the proportion of non-marine source waters (Topper et al., 2011; Roveri et al., 2014a). At first glance it seems that each MSC Stage was characterized by a well-defined range of Sr ratios (Fig. 10a inset), an observation that led Roveri et al. (2014b) to attribute a chronostratigraphic value to MSC $^{87}$Sr/$^{86}$Sr ratios. A closer look, however, shows that MSC substages are anything but homogeneous with respect to $^{87}$Sr/$^{86}$Sr ratios. At least in the marginal basins, local lithological differences in the catchments (each lithology carries a unique $^{87}$Sr/$^{86}$Sr fingerprint; see subsection 8.1.1) explain the different Sr isotopic compositions from basin to basin (see Schildgen et al., 2014; Modestou et al., 2017; Andreetto et al., 2021), therefore arguing against the use of $^{87}$Sr/$^{86}$Sr ratios for chronostratigraphic purposes in the MSC record.

Most of the data characterizing substage 3.1 (Fig. 10a) are from the Eraclea Minoa gypsum (Fig. 5h). These data define a narrow range of Sr isotope ratios between 0.708747 and 0.708793 (García-Veigas et al., 2018). Similar values were reported from both Eraclea Minoa and the nearby Siculiana Marina section (0.708710-0.708760; Keogh and Butler, 1999; Fig. 5i). The dominance of Sicily samples gives the appearance of a consistent Sr isotope signal for gypsum beds. However, data points from elsewhere (Cyprus, Manzi et al., 2016a; DSDPs 122, 371 and 372 in the Algero-Balearic Basin, ODPs 652, 653 and 654 in the Tyrrhenian Basin, DSDP 374 in the Ionian Basin; Müller et al., 1990; Müller and Mueller, 1991; Roveri et al., 2014b) display a wider range (from ~0.7087 to 0.708847; Fig. 10a) that may indicate a different hydrological regime for each basin (e.g. Müller et al., 1990; Müller and Mueller, 1991; Ryan, 2009). The one published Sr isotope value for ostracods found within one of the marl interbeds at Eraclea Minoa also has a lower value outside the typical Sicily gypsum range (Grossi et al., 2015). This suggests that a different hydrological regime may also have characterised precession minima stages of the precessional cycle.

The Sr isotope dataset for the Lago-Mare phase includes the lowest values measured on MSC sediments (~0.7085 from between gypsum VI and VII at Eraclea Minoa; Fig. 3a; Grossi et al., 2015) and the widest range of ratios spanning from 0.7085 to 0.7091, which is above coeval oceanic values (Fig. 10a). Again, the conspicuously high Sr isotope values in substage 3.2 come from two areas, the marginal basins of southern Spain (Andreetto et al., 2021 and references therein; Figs. 5a-c) and the intermediate Polemi Basin on Cyprus (McCulloch and De Deckker, 1989). The lower values are drawn from right across the intermediate-deep Mediterranean (Algero-Balearic, Sicily, Levant; Fig. 2a) and are therefore more likely to represent a Mediterranean-wide Sr isotope signal.

New Sr isotope data from Eastern Parathethys (i.e. Dacian and Caspian basins; Fig. 2b) are now available for the interval corresponding to MSC Stage 3 (inset Fig. 10a). The $^{87}$Sr/$^{86}$Sr ratios of the Dacian Basin (0.708865-0.708982; Vasiliev et al., 2010; Grothe, 2016) are slightly lower than coeval ocean water (0.709020), but much higher than coeval Mediterranean values. However, the Dacian Basin is regarded as highly restricted from the Mediterranean throughout the MSC (Vasiliev et al., 2010). By contrast, the Caspian has very low values (0.708402 to 0.708473, Grothe et al., 2020) which are thought to reflect both the very low Sr isotope ratio of the Volga river (0.708020; Vasiliev et al., 2010 and references therein) and some input from the Mediterranean (Grothe et al., 2020).

## 6.2. Sulfate isotopes

Sulfur isotopic investigations have been carried out only on sulfate minerals (gypsum and more rarely anhydrite) of the MSC Stage 3 deposits with samples drawn from both onshore intermediate sequences (Caltanissetta Basin and Cypriot basins) and deep basinal records (Sites 122, 124, 125A, 132, 134, 372, 374, 375, 376, 652, 653, 654, 968, 969, 970; Fig. 10b; Fontes et al., 1973; Pierre, 1974, 1982; Pierre and Fontes, 1978; Ricchiuto and McKenzie, 1978; Pierre and Rouchy, 1990; Blanc-Valleron et al., 1998). Because the incorporation of dissolved sulfate into gypsum produces a nearly constant fractionation of $\delta^{18}$O (+3.5‰) and

$\delta^{34}$S (+1.65‰) at earth surface temperatures (Thode and Monster, 1965; Lloyd, 1968; Warren, 2016), $\delta^{18}$O and $\delta^{34}$S isotopic values measured in gypsum should be corrected with the above mentioned fractionation factors in order to reconstruct the sulfate isotopic composition of the basin waters at the time of gypsum formation.

