5.1.1 Processes affecting O and S isotopic compositions

Firstly, gypsum precipitation would fractionate both O and S isotopes following a slope of 0.67±0.02 (Fig. 4) when considering fractionation factors for ¹⁸O∕¹⁶O between the dissolved sulfate and the gypsum of ∼1.002 or 1.0036 (experimental and natural values, respectively) (Lloyd, 1968). For ³⁴S∕³²S, the ranges would be between 1.000 and 1.0024 (Ault and Kulp, 1959; Nielsen, 1974; Raab and Spiro, 1991; Thode et al., 1961), and a Rayleigh-type process in which black crusts represents the cumulated (precipitated) product at different residual fraction F of dissolved sulfates that are leached. However, the slope defined by the samples is steeper, ∼1.52 (R²=0.58), implying that the gypsum precipitation is not the main mechanism driving δ³⁴S and δ¹⁸O variations in black crusts. Another process that could affect δ³⁴S–δ¹⁸O values is the partial oxidation of SO₂ by different oxidants (e.g., O₂ TMI, H₂O₂, O₃, OH). Using the fractionation factors ${}^{\mathrm{34}}{\mathit{\alpha }}_{{\mathrm{SO}}_{\mathrm{4}}-{\mathrm{SO}}_{\mathrm{2}}}$ obtained experimentally at 19 ^∘C by Harris et al. (2012a) for each oxidant (${}^{\mathrm{34}}{\mathit{\alpha }}_{{\mathrm{SO}}_{\mathrm{4}}-{\mathrm{SO}}_{\mathrm{2}}}$ (OH) $=\mathrm{1.0113}\pm \mathrm{0.0024}$; ${}^{\mathrm{34}}{\mathit{\alpha }}_{{\mathrm{SO}}_{\mathrm{4}}-{\mathrm{SO}}_{\mathrm{2}}}$ (H₂O₂) $=\mathrm{1.0151}\pm \mathrm{0.0013}$; ${}^{\mathrm{34}}{\mathit{\alpha }}_{{\mathrm{SO}}_{\mathrm{4}}-{\mathrm{SO}}_{\mathrm{2}}}$ (O₂ TMI) $=\mathrm{0.9894}\pm \mathrm{0.0043}$; ${}^{\mathrm{34}}{\mathit{\alpha }}_{{\mathrm{SO}}_{\mathrm{4}}-{\mathrm{SO}}_{\mathrm{2}}}$ (O₃) $=\mathrm{1.0174}\pm \mathrm{0.0028}$) and their respective proportions from Sofen et al. (2011), usually cited in the literature for present-day atmosphere (27 % OH; 18 % O₂ TMI; 50 % H₂O₂; 5 % O₃), we calculated a global fractionation factor of 1.0097. It is worth noting that these values produce a Δ¹⁷O of 0.66 ‰ in agreement with our mean Δ¹⁷O in black crusts. Consequently, following a Rayleigh distillation model and an initial SO₂ with δ³⁴S =0 ‰, the cumulated sulfates representing the black crusts would increase up to ∼9 ‰ at maximum when < 10 % SO₂ is oxidized, which cannot explain the ∼17 ‰ δ³⁴S variation, especially since 40 % oxidized SO₂ is reported (Chin et al., 2000). To generate δ³⁴S values as high as 17 ‰, O₃ and H₂O₂ oxidation pathways should increase drastically (i.e., requiring the absence of an O₂ TMI pathway), predicting an increase of Δ¹⁷O up to ∼6.5 ‰, which is not consistent with Δ¹⁷O ∼0 ‰ associated with high δ³⁴S (Fig. 5). Therefore, SO₂ partial oxidation can explain a part of the data but not the whole isotope variations. The large δ³⁴S range could also reflect temporal variation, since in Greenland ice cores δ³⁴S was > 10 ‰ before the industrial period (Patris et al., 2002), which is dominated by SO₂ from DMS (Sofen et al., 2011) and then decreased < 4 ‰ in the 1960s, which is dominated by anthropogenic SO₂. Following this variation, black crusts on recently renovated churches should display low δ³⁴S and those renovated before the industrial period should display higher δ³⁴S. However, samples ME77-2 (δ³⁴S $=-\mathrm{0.54}$ ‰) and EV27-1 (δ³⁴S =6.60 ‰) compared to PY89-1 (δ³⁴S =0.46 ‰) gathered on churches restored after World War II and in 1772, respectively, present no significant temporal variation, which might be due to higher proportions of anthropogenic SO₂ emitted recently (0.5 Tg S yr⁻¹ before the industrial period and up to 69 Tg S yr⁻¹ for the present day; Sofen et al., 2011, and references therein). Thus, black crusts do not seem to record temporal isotopic variation, even if samples with δ³⁴S $=-\mathrm{2.66}$ ‰ and δ³⁴S =13.99 ‰ should be dated to confirm this assumption. Alternatively, with well-exposed surfaces to precipitation emphasizing wash-out and subsequent reprecipitation, black crusts could rather probe “recent” SO₂ oxidation. So far, no known processes seem to affect the isotopic compositions, which rather probe the source signatures.

