of further research.