The deep Mediterranean samples exhibit a wide range of $\delta^{34}$SSO4, but the majority of samples display $\delta^{34}$SSO4 values between 18 and 22‰, strongly indicative of a marine origin of the sulfate forming the gypsum (Fig. 10b; Fontes et al., 1973; Pierre, 1974, 1982; Pierre and Fontes, 1978; Pierre and Rouchy, 1990; Blanc-Valleron et al., 1998). The $\delta^{34}$SSO4 values lower than marine sulfate in the dataset are generally considered to represent a greater influence of continental sulfate input to the basin (Fig. 10b; Pierre, 1974; Pierre and Fontes, 1978; Pierre and Rouchy, 1990). By contrast, the data display $\delta^{18}$Oso4 isotopic values that deviate substantially from marine $\delta^{18}$Oso4 values towards higher values (Fig. 10). This is consistent with the influence of sulfate produced by reoxidation of reduced sulfur compounds generated by microbial sulfate reduction (MSR; Kaplan and Rittenberg, 1964; Brunner and Bernasconi, 2005; Sim et al., 2011; Leavitt et al., 2013). The microbial use of SO4 leads to an equilibration of $\delta^{18}$Oso4 with ambient water oxygen, whereas the $\delta^{34}$SSO4 returns towards its initial value as a higher fraction of sulfide produced by MSR is re-oxidated. This mechanism has been suggested for Sites in the Algero-Balearic, Tyrrhenian and Ionian basins and offshore Cyprus (Pierre, 1974; Pierre and Fontes, 1978; Pierre and Rouchy, 1990). Although some authors have suggested that partial equilibration of sulfate oxygen toward $\delta^{18}$O$_{H_2O}$ values of the basin enriched in heavy oxygen isotopes by evaporation have led to an increase in $\delta^{18}$Oso4 values without significant changes in $\delta^{34}$SSO4 (Fontes et al., 1973; Pierre, 1974; Ricchiuto and McKenzie, 1978), this hypothesis seems highly unlikely as the abiotic equilibration between sulfate and water oxygen take about 20 Myr at normal marine pH (Lloyd, 1968; Longinelli and Craig, 1967; Turchyn et al., 2006). Moreover, the microbial sulfate reduction process is supported by the presence of pyrite at Sites 132, 654A and 968 (Pierre, 1982; Pierre and Rouchy, 1990; Blanc-Valleron et al., 1998) and the existence of filaments of possible microbial origin at Site 654A (Pierre and Rouchy, 1990).

The sulfate isotopic values reported by Longinelli (1979) and Pierre (1982) from the Upper Gypsum of Eraclea Minoa (Caltanissetta Basin, Sicily) are considerably more scattered than those from a recent study by García-Veigas et al., 2018; Fig. 10b). Such discrepancies are probably a consequence of different sample selection: García-Veigas et al. (2018) analyzed only pristine whitish selenite and balatino samples, while Longinelli (1979) and Pierre (1982) analyzed all types of gypsum-bearing samples such as “gypsiferous marl” and gypsum laminae intercalated in carbonate or diatomaceous intervals. These less pristine samples probably contain high quantities of $^{34}$S-depleted solid sulfides or diagenetic gypsum formed by oxidation of sulfides (see Liu et al., 2017 for more details on this process) and are therefore unlikely to be representative of the primary gypsum facies. Once these data are excluded, the Eraclea Minoa sulfate values ($\delta^{18}$Oso4 from 12.4 to 14.6‰ and $\delta^{34}$SSO4 from 21.0 to 22.3‰) suggest a marine origin of the sulfate and stable redox conditions during gypsum deposition (Fig 9.b; García-Veigas et al., 2018). Interestingly, the Eraclea Minoa sulfate values are in compliance with the isotopic values ($\delta^{18}$Oso4=15.2 to 16.8‰; $\delta^{34}$SSO4=20.4 to 21.9‰) measured by Pierre (1982) in the Polemi Basin (Cyprus).

## 6.3. Hydrogen isotopes on molecular biomarkers

From the point of view of the application of organic geochemistry proxies, the Miocene Mediterranean Basin received little attention so far, with biomarker-based proxies that have been mostly applied to (a limited number of) pre-MSC sequences (Tzanova et al., 2015; Herbert et al., 2016; Mayser et al., 2017; Natalicchio et al., 2017, 2019; Vasiliev et al., 2019) and pre-Stage 3 sedimentary records (Lower Evaporites on Sicily, Andersen et al., 2001; Vena del Gesso Basin, Sinninghe Damsté
et al., 1995 and Vasiliev et al., 2017; Levant Basin, Meilijson et al., 2019). To date, only one study analyzed Stage 3 samples (Vasiliev et al., 2017). This study used compound specific hydrogen isotope (δD) analyses, measured on both terrestrial (long chain C29 and C31n-alkanes; Sachse et al., 2006) and aquatic (alkenones; Englebrecht and Sachs, 2005) biomarkers from the gypsum beds of the Upper Gypsum at Eraclea Minoa to reconstruct the hydrological cycle during gypsum precipitation.

Both δDC29n-alkane and δDalkenones results (Fig. 10c) suggested that conditions in Sicily were significantly dryer than today, with highly enriched values of δDC29n-alkanes (up to -125‰). The δDalkenones varied between values suggesting evaporative conditions (-125‰) and values typical for present-day δDalkenones in the Mediterranean (-203‰) (Vasiliev et al., 2017).