Figure 5Evolution of δ³⁴S with Δ¹⁷O in black crust (BC) sulfates and sulfate aerosols (SA). The limit between mass-dependent and mass-independent fractionation (dashed line) is defined for Δ¹⁷O ∼0.65 ‰, where H₂O₂ will be the major oxidant, giving its O anomaly to sulfates (Savarino et al., 2000). When Δ¹⁷O < 0.65 ‰, black crust sulfates result from a mixing between primary sulfates (gypsum plaster and CAS and/or anthropogenic sulfur) and secondary sulfate aerosols, where SO₂ is oxidized by H₂O₂ (Δ¹⁷O =0.65 ‰), OH (Δ¹⁷O =0 ‰), or O₂ TMI (Δ¹⁷O $=-\mathrm{0.09}$ ‰) mainly (gray bar). When Δ¹⁷O > 0.65 ‰, black crust sulfates represent secondary sulfate aerosols mainly, resulting from a mixing between SO₂ oxidized by O₃ (Δ¹⁷O =8.75 ‰) and H₂O₂ (red bar). The yellow array represents gypsum plaster and CAS isotopic compositions. The dashed fields represent the sulfur isotopic composition of the two anthropogenic (An) and CAS/plaster (CAS/PL) end-members.

Download

5.1.2 Source effects

If δ³⁴S–δ¹⁸O variation reflects mixing of sources, at least two end-members are required. Determined graphically in Figs. 4 and 5, a first one would be ¹⁸O–³⁴S enriched both around 18 ‰ with a near-zero Δ¹⁷O, which in view of the sampling cross section from NW to SE and west-dominating winds could correspond to the sea-spray isotopic signature, but available data usually display δ¹⁸O of ∼9 ‰ (Markovic et al., 2016) and δ³⁴S ∼21 ‰ (Rees et al., 1978), ruling out the occurrence of sea-sprays. With the DMS produced by phytoplankton and oxidized in the atmosphere (1125 Tg S yr⁻¹) being higher than sea-salt emissions (6–12 Tg S yr⁻¹ (Alexander et al., 2005, and references therein), with δ³⁴S of 15 ‰–20 ‰ (Calhoun et al., 1991), sulfate aerosols deriving from DMS oxidation could rather represent this ¹⁸O–³⁴S-enriched end-member. However, the absence of a correlation between δ³⁴S, δ¹⁸O, Δ¹⁷O and the distance from coastline (Fig. S1) and near-zero Δ¹⁷O for high δ³⁴S values (Fig. 5) is not consistent with a significant DMS contribution, mostly oxidized by O₃ (see above; Alexander et al., 2012). Despite some isolated halite crystals observed in one sample (Fig. 3e), we conclude that, overall, marine aerosols (DMS, sea-salt sulfates) do not relate to the high δ³⁴S–δ¹⁸O end-member. Major-element contents (e.g., Na, Cl) have not been measured here even if they could further constrain and quantify the presence of marine aerosols. The structural analyses of black crusts emphasize dissolution of the underlying carbonate. Carbonate-associated sulfates (with S abundances varying between a few tens to thousands of parts per million (ppm; Kampschulte and Strauss, 2004, and references therein) would also be dissolved and reprecipitated in black crusts and may well represent the enriched δ³⁴S–δ¹⁸O end-member with near-zero Δ¹⁷O. CAS analyses from Atlantic and Pacific oceans over the last 25 Myr and in the middle Cretaceous Tethys ocean show δ³⁴S from 11 ‰ to 24 ‰, with δ¹⁸O from 5 ‰ to 21 ‰ similar to barite isotopic composition (Rennie and Turchyn, 2014; Turchyn et al., 2009). Furthermore, marine sulfates typically have Δ¹⁷O of ∼0 and > −0.2 ‰ in the geological record (Bao et al., 2008). CAS would thus perfectly match this end-member. Plaster used to seal blocks of carbonate stones is made through Lutetian gypsum dehydration and could also well represent the ¹⁸O–³⁴S-rich end-member. Indeed, Kloppmann et al. (2011) measured δ³⁴S between 12.6 ‰ and 18.3 ‰ and δ¹⁸O from 14.6 ‰ to 21.5 ‰ for mortars and plasters from French churches and castles. Thus, the surrounding plaster also matches this end-member, which is named CAS/Pl in Figs. 4 and 5.

The depleted end-member is graphically characterized by δ³⁴S < −3 ‰ with little constrained δ¹⁸O from 5 ‰ to 15 ‰ (Fig. 4, dashed box “An”) and Δ¹⁷O from ∼0 ‰ to 2.6 ‰ (Fig. 5, dashed box An). Sulfates from dissolved and oxidized sedimentary pyrites contained in the building carbonate stone are known to have δ³⁴S < −12 ‰ (since at least the last 500 Myr; Canfield, 2004). Despite a sulfide content that can vary between a few tens to thousands of parts per million (Thomazo et al., 2018), our sampled carbonate stones are very whitish, suggesting a low sulfide content. Even if it would certainly not affect the mass balance, we took into account pyrite oxidation, as other studies did on black crusts (Kramar et al., 2011; Vallet et al., 2006). With the S isotope fractionation factor during pyrite oxidation being negligible (between 0.996 and 1; Thurston et al., 2010, and references therein) compared to O isotopes, we modeled the δ¹⁸O variation according to a Rayleigh distillation to represent the sulfide oxidation, commonly occurring via O₂+H₂O at the atmosphere–carbonate building stone interface. With an initial δ¹⁸O of $\sim -\mathrm{6}$ ‰ of rainwater in the Paris Basin and a mean ¹⁸α_{water-sulfate} of 1.010 (Gomes and Johnston, 2017), sulfates from pyrite oxidation would have δ¹⁸O of ∼4 ‰ and as low as −6 ‰ if water would be in limited amounts (i.e., residual fraction of water F∼0). Very recently, pyrite oxidation was hypothesized to occur via O₃ (Hemingway et al., 2019), which would lead to positive Δ¹⁷O of sulfates with low δ³⁴S, explaining the depleted end-member. However, our data are strikingly higher than for black crusts from Ljubljana (Slovenia; Kramar et al., 2011), which show δ³⁴S as low as −20 ‰ and δ¹⁸O between −2 ‰ and 5 ‰ (Fig. 4) that would be typical for pyrite oxidation. Besides, there is so far no evidence for a higher oxidation flux of pyrite via O₃ than major constituents as H₂O and O₂. This means that another source should have negative δ³⁴S. Anthropogenic sulfur represents ∼60 % of the total sulfur released worldwide and includes primary sulfates as oil, coal, and biomass combustion products as well as SO₂ emission that can be oxidized into secondary sulfates. When considering coal and oil combustion, δ³⁴S can vary largely between −30 ‰ and 32 ‰ (e.g., Faure, 1986). More locally, a recent study reported a narrow range from −0.57 ‰ to 11.33 ‰ for sulfur emitted by transport and industries in Paris (Au Yang et al., 2020). Lee et al. (2002) carried out vegetation and diesel combustion experiments, resulting in δ³⁴S and δ¹⁸O values between 9.55 ‰ and 16.42 ‰ and between 5.5 and 10.5 ‰, respectively, with Δ¹⁷O of ∼0 ‰, forming primary sulfates without mass-independent signatures. Because sulfate aerosols can be either primary or secondary with various SO₂ oxidation pathways having distinct δ¹⁸O–Δ¹⁷O values and O-fractionation factors, atmospheric aerosols would result in variable δ¹⁸O–Δ¹⁷O values. Therefore, the depleted end-member with δ³⁴S < −3 ‰ and 0 < Δ¹⁷O < 0.65 ‰ could be typified by primary anthropogenic sulfate aerosols, or SO₂ oxidized by OH or O₂ TMI, and/or a subtle mixing of oxidation pathways to yield near-zero Δ¹⁷O, whereas samples with Δ¹⁷O > 0.65 ‰ rather point to a significant anthropogenic SO₂ fraction oxidized by O₃ and H₂O₂ or by O₃ and to a lesser extent by O₂ TMI and OH, depending on the water pH (Lee and Thiemens, 2001), corresponding to secondary sulfate aerosols (named An in Figs. 4 and 5). As the distinction between primary and secondary sulfate aerosols having near-zero Δ¹⁷O is not possible, we assume a mixing with only two end-members, CAS/Plaster and anthropogenic sulfur (primary and secondary). Furthermore, in view of the O isotope variability caused by the oxidation, mixing proportions were calculated based only on δ³⁴S values. We chose the end-members graphically and in agreement with the literature, i.e., a CAS/PL δ³⁴S value of 18 ‰, in the range from 11 ‰ to 24 ‰ (Kloppmann et al., 2011; Rennie and Turchyn, 2014; Turchyn et al., 2009) (Fig. 4) and an An δ³⁴S value of −3 ‰, similar to Montana et al. (2008) as well the closest to sulfates measured in Paris. CAS/PL proportions range from 2 % to 81 % with an average of ∼32 %. With an extreme δ³⁴S of −10 ‰ for the An end-member, encompassing black crusts from Antwerp, the CAS/PL proportion averaged 49 %. This highlights that host-rock sulfate is on average not the main S provider, and that black crusts record atmospheric sulfate aerosols. Excluding the most “contaminated” samples by CAS/PL and assuming that those having Δ¹⁷O > 0.65 ‰ obviously represent SO₂ oxidized by O₃ and H₂O₂, the minimum proportion of MIF-bearing sulfates, and hence secondary sulfates, can be estimated at ∼63 %, which is close to estimations by Lee and Thiemens (2001) and Sofen et al. (2011). In summary, black crusts sample significant amounts of atmospheric SO₂ and complement existing sampling such as aerosols, which allow us to address the origin of the Δ³³S anomaly.