No time-equivalent biomarker data from the open ocean settings are currently available. Instead, Vasiliev et al. (2017) compared their Mediterranean data with data from the Black Sea (DSDP 42B Hole 380 and Taman peninsula; Vasiliev et al., 2013, 2015). The Upper Gypsum δDn-alkanes were more enriched when compared to their time equivalent deposits of the DSDP 42B 380 borehole of the Black Sea (-180‰). This probably reflects the more intracontinental position of the Black Sea which commonly translates into more depleted values for δDprecipitation used by the vegetation, resulting in more depleted δDC29n-alkanes. However, there is a 30 to 40‰ enrichment relative to present in the δDn-alkanes (i.e. δDprecipitation) in both Mediterranean and Paratethys domains, indicating concurrent changes in both areas during the latest phase of the MSC.

Both the Mediterranean and Paratethyan samples contain δDalkenones with low values (~-200‰) (Fig. 10c) leading Vasiliev et al. (2017) to suggest that either the surface water from the Upper Gypsum was derived from the Black Sea, or that the Mediterranean and Paratethys were exchanging surface water during gypsum precipitation. Similarity between the relative contribution of the C37, C38 and C39 alkenones at Eraclea Minoa and one of the Black Sea samples may suggest common alkenone producers for the two areas, again supporting the idea of a Mediterranean-Paratethys connection during Stage 3 (Vasiliev et al., 2017).

A final speculative insight from this biomarker dataset is that the relative contribution of alkenones found in the Upper Gypsum of Eraclea Minoa is strikingly similar to present-day open marine samples, even though *Emiliania huxleyi*, the principal ocean alkenone producer today, did not exist in the late Miocene. Vasiliev et al. (2017) suggested that this could imply the existence of a connection to the open ocean during Upper Gypsum deposition in Sicily (i.e. throughout Stage 3; Fig. 3a).

## 6.4. Oxygen isotopes

Oxygen stable isotope data (δ¹⁸O) are available from bulk samples (Rouchy et al., 2001, 2003, 2007; Pierre et al., 2006; Cosentino et al., 2012), gypsum (Pierre and Fontes, 1978; Ricchiuto and McKenzie, 1978; Lugli et al., 2007), mollusk shells (Carnevale et al., 2008; Grunert et al., 2016) and ostracod valves (Cosentino et al., 2012; Grossi et al., 2015).

For all the sub-basins for which there is latest Messinian data (e.g. Sites 974 and 975; Eraclea Minoa section, Sicily; Aghios Stefanos section, Corfu; Kalamaki section, Zakynthos; Pissouri Basin, Cyprus; Rouchy et al., 2001, Pierre et al., 2006), each has its own range of oxygen isotopic compositions and its own degree of variability. Values from above the Miocene/Pliocene boundary regain seawater values of 0.3 to 1 ‰ (e.g. Pierre et al., 2006).

In marginal marine settings and lakes, the controls over δ¹⁸O are poorly constrained as oxygen does not respond simply to the freshwater flux, but to a combination of variables such as temperature, rainfall and evaporation (e.g. Placzek et al., 2011). Freshwater input may contribute to the signal, resulting in δ¹⁸O more negative than seawater (0.3‰ to 0.8‰ SMOW; Dettman et al., 2004), but under prevailing evaporating conditions it is likely that the δ¹⁸O will be primarily influenced by

evaporation, leading to δ¹⁸O more positive than seawater (e.g. Dettman et al., 2004), making any data very difficult to interpret. Furthermore, the lack of a unique δ¹⁸O signature for each water source makes oxygen isotopes a difficult tracer proxy to use.

## 6.5. Summary of the Stage 3 geochemical dataset

The variety of paleoenvironmental and connectivity proxies applied to MSC Stage 3 record provide valuable insights into the hydrological conditions during Stage 3. The more outstanding results from all discussed proxies are that:

1) Paleodepositional subaqueous environments where gypsum was precipitating and ostracods and biomarker-producers were thriving were strongly dominated by non-oceanic inputs;

2) an indisputable marine signal is absent and only regained above the M/P boundary.

Sulfate and oxygen isotopes are currently difficult to use for water provenance reconstruction because the non-marine sources (local and major rivers and Eastern Paratethys) that are likely to be of influence lack distinctive isotopic signatures and, especially for oxygen, respond to a combination of controls (e.g. temperature, rainfall, evaporation) with local variability. δ³⁴S₅₀ are claimed by several authors to be an evidence of the presence of an Atlantic inflow (δ³⁴S₅₀=22‰; Turchyn and Schrag, 2004) in a Mediterranean strongly affected by non-marine waters (Manzi et al., 2009, 2016a; García-Veigas et al., 2018 among others). However, the same values can be obtained by means of the recycling of PLG deposits (~23‰; Lu et al., 2001; Lugli et al., 2010; García-Veigas et al., 2018).