5.2.1 Processes implicated in black crust formation

Δ³³S–Δ³⁶S values recorded by black crust sulfates range from −0.34 ‰ to 0.00±0.01 ‰ for Δ³³S and from −0.76 ‰ to $-\mathrm{0.22}\pm \mathrm{0.20}$ ‰ (2σ) for Δ³⁶S (Table 2). These values are quite unusual compared with anthropogenic and natural aerosols. As illustrated by Fig. 6, black crust sulfate Δ³³S values are all negative, and it is worth noting that this depletion occurs with near-constant Δ³⁶S values. This is somewhat distinct from most aerosols, which display almost exclusively positive Δ³³S values up to ∼0.5 ‰ and both positive and negative Δ³⁶S values (Au Yang et al., 2019; Guo et al., 2010; Lin et al., 2018b; Romero and Thiemens, 2003; Shaheen et al., 2014). So far, the only negative Δ³³S values down to −0.6 ‰ were measured in sulfate aerosols from Beijing (China) during winter months (Han et al., 2017) (no Δ³⁶S values provided), and these values were assumed to result from incomplete combustion of coal. This assumption ultimately relies on the work of Lee et al. (2002), which showed that primary anthropogenic aerosols formed by high-temperature combustion (e.g., diesel) result in near-zero Δ³³S–Δ³⁶S values, whereas those formed by low-temperature combustion (e.g., biomass burning) result in Δ³³S down to −0.2 ‰ and Δ³⁶S values varying between −1.9 ‰ and 0.2 ‰ (data recalculated with ³⁶β=1.9). Negative Δ³⁶S values well correlated with biomass burning proxies are also reported in East China (Lin et al., 2018b), although Δ³³S was ∼0 ‰. As in many other cities, Paris has long been affected by coal and wood burning, we can hypothesize that Δ³³S–Δ³⁶S variations result from high- and/or low-temperature combustion processes. Some black crust sulfates with near-zero Δ³³S-Δ³⁶S values could result from high-temperature combustion, but this would not explain negative Δ³³S–Δ³⁶S values. Furthermore, according to Lin et al. (2018b), low-temperature combustion would preferentially fractionate ³⁶S over ³³S, which should result in a steep slope in Δ³³S–Δ³⁶S space. The trend defined by our black crust samples shows higher ³³S fractionation than ³⁶S with Δ³³S values lower than those obtained by available low-temperature combustion experiments (< −0.2 ‰; Lee et al., 2002) and with Δ³⁶S values in the range of aerosols. Furthermore, no Δ³³S evolution is observed in black crusts sampled on churches with different ages of renovation (see Sect. 5.1; ME77-2 Δ³³S $=-\mathrm{0.21}$ ‰; EV27-1 Δ³³S $=-\mathrm{0.05}$ ‰; and PY89-1 Δ³³S $=-\mathrm{0.21}$ ‰), whereas we would expect a Δ³³S increase in black crusts from −0.2 ‰ to 0 ‰ due to the reduction of sulfur emission from low-temperature replaced by high-temperature combustion processes. Therefore, available data highlight that neither high- nor low-temperature combustion processes are responsible for low Δ³³S measured in black crusts.

Figure 6Δ³³S and Δ³⁶S of the black crust, compared to sulfates formed by different oxidation pathways and by a mixing of them in the proportions estimated by Sofen et al. (2011). We took ${}^{\mathrm{33}}{\mathit{\beta }}_{{\mathrm{H}}_{\mathrm{2}}{\mathrm{O}}_{\mathrm{2}}/{\mathrm{O}}_{\mathrm{3}}}=\mathrm{0.511}$, ³³β_OH=0.503, ${}^{\mathrm{33}}{\mathit{\beta }}_{{\mathrm{O}}_{\mathrm{2}}\text{ TMI}}=\mathrm{0.498}$ (for T < 20 ^∘C), ${}^{\mathrm{33}}{\mathit{\beta }}_{{\mathrm{O}}_{\mathrm{2}}\text{ TMI}}=\mathrm{0.547}$ (for T > 20 ^∘C), ${}^{\mathrm{33}}{\mathit{\beta }}_{{\mathrm{NO}}_{\mathrm{2}}}=\mathrm{0.514}$ and ${}^{\mathrm{36}}{\mathit{\beta }}_{{\mathrm{H}}_{\mathrm{2}}{\mathrm{O}}_{\mathrm{2}}/{\mathrm{O}}_{\mathrm{3}}}=\mathrm{1.82}$, ³⁶β_OH=1.97, ${}^{\mathrm{36}}{\mathit{\beta }}_{\mathrm{O}+\mathrm{2}\text{ TMI}}=\mathrm{1.98}$ (for T < 20 and T > 20 ^∘C), and ${}^{\mathrm{36}}{\mathit{\beta }}_{{\mathrm{NO}}_{\mathrm{2}}}=\mathrm{1.90}$ (Au Yang et al., 2018; Harris et al., 2013b). As ${}^{\mathrm{33}}{\mathit{\beta }}_{{\mathrm{O}}_{\mathrm{3}}}$ and ${}^{\mathrm{36}}{\mathit{\beta }}_{{\mathrm{O}}_{\mathrm{3}}}$ are unknown, we modified the proportions by Sofen et al. (2011) as follows: 27 % OH, 18 % O₂ TMI, 55 % H₂O₂, and 0 % O₃. The urban aerosol isotopic compositions are a compilation from Au Yang et al. (2019); Guo et al. (2010); Lin et al. (2018b); Romero and Thiemens (2003); and Shaheen et al. (2014), while the combustion process reflects samples from Lee et al. (2002). Modeled Δ³³S–Δ³⁶S values of cumulated black crusts (BC) sulfates – formed by SO₂ wet and dry deposition with a magnetic isotope effect (MIE; ³³S depletion compared to initial SO₂ with constant negative Δ³⁶S) and of cumulated secondary aerosols formed by SO₂ oxidation by O₃, O₂ TMI, OH, and H₂O₂ (³³S enrichment compared to initial SO₂) from an initial SO₂ with Δ³³S–Δ³⁶S =0 ‰ – are reported with corresponding β exponents (see Sect. 5.2.2 for model explanation). Residual SO₂ and global cumulated BC plus secondary aerosol isotopic compositions were not reported for better readability. Percentages indicate the fraction of produced cumulated BC and secondary aerosols.