Similarities between the δDalkenones of the Upper Gypsum at Eraclea Minoa and coeval Black Sea sediments and δDn-alkanes similar to present-day marine settings, suggest that Eastern Paratethys and the Atlantic were simultaneously contributing to the Mediterranean hydrological budget. ⁸⁷Sr/⁸⁶Sr isotope ratios are a useful water-mass tracer because each water body carries a unique Sr isotope fingerprint (see subsection 8.1.1). Our plotting of Stage 3 ⁸⁷Sr/⁸⁶Sr isotope values (Fig. 10a) highlights the large geographical variability of the values and the sharp division between Sr isotope ratios measured in marginal basins versus those in intermediate-deep water locations. This is only noticeable in substage 3.2, since no (or not enough) material suitable for Sr analysis is present in substage 3.1 deposits from the marginal basins. Some authors see this variability as an indication of isolated subbasins with unique hydrological conditions driven by their catchment rivers (e.g. Müller et al., 1990; Müller and Mueller, 1991; Ryan, 2009). If some degree of connection was present, it involved only neighbouring basins (e.g. Tyrrhenian subbasins; Müller et al., 1990; Müller and Mueller, 1991). A recent comparison of the Sr isotope record of the Spanish marginal basins of Sorbas, Nijar and Vera with the Sr isotope ratios likely to have typified the local riverine sources demonstrated that a local sources-mixed signal expected from an endorheic lake in that location is absent. In this instance mixing of intrabasinal water sources with a non-marine Mediterranean water mass is used to explain the measured values (Andreetto et al., 2021). If this explanation is more widely applicable, then it may result in a re-interpretation of the spread of Sr isotope data from the latest Messinian interval.

To conclude, geochemical proxies have great potential to test the different scenarios, but data are currently too numerically and geographically limited to be robust.

## 7. Paleoenvironmental scenarios for freshening the salt giant: desiccated versus full Mediterranean

The riddle of the Mediterranean environmental and hydrological conditions during Stage 3 is a highly debated topic and it is key to understanding the means by which open marine conditions were restored
at the base of the Zanclean and on the potential impact that the Atlantic-Mediterranean re-connection had on the Atlantic and global climate (Flecker et al., 2015; Capella et al., 2019). In this chapter, the paleo-environmental scenarios, in terms of base-level position (desiccated or full Mediterranean) and hydrological configuration (connections to the Atlantic and/or Paratethys), proposed for the Mediterranean during Stage 3 are described, as well as the different timings of the reflooding (instantaneous, gradual, step-like increments). The low-salinity Stage 3 followed the hypersaline Stage 2 and the transition between the two likely influences the plausibility of the various paleoenvironmental scenarios proposed for the terminal stage. We therefore first summarize the current understanding of the configuration of the Mediterranean during Stage 2 and the enduring controversies (see Roveri et al., 2014a for a more extensive review).

## 7.1. Stage 2 (5.59-5.55 Ma): formation of the Mediterranean salt giant

Numerical modelling based on hydrological budget calculations shows that in order to reach salinity levels compatible with halite saturation and to accumulate the substantial thicknesses of halite observed in the seismic profiles (Ryan, 1973; Haq et al., 2020), the Atlantic-Mediterranean gateway needs to have permitted inflow from the Atlantic, but may have completely blocked outflow (Blanc, 2002; Krijgsman and Meijer, 2008). Numerical models also showed that without Atlantic inflow into the Mediterranean Sea its base level is forced to drop on time scales in the order of a few thousand years by virtue of the basin's negative hydrological budget, where more water is lost to the atmosphere by evaporation than is received from rainfall and river runoff (e.g. Meijer and Krijgsman, 2005; Krijgsman and Meijer, 2008; Simon et al., 2017). The idea of a drawdown is supported by several arguments: (1) the widespread presence, from the margins to the slopes, of the Messinian Erosional Surface cutting through Stage 1 and pre-MSC deposits and canyon incisions following today's drainage networks (e.g. Chumakov, 1973; Clauzon, 1982; Lofi et al., 2005, 2011a, 2011b; Loget et al., 2006; Maillard et al., 2006, 2020; Estrada et al., 2011; Just et al., 2011; Urgeles et al., 2011; Amadori et al., 2018; Lymer et al., 2018; Cazzini et al., 2020; Figs. 5e, 7e); (2) their morphology

interpreted as subaerial in origin; (3) the clastic fans at the outlet of the valleys onlapped by Stage 3 deposits and interpreted as fluvial accumulations (e.g. Lofi et al., 2005; Maillard et al., 2006; Pellen et al., 2019). A number of studies have tried to quantify the magnitude of the sea-level fall by compensating for the isostatic vertical motion since the Messinian to obtain the original depth of the erosional features and Messinian deposits. However, this depends on the assumptions about when the drawdown occurred relative to the halite precipitation: before (e.g. Cartwright and Jackson, 2008; Bache et al., 2009, 2012), during (e.g. Ryan, 2008, 2009) or after (e.g. Ryan, 1978; Bertoni and Cartwright, 2007; Lofi et al., 2011a, 2011b). How shallow the Mediterranean became during Stage 2 is also a matter of disagreement. Estimates in the Western Mediterranean vary from a maximum drawdown of 2500 m (Ryan, 1976) to 1000-1500 m (Bache et al., 2012) in the Gulf of Lion, 800-1200 m in the Balearic promontory (Mas et al., 2018b) and 400 m in the Ebro delta region (Frey-Martinez et al., 2004). A later backstripping analysis of this delta yielded a drawdown of ~1300 m (Urgeles et al., 2011). East of the Sicily sill, backstripping studies estimated base-level drops of 1800-2000 m in the Ionian basin (Micallef et al., 2018, 2019; Camerlenghi et al., 2019; Spatola et al., 2020), 800-900 m in the Adriatic foredeep and Po plain (Ghielmi et al., 2013; Amadori et al., 2018), 800-1300 m (Ben-Gal et al., 2005), 600 (Druckman et al., 1995) and 800 m (Cartwright and Jackson, 2008) in the Levant Basin.