Download

With a part of black crust sulfates being atmospheric in origin, isotopic effects during SO₂ oxidation could be responsible for Δ³³S–Δ³⁶S variations. To better address this issue, we calculated the Δ³³S–Δ³⁶S values of sulfates predicted by each of the main SO₂ oxidation pathways and by a mixing of them in the proportions given by Sofen et al. (2011). We used ³³β and ³⁶β values determined by experiments of SO₂ oxidation by O₂ TMI, H₂O₂, O₃, OH (Harris et al., 2013b), and values cited in Au Yang et al. (2018) (see caption text). We also used NO₂ values (Au Yang et al., 2018) and T-dependent equations determined by Harris et al. (2013b) to calculate each ³⁴α with initial sulfur dioxide Δ³³S and Δ³⁶S of 0 ‰ (Lin et al., 2018b). As mentioned earlier (Au Yang et al., 2018; Harris et al., 2013b), none of these models can account for anomalous Δ³³S–Δ³⁶S values in either aerosols or in black crusts (Fig. 6). Although oxidation with O₂ TMI at T=50 ^∘C could produce negative Δ³³S down to −0.37 ‰, which would account for the lowest Δ³³S observed in black crusts, this oxidation pathway would also produce larger Δ³⁶S down to −1.50 ‰ at odds with the Δ³⁶S reported in the black crust. Their potential combination cannot account for sulfate aerosol data from the literature (Au Yang et al., 2019; Guo et al., 2010; Lin et al., 2018b; Romero and Thiemens, 2003; Shaheen et al., 2014) or for the black crust as it would result in slightly negative Δ³³S–Δ³⁶S values that could not explain the Δ³³S as low as −0.34 ‰ (yellow frames in Fig. 6). Available literature data are therefore not consistent with the anomalous Δ³³S–Δ³⁶S values recorded in black crust sulfates.

Mass-dependent processes can also result in small Δ³³S–Δ³⁶S variations, depending on the magnitude of the ³⁴S fractionation (Ono et al., 2006a). As mentioned in Sect. 5.1.2, a mixing between a ^33,34S-depleted end-member (An) consisting of anthropogenic sulfur (δ³⁴S $=-\mathrm{3}$ ‰, Δ³³S =0 ‰) and a ^33,34S-enriched sulfates end-member (CAS/PL) from plaster or CAS (δ³⁴S =18 ‰, Δ³³S =0 ‰) would result in small Δ³³S values of −0.01 ‰ for 50 % mixing, which is far from the maximum measured Δ³³S of $\sim -\mathrm{0.34}$. Moreover, the slope between Δ³³S–Δ³⁶S would be about −7, which is at odds with our observations. Therefore, we conclude that mixing cannot account for the black crust Δ³³S–Δ³⁶S variations.