None of these quantifications could unequivocally constrain the timing of the drawdown within the MSC sequence, but numerical modeling studies show that, if the blocking of the outflow was controlled by a tectonic uplift counteracted by inflow erosion across the Strait of Gibraltar, then the expected drawdown of the Mediterranean Sea should be moderate (< 400 m; and possibly harmonic) due to an equilibrium between incision and uplift before the complete blocking of inflow and larger (up to complete desiccation) only after tectonic uplift overcame incision rates (Garcia-Castellanos and Villaseñor, 2011). The same model suggests that the initiation of halite precipitation might overlap in time with the late primary gypsum deposition, right before the full disconnection from the Atlantic Ocean.

The interpretation of the deep evaporites and their associated seismic markers (erosional surfaces and deep engravings along the shelf-slope

<figure><img src="image_24.png" /></figure>

**Fig. 11.** (a), (b) Schematic W-E profiles across the Mediterranean Basin showing the contrasting paleoenvironmental, paleohydrological and paleoconnectivity interpretations proposed for Stage 3. When a water flow is present (green arrow) from and/or to an extra-Mediterranean water mass (i.e., A: Atlantic Ocean; I: Indian Ocean; P: Eastern Paratethys), the direction of the arrow gives the direction of flow. For simplicity, water added by the major and local rivers is not shown, but it adds to the hydrological budget at any time in each scenario. Note the main difference between the isolated (a) and density-stratified (b) scenario lies in the connectivity framework (Atlantic connection closed and negligible influence from the Paratethys in the isolated scenario; influence from both Atlantic and Paratethys in the density-stratified scenario), which affects the position of the base level of the Mediterranean water mass and its hydrochemistry (see extensive discussion in subsection 7.2). Abbreviations: Sp.: SE Spain; V-B: Valencia Basin; Tyr: Tyrhenian Basin; Calt: Caltanissetta Basin; IAP: Ionian Abyssal Plain; GS: Gulf of Sirt; Cyp: Cyprus; Ada: Adana Basin. See Fig. 2 for the geographic position of each basin. (c) Schematic plot showing the evolution of the Mediterranean base-level during Stage 3 according to both the isolated (red line) and half-full (black line) scenarios. The critical sills for controlling intra- and extra-Mediterranean connectivity are also shown.
systems) is not straightforward. Recently, it was suggested that the deep evaporitic facies and the seismic morphological features could have been produced without a significant drop of the Mediterranean base-level, therefore promoting the persistence of a relatively deep-water Mediterranean basin even during halite deposition (Lugli et al., 2013, 2015; Roveri et al., 2014b). For example, Roveri et al. (2014c) proposed that downslope flows of dense, hypersaline waters sourced from evaporation in shallower water areas could have generated both the observed shelf-slope erosion and have created a deep brine, supersaturated in the ions necessary for precipitating halite. These subaqueous hyperpycnal flows are consistent with the observed clastic evaporites that filled the Levant margin canyons (Lugli et al., 2013) and, more generally, with the widespread presence of Complex Units at the outlet of the MES drainage systems (see Lofi et al., 2005, 2011a, 2011b; Lofi, 2018). These sediments are dominated by reworked PLG that would have been exposed by a sea-level fall as little as 200 m (Lugli et al., 2010). However, the hypersaline environment that is presumed to be established by these hyperpycnal flows during the deposition of the RLG is in contrast with the occurrence of the Paratethyan ostracod *L. muelleri* within the clastic evaporites (RLG) in several marginal sections (e.g. Adana Basin, Faranda et al., 2013; Radeff et al., 2016, 2017).

Whatever the state of Mediterranean base-level during Stage 2, the more commonly used chronostratigraphic model for the MSC (Fig. 1a; Roveri et al., 2014a) states that massive halite precipitation ceased at 5.55 Ma and was superseded by an environment that, with precession periodicity (Fig. 3a), cycled between gypsum precipitation and conditions that saw fresh-brackish organisms thriving. The question is whether these conditions cycled homogeneously in several isolated lakes or in basins largely connected to the same Atlantic and Eastern Paratethys-influenced water mass (Fig. 11).