5.2.2 A new oxidation pathway implying magnetic isotope effect

Several studies proposed that positive Δ³³S measured in sulfate aerosols, with Δ³³S up to 0.5 ‰, from, for example, East China and California could result from stratospheric fallout of SO₂ (with Δ³³S potentially up to 10 ‰ higher; Ono et al., 2013), which underwent UV photolysis by short-wavelength radiation (Romero and Thiemens, 2003; Lin et al., 2018a, b). This suggestion primarily relies on the similarities between Δ³³S–Δ³⁶S values of sulfate aerosols and laboratory experiments of SO₂ photolysis conducted at different wavelengths (Romero and Thiemens, 2003) and on the correlation between ³⁵S-specific activity and Δ³³S values (Lin et al., 2018b). However, these studies never addressed the absence of the complementary negative Δ³³S reservoir, which is required to balance the positive Δ³³S reservoir (see Au Yang et al., 2019). In this respect, it is worth mentioning that volcanic and stratospheric aerosols trapped in Antarctic ice cores (see Gautier et al., 2018, and references therein) show both positive Δ³³S (up to ∼2 ‰) and complementary negative Δ³³S values (down to −1 ‰) and weighed average Δ³³S ≠0 ‰ explained by prior partial deposition. Stratospheric fluxes are actually too low to account for Δ³³S > 0.1 ‰ (Lin et al., 2016; Au Yang et al., 2019). Accordingly, some other authors rather tried to explain the positive anomalies of most aerosols with “tropospheric” chemical reactions, which are SO₂ oxidation by the main oxidant including NO₂, H₂O₂, OH, O₃, and O₂ TMI, but experimental data result in a maximum Δ³³S values of ∼0.2 ‰ for all studied reactions (Au Yang et al., 2018; Harris et al., 2013b). Isotope effects associated with SO₂ oxidation by minor species, such as Criegee radicals, remains to be investigated (Au Yang et al., 2018). In summary, whatever the stratospheric vs. tropospheric origin of positive Δ³³S values recorded by most aerosols, there is a ³³S isotope imbalance and a missing reservoir with negative Δ³³S that must exist. Han et al. (2017) reported Δ³³S values down to −0.6 ‰ in sulfate aerosols from Beijing. As discussed above, the authors' suggestion calling for low-temperature combustion is little supported by available data, and clearly the very restricted location and time interval, over a month, where these anomalies occurred cannot counterbalance, both spatially and temporally, the common positive Δ³³S values of most aerosols; the missing reaction or reservoir requires, instead, to be ubiquitous worldwide.

In this study, black crust sulfates display negative Δ³³S values (from ∼0 ‰ down to −0.34 ‰). These values are certainly produced by tropospheric chemical reactions. Otherwise, they would have, according to the stratospheric origin model, the same sign as those measured among aerosols. Furthermore, the chemical reactions (or reaction) involved in the formation of black crusts must be distinct compared to those leading to the formation of tropospheric aerosols. As developed thoroughly, black crusts could well represent the missing sulfur reservoir.

An additional observation is that negative Δ³³S values occur with near-constant Δ³⁶S (from −0.76 ‰ to $-\mathrm{0.22}\pm \mathrm{0.20}$ ‰; Fig. 6). This signature is typical of the magnetic isotope effect (MIE), which involve a radical pair, where coupling between the nuclear magnetic moment of the nucleus of odd isotopes and the electron occurs, allowing for electron spin transition from singlet to triplet (or vice versa) (Buchachenko et al., 1976). This leads to distinct half-lives between odd and even isotopes, resulting in specific odd over even isotope enrichment (or depletion). MIE has been so far reported for various reactions occurring on a surface (Buchachenko, 2001, 2000; Turro, 1983) such as sulfate thermochemical reduction (Oduro et al., 2011) or Fe reduction in magneto-tactic bacteria (Amor et al., 2016), which are the most geologically relevant. It is worth pointing out that MIE could also be responsible for positive Δ¹⁷O values measured in black crusts, i.e., as opposed to the Δ¹⁷O anomaly being inherited from SO₂ oxidants. However, Lee et al. (2002) also measured the O multi-isotope compositions of sulfate aerosols (i.e., from the atmosphere as opposed to reaction on a solid substrate) from Paris and obtained Δ¹⁷O =0.2 and 0.8 ‰ for the Paris highway and in the 13th zone, respectively, which is in good agreement with our three samples collected in Paris (from 0.17 ‰ to 0.89 ‰). Thus, this is consistent with black crust formation recording mostly an atmospheric signal and no significant magnetic isotope effect on Δ¹⁷O.

The magnetic effect could occur on a surface such as on mineral dust suspended in the atmosphere during aerosol formation, leading to residual ³³S-depleted atmospheric SO₂ from which black crusts would subsequently form. This model would, however, predict that some sulfate aerosols subsequently formed to display negative Δ³³S values: such values are extremely uncommon, being primarily restricted to the Beijing winter months (Han et al., 2017). Instead, the magnetic effect could occur during black crust formation, on the carbonate building stone, leading to residual ³³S-enriched atmospheric SO₂ from which tropospheric aerosols would subsequently be formed, which is consistent with available observations. This model would, however, predict some black crust subsequently formed to display positive Δ³³S: such values have not been found yet and this may well reflect sample bias, with our data being the first reported for such samples. Both scenarios imply nonzero Δ³³S values of residual atmospheric SO₂, which contrast with the data by Lin et al. (2018b) showing Δ³³S of ∼0 ‰ (n=5, Δ³³S varying from −0.04 ‰ to 0.01±0.01 ‰). Given that, in the study of Lin et al. (2018b), SO₂ was sampled close to the third largest Chinese megacity, such nonzero Δ³³S values may thus be rather symptomatic of emitted (i.e., anthropogenic) SO₂ rather than residual or background (i.e., after significant black crust and aerosols formation). SO₂ in the Paris Basin still has to be measured to confirm this assumption, but, so far, this could be consistent with the interpretation that nonzero Δ³³S values of residual or background atmospheric SO₂ are erased by anthropogenic SO₂ having zero Δ³³S values (Au Yang et al., 2019) moving towards the local source(s) of anthropogenic SO₂.