## 7.2. Stage 3 (5.55-5.33 Ma): resumption of (upper) gypsum precipitation and Paratethys fauna invasion

### 7.2.1. An isolated Mediterranean dotted by sabkhas and lakes

The first and long-lasting paleoenvironmental interpretation of the evaporite-bearing UG/UU units and (possibly) time-equivalent evaporite-free units (e.g. LM Unit in Malaga, Sorbas and Zorreras Mb. in Sorbas, Feos Fm. in Nijar, Cassano Spinola Conglomerates in Piedmont, San Donato/Colombacci fms. in the Apennines, Handere Fm. in Turkey) envisaged their sedimentation in a Mediterranean mostly isolated from the Paratethys (which may have added water only to some basins in the Eastern Mediterranean) and totally isolated from the Atlantic where, in each subbasin, continental settings (e.g. alluvial plains, river channels, alluvial fans, playa lakes, sabkhas) alternated/interfingered with shallow, endorheic lakes (Figs. 11a, c; e.g. Ruggieri, 1962, 1967; Decima and Sprovieri, 1973; Decima and Wezel, 1973; Friedman, 1973; Hsü et al., 1973a, 1973b, 1973c, Hsü et al., 1978a, 1978b; Ryan et al., 1973; Selli, 1973; Sturani, 1973; Sissingh, Müller and Mueller, 1991; Orszag-Sperber et al., 2000; Rouchy et al., 2001, 2003, 2007; Blanc, 2002; Lofi et al., 2005, Lofi et al., 2011b; Bassetti et al., 2006; Rouchy and Caruso, 2006; Bertoni and Cartwright, 2007; Cameselle and Urgeles, 2017; Amadori et al., 2018; Camerlenghi et al., 2019; Kartveit et al., 2019; Madof et al., 2019; Ben Moshe et al., 2020; Caruso et al., 2020; Cazzini et al., 2020; Raad et al., 2021). The full disconnection is also supported by observations that support an abrupt Zanclean reflooding (e.g. Blanc, 2002; Micallef et al., 2018, 2019; Garcia-Castellanos et al., 2020; Spatola et al., 2020), since a rapid outburst flood requires a large sea level difference prior to the flood that can only be developed in a scenario of a full Mediterranean-Atlantic disconnection (Garcia-Castellanos et al., 2009; Garcia-Castellanos and Villaseñor, 2011). Although rarely explicitly stated, all these studies must assume that:

1) all Paratethyan biota (and possibly other organisms of undisclosed provenance like diatoms) migrated passively via aquatic migratory birds across the entire Mediterranean (Fig. 11a; Benson, 1978; Benson and Rakic-El Bied, 1991; Caruso et al., 2020);

2) chemical and physical conditions (brackish water and water depth not exceeding 100 m; e.g. Hajós, 1973; Gliozzi and Grossi, 2008) that allowed alternated conditions suitable for gypsum to precipitate and Paratethyan biota and euryhaline benthic foraminifera to thrive were related to changes in the local freshwater budget;

3) The marine isotopic signals in UU/UG gypsum (Fig. 10) are entirely the reflection of the lithologies that are leached by continental waters in surficial and/or underground drainage areas (e.g. Ryan, 2009; Raad et al., 2021);

4) Stage 3 gypsum precipitated in extremely shallow-water (playa lakes) to completely dried environments (sabkhas) and the excessive sulfate necessary is completely derived from “clastic reworking, dissolution, re-precipitation and diagenesis of materials belonging to the PLG and halite of the previous MSC Stage 2” (Ryan, 2009).

Observations supporting a Mediterranean isolated throughout Stage 3 and only at the mercy of local freshwater inputs (Fig. 11a) are: (1) the lack of evidence for in situ marine fauna and flora in UU (e.g. Ryan et al., 1973; Hsü et al., 1978a; Cita et al., 1990; Ryan, 2009; Lofi et al., 2011a); (2) the shallow-water mode of life and highly likely in-situ nature of ostracods and euryhaline, shallow-water benthic foraminifera observed in DSDP/ODP wells from intermediate and deep basins (e.g. Cita et al., 1978; Iaccarino and Bossio, 1999; Figs. 9a-c); (3) the bathymetric contrast (up to several hundred meters) between the late Messinian paleoenvironments and the marine Zanclean on top (e.g. Cita and Colombo, 1979; Bonaduce and Sgarrella, 1999; Caruso et al., 2020); (4) the presence of paleosols in Cyprus (Orszag-Sperber et al., 2000; Rouchy et al., 2001) and on the crest of the Eratosthenes seamount (Robertson, 1998a, 1998b); (5) the erosional features preserved both offshore on the continental shelves and lower-middle slope domain and interpreted in most seismic stratigraphic studies as the result of subaerial exposure (e.g. Lofi et al., 2005; Lofi et al., 2011b; Lymer et al., 2018; Ben Moshe et al., 2020); (6) the pinching out of the UU/BU units towards evaporite-free pre-Messinian structural highs (Figs. 7b-g; Figs. 8a, e; Ryan, 2009; Lymer et al., 2018; Camerlenghi et al., 2019; Raad et al., 2021); (7) the more abundant terrigenous clasts and reworked calcareous fossils in Stage 3 samples compared to the overlying, deep-water Pliocene (Ryan et al., 1973; Hsü et al., 1978b; Ryan, 2009); (8) the erosional nature of the M-reflector/TES/IMTS in the Levant Basin (Figs. 8e-g), by some linked to subaerial exposure of the Levant seafloor (e.g. Bertoni and Cartwright, 2007; Lofi et al., 2011a, 2011b; Maillard et al., 2011a) before the emplacement of deposits interpreted as fluvial from seismic observations (Bowman, 2012; Radeff et al., 2017; Leila et al., 2018; Kartveit et al., 2019; Madof et al., 2019). Furthermore, (9) isolated hydrological circuits with unique chemical composition are regarded by Camerlenghi et al. (2019) as the most plausible explanation for the W-E change in the MSC sedimentary expression in the deep basins, represented by the trilogy LU-MU-UU in the Algero-Balearic and Liguro-Provençal basins, missing the LU in the Tyrrhenian and (possibly) Ionian basins, by terrigenous deposits with hiatuses in the WAB and Adriatic foredeep and by halite, anhydrite and clastics in the Levant Basin (Interbedded and Argillaceous evaporites of Meilijson et al., 2019; Fig. 3b).

The main problems with the isolated scenario lasting throughout Stage 3 are: (1) it does not provide an explanation neither for the homogeneity of Paratethyan ostracod assemblages in the marginal basins (e.g. Gliozzi et al., 2007; Stoica et al., 2016), an aspect difficult to explain when fauna migration takes place passively via either birds or wind, nor for the biomarkers (Vasiliev et al., 2017), which cannot be transported effectively by aquatic birds; (2) it does not explain the mismatch between $^{87}$Sr/$^{86}$Sr isotope ratios measured on marginal ostracods and Sr values expected from endorheic lakes fed with local
freshwaters (e.g. Andreetto et al., 2021); 3) it misses to substantiate, with geochemical arguments, the precipitation of gypsum in lakes, a process that is everything but straightforward (see Warren, 2016 for insights); 4) except for the salt-bearing basins, the source(s) of solutes which makes freshwater-fed endorheic lakes brackish and causes similar physico-chemical conditions to exist in each lake is also difficult to explain in the context of a Mediterranean only at the mercy of local rivers.

### 7.2.2. The half-full, density-stratified Mediterranean scenarios

An alternative concept to the isolated scenario envisages the Mediterranean connected with the Atlantic and/or the Eastern Paratethys and relatively full of water connecting the different subbasins (Fig. 11b). To our knowledge, this scenario was first developed by McCulloch and De Deckker (1989) on the basis of the similar $^{87}$Sr/$^{86}$Sr ratios from marginal (Spain and Cyprus) and deep (Levantine and Algero-Balearic) basins. This intuition was a significant departure from the far more in vogue desiccated scenario (see conclusion of Hsü et al., 1973b), and for this was long overlooked. Sr isotope ratios lower than contemporary ocean water led McCulloch and De Deckker (1989) to conclude that a brackish water mass created by the mixing of water from the peri-Mediterranean rivers (e.g. Nile, Rhône and African rivers that no longer flow today, etc.; see Griffin, 2002 and Gladstone et al., 2007) with water of the Eastern Paratethys filled the Mediterranean, resembling the Caspian Sea today. This conclusion is consistent with the impoverished (or absent) marine fauna and flora of Stage 3 sediments and the enhanced assemblage of fresh-brackish water biota (see subsection 5.7; Figs. 9a-c), but is problematic as a viable origin for Stage 3 gypsum to precipitate at depth. Furthermore, climate models for the late Miocene fail to fill the Mediterranean Basin with fluvial and Paratethys waters alone (Gladstone et al., 2007; Marzocchi et al., 2016, 2019; Simon et al., 2017). A marine contribution is therefore required to fill the Mediterranean (Marzocchi et al., 2016). In the event, the contribution is most likely to have derived from the Atlantic via the Gibraltar Corridor (Flecker et al., 2015; Booth-Rea et al., 2018; Krijgsman et al., 2018) either through a karst system (Krijgsman et al., 2018) or an emerged volcanic archipelago in the Alborán Basin (Booth-Rea et al., 2018). In fact, although an Indian Ocean contribution was proposed (Cita et al., 1978; Hsü et al., 1978a) and the possibility discussed (Ryan, 2009; Vai, 2016), palinspastic reconstructions concluded that the Neo-Tethys Mediterranean-Indian Ocean connection via southern Turkey and Iran already closed before the Tortonian (Rögl, 1998; Popov et al., 2004; Gargani et al., 2008; Bialik et al., 2019; Gülyüz et al., 2020), while a seaway via the Red Sea and Gulf of Aden, although not completely ruled out (e.g. Schütz, 1994; Bosworth et al., 2005; Gargani et al., 2008; Ryan, 2009), is highly contested (e.g. Meulenkamp and Sissing, 2003; Segev et al., 2017).

In light of this, Roveri et al. (2014c), Gvirtzman et al. (2017), Vasiliev et al. (2017), García-Veigas et al. (2018) and Grothe et al. (2020) suggested that the Mediterranean was likely density-stratified during this interval as a result of the simultaneous influx of isotopically-different marine and non-marine (major Mediterranean rivers and Eastern Paratethys) water sources (Fig. 11b). This connectivity framework resulted in a brackish layer carrying low-salinity (mostly Paratethyan) biota (Gliozzi et al., 2007; Stoica et al., 2016; Grothe et al., 2018, 2020; Figs. 9a-b) to lay on top of a more saline layer formed by Atlantic-derived seawater from which UU/UG gypsum (Figs. 5h-j, 7b-g, 8a-d), that facies analyses demonstrated to result from subaqueous deposition (Hardie and Lowenstein, 2004; Lugli et al., 2015), precipitated at intermediate and greater depths (e.g. García-Veigas et al., 2018). A dense, anoxic deep-water mass, possibly inherited from Stage 2, is envisaged at the bottom of the Mediterranean by Marzocchi et al. (2016) and García-Veigas et al. (2018), albeit without conclusive arguments, and by Gvirtzman et al. (2017) following the observation that the tilted halite body of the Levant Basin was simultaneously eroded landward and preserved basinward (Fig. 8f).