Figure 7Modeled δ³⁴S and Δ³³S values of black crust (BC) sulfates (instantaneous and cumulated) formed by SO₂ wet and dry deposition with a MIE (³³S depletion compared to initial SO₂) and of secondary aerosols (instantaneous and cumulated) formed by SO₂ oxidation by O₃, O₂ TMI, OH, and H₂O₂ (³³S enrichment compared to initial SO₂) from an initial SO₂ with δ³⁴S =1 ‰ and Δ³³S =0 ‰ (red point). ³³S enrichment of residual SO₂ and global cumulated BC plus secondary aerosol isotopic compositions are also reported. Percentages indicate the fraction of produced cumulated BC and secondary aerosols (see Sect. 5.2.2 for model explanation). Black crust sulfate isotopic compositions (dark gray points) can be explained by a mixing (red triangle) between sulfates from CAS/plaster (dashed square, δ³⁴S =18 ‰ and Δ³³S =0 ‰; see Sect. 5.1), primary anthropogenic sulfates (dashed square, δ³⁴S $=-\mathrm{3}$ ‰ and Δ³³S =0 ‰), and sulfates formed by wet and dry deposition of SO₂ undergoing a MIE and oxidized SO₂ forming secondary aerosols (cumulated BC sulfates and cumulated BC + secondary aerosols). Urban aerosol (light gray points) isotopic compositions are a compilation from Au Yang et al. (2019); Guo et al. (2010); Lin et al. (2018b); Romero and Thiemens (2003); and Shaheen et al. (2014).

Download

In the absence of additional observations, proposing a chemical reaction, and hence a radical pair that breaks and recombines, would be very speculative, but our data clearly point towards the occurrence of a magnetic effect during the formation of black crusts, involving ubiquitous heterogeneous chemical reactions. This is supported by previous recognition of sulfur radicals such as ${\mathrm{SO}}_x^{-}$ (Herrmann, 2003) or S–S (see Babikov, 2017. But note that their Δ³⁶S∕Δ³³S slope is distinct from ours). Clearly, the reaction does not occur after sulfate formation such as during dissolution–precipitation mechanisms, which do not involve any radical species. As mentioned above, magneto-tactic bacteria can produce MIE when reducing Fe (Amor et al., 2016). With microbial activity being sometimes present on black crusts (Gaylarde et al., 2007; Sáiz-Jiménez, 1995; Scheerer et al., 2009; Schiavon, 2002; Tiano, 2002), the involvement of microorganisms, affecting only the sulfur isotopes as the most negative Δ³³S, does not correspond to the most negative Δ¹⁷O, which represents another possibility to investigate. Another implication that can be tested in future work is that the kinetics of heterogeneous reactions leading to sulfate and black crust formation should be comparable or faster than those leading to aerosol formation. So far, Li et al. (2006) showed comparable loss of atmospheric SO₂ by heterogeneous oxidation on calcium carbonate substrates and by gas-phase oxidation. Our conclusions show a strong analogy with the model of Au Yang et al. (2019), who suggest that SO₂ photooxidation on mineral dust could form sulfate aerosols depleted in ³³S that would then be deposited. The residual SO₂ would be subsequently enriched in ³³S, then be oxidized by common O₃, H₂O₂, O₂, or OH oxidants. Their ³³S-depletion mechanism was not further constrained, except that it was speculated to be photochemical in origin.

If correct, this view requires reassessing the overall S isotope fractionation during SO₂ atmospheric reaction. So far, previous studies assumed that the overall sulfur isotopic fractionation between the wet and dry deposit and oxidized SO₂ was equal to 1 (i.e., no isotope effect), but negative Δ³³S in black crusts is inconsistent with such an assumption.

Starting with an SO₂ Δ³³S value of 0 ‰ (Au Yang et al., 2018; Lin et al., 2018b) and forming oxidized (sampled by secondary aerosols) and wet and dry deposit (sampled by black crusts) reservoirs with Δ³³S values up to 0.50 ‰ down to −0.34 ‰, respectively, mass balance implies that SO₂ dry and wet depositions and secondary sulfate aerosols represent ∼60 % and 40 %, respectively. This is in good agreement with proportions obtained by Chin et al. (2000) and quoted by Harris et al. (2013b). Therefore, we conclude that MIE happening during SO₂ dry and wet depositions could be a viable mechanism responsible for ³³S enrichment of secondary sulfate aerosols and that black crusts could represent the ³³S-negative complementary reservoir.

Figure 8Scheme summarizing the sulfur sources and processes that lead to black crust formation. Sulfur dioxide releases by anthropogenic activities can either be oxidized in the atmosphere by H₂O₂, O₃, OH, O₂ TMI and formed secondary sulfate aerosols that will react with the carbonate building stone to produce ³³S-enriched black crust sulfates or be deposited, as dry and wet deposit, on the carbonate substrate where its oxidation into ${\mathrm{SO}}_x^{-}$ then sulfates through MIE will produce ³³S-depleted black crust sulfates and a ³³S-enriched residual SO₂ (source 3). Primary sulfates emitted by anthropogenic activities (source 2) or carbonate-associated sulfates and/or plaster of the host rock (source 1) are also likely sources contributing to black crust formation.

Download

In order to better apprehend the aerosols and black crust complementarity, we modeled the S isotopic fractionation of both black crusts and aerosols during SO₂ oxidation (Figs. 6 and 7). We assumed a Rayleigh distillation model to represent the atmosphere–building-stone interface open system. The global fractionation factor between residual SO₂ and oxidized (secondary aerosols) plus deposited (black crusts) SO₂ is defined as

with A and B being the proportions of SO₂ deposited and oxidized, being equal to 60 % and 40 %, respectively. This allows us to deduce the ${}^{\mathrm{33},\mathrm{34},\mathrm{36}}{\mathit{\alpha }}_{\text{BC-}{\mathrm{SO}}_{\mathrm{2}}}$ and the associated ^33,36β factors. A δ³⁴S of 1 ‰ for the initial SO₂ was considered to obtain black crusts of at least −3 ‰ (see Sect. 5.1.2) and Δ³³S–Δ³⁶S =0 ‰; ${}^{\mathrm{34}}{\mathit{\alpha }}_{\text{aerosols-}{\mathrm{SO}}_{\mathrm{2}}}$ was taken as 1.0097 as calculated using the different oxidation channel proportions of Sofen et al. (2011). With the oxidation being mass dependent, we chose ${}^{\mathrm{33},\mathrm{36}}{\mathit{\beta }}_{\text{aerosols-}{\mathrm{SO}}_{\mathrm{2}}}$ values of 0.515 and 1.9, respectively (Harris et al., 2012b). The best fit is obtained for ${}^{\mathrm{34}}{\mathit{\alpha }}_{\text{BC-}{\mathrm{SO}}_{\mathrm{2}}}=\mathrm{0.9985}$, ${}^{\mathrm{33}}{\mathit{\alpha }}_{\text{BC-}{\mathrm{SO}}_{\mathrm{2}}}=\mathrm{0.9986}$, and ${}^{\mathrm{36}}{\mathit{\alpha }}_{\text{BC-}{\mathrm{SO}}_{\mathrm{2}}}=\mathrm{0.9972}$ with ${}^{\mathrm{33},\mathrm{36}}\mathit{\beta }=\mathrm{0.9}$ and 1.9, respectively. The ³³S enrichment in secondary sulfate aerosols is well represented by this parameterization (instantaneous and cumulated products; Figs. 6 and 7). The concomitant ³³S depletion in modeled cumulated deposit is also well represented. The ³³S isotopic fractionation occurring during the MIE is higher than the one observed in black crusts. To a first-order approximation, this model works, predicting the total cumulated products of black crusts and aerosols to have Δ³³S values of −0.23 ‰ and 0.35 ‰, respectively. The match is not perfect, which does not entirely capture the black crust isotopic compositions. But remember that black crusts are produced from anthropogenic Parisian SO₂, whereas aerosols formed in other locations possibly formed from distinct anthropogenic SO₂ δ³⁴S values. In addition, we are aware that our model strongly depends on oxidation pathways estimated by Sofen et al. (2011), which vary spatially and temporally. The main weakness is the poorly constrained estimate of intrinsic S-bearing compounds (CAS/plaster end-member) in the host rock as well as sulfate aerosols (Δ³³S > 0 ‰), which dilutes the ³³S anomaly, lowering the overall black crust Δ³³S. Ultimately, black crusts result mainly from the deposition followed by oxidation of SO₂ on the building stone rather than aerosol accumulation. The δ³⁴S value of initial anthropogenic SO₂ is another poorly constrained parameter whose variability might be difficult to estimate both spatially and temporally.

In conclusion, black crusts could represent the complementary sulfur end-member to sulfate aerosols. Its fractionation factor is relatively restricted (−1.5 ‰) and is thus likely identifiable from its negative Δ³³S values. Our model is actually consistent with the assumption that the global SO₂ oxidation occurs with little fractionation (³⁴α_global=1.00298) as commonly done in the literature. Finally, Fig. 8 summarizes the different sulfur sources involved in the black crusts as well as the processes leading to their formation. Black crust isotopic compositions could thus be explained by a mixing between sulfates from CAS/plaster (δ³⁴S ∼18 ‰ and Δ³³S =0 ‰; see Sect. 5.1 and 1 in Fig. 8), primary anthropogenic sulfates (δ³⁴S $\sim -\mathrm{3}$ ‰ and Δ³³S =0 ‰; see 2 in Fig. 8), and wet and dry deposition of SO₂ undergoing MIE during its oxidation on the building stone combined with secondary aerosols (see red triangle in Fig. 7 and process 3 in Fig. 8).