This scenario accounts for the erosive/non-depositional features (Figs. 5e, 6a, e) and continental/lacustrine facies (Figs. 5a-b, d-g) widespread around the margins and shelves and suggestive of a Mediterranean base-level somewhat lower than the Atlantic level suggesting a one-way inflow from both the Atlantic and the Eastern Paratethys after Stage 2 (e.g. Marzocchi et al., 2016; Figs. 11b, c), a connectivity configuration that effectively translates in a half-full Mediterranean (e.g. Krijgsman and Meijer, 2008). Refilling as a result of persistent Atlantic inflow, in part perhaps because of the latest Messinian deglaciation (see subsection 2.2; Van der Laan et al., 2006; Hilgen et al., 2007), would have resulted in the establishment of two-way exchange first with the Paratethys at some point during the Lago-Mare phase and later, i.e. slightly before or at the Messinian/Zanclean boundary, with the Atlantic Ocean (Fig. 11c; Marzocchi et al., 2016). The moment the Mediterranean base-level reached the sill with the adjacent water body (Paratethys and Atlantic) and a two-way exchange was initiated, the density contrast will have prompted an enhanced inflow into the Mediterranean (Marzocchi et al., 2016). The overall transgressive trend leading to the Zanclean marine replenishment was accompanied by base-level fluctuations in the order of 400±100 m every precessional cycle (Fig. 11c; Fortuin and Krijgsman, 2003; Ben Moshe et al., 2020; Andreetto et al., 2021). These fluctuations are ascribed to switch in the Mediterranean freshwater budget driven by the African summer monsoon and Atlantic winter storms (e.g. Marzocchi et al., 2015, 2019; Simon et al., 2017). Since higher freshwater discharge rates occur at precession minima times and their Stage 3 sedimentary expression is considered to be the mudstone intervals (Fig. 3a; Manzi et al., 2009), mudstone interbeds (both onshore and offshore; e.g. Figs. 5h-j) represent the highstand episodes (e.g. Manzi et al., 2009; Roveri et al., 2008a; Omodeo-Salé et al., 2012; Fig. 3), while continental facies onshore (e.g. conglomerates in the Apennines; Fig. 5g) and offshore (clastic beds in the Levant Basin) and gypsum beds (Algero-Balearic, Liguro-Provençal, CMD, Tyrrenian, Caltanissetta, Ionian, Sirte and Polemi-Pissouri basins; Figs. 5h-j) represent the lowstand (e.g. Roveri et al., 2008a; Manzi et al., 2009; Meilijson et al., 2019; Fig. 3). If Atlantic was the major source of sulfate for Stage 3 gypsum (e.g. García-Veigas et al., 2018) and an intervening, relatively shallow (Sicily) sill was establishing Western and Eastern Mediterranean division during the MSC (e.g. García-Castellanos et al., 2009, 2020; Micallef et al., 2018), the presence of Stage 3 gypsum to the east of the Sicily sill (Fig. 2b) implies that the Mediterranean base level never dropped below the (maximum estimated) paleodepth of the sill (i. e. ~430 m; García-Castellanos et al., 2009) during Stage 3 and Western and Eastern Mediterranean remained connected also during the arid (lowstand) phases of the precession cycles.

A Mediterranean step-wise refilled and at times filled with water up to the marginal belt agrees with: (1) Paratethyan biota being present only in intermediate-deeper settings during substage 3.1, but more widespread including marginal settings during substage 3.2; (2) the W-E homogeneity of Paratethyan ostracod assemblages around the Mediterranean marginal belt (Gliozzi et al., 2007; Stoica et al., 2016; Sciuto et al., 2018; Sciuto and Baldanza, 2020; Fig. 9a); (3) the presence, in marginal basins, of Paratethyan fish (Bannikov et al., 2018; Schwarz-hans et al., 2020), dinocysts (e.g. Pellen et al., 2017; Fig. 9b) and biomarkers (Vasiliev et al., 2017; Fig. 10c); (4) the occurrence of a monospecific assemblage of abundant *Sphenolithus* spp. just below the M/P boundary at ODP Sites 978, 975 and 967 (Castradori, 1998); (5) the requirement of water from the Mediterranean to explain the Sr isotope ratios of ostracods that inhabited marginal subaqueous environments (Andreetto et al., 2021); (6) the Atlantic-like sulfate values (although variably diluted and affected by microbial processes; Fig. 10b) of the UU/UG gypsum beds (García-Veigas et al., 2018); (7) the presence of long chain alkenones in the Sicilian UG beds similar to those observed in present-day marine settings (Fig. 10c; Vasiliev et al., 2017).

Major problems also exist with the half-full stratified scenario: (1) it does not provide a proper mechanism for gypsum precipitation at several hundreds, or thousands, meters water depth; (2) it fails to explain how unquestionable shallow-water (< 50 m) benthic organisms
such as *Ammonia tepida* and *Cyprideis* sp. could survive at hundreds of meters of depth and beyond; (3) it does not provide an explanation for the high abundance of coarse-grained detritus at intermediate and deep-water locations, especially when compared to deep-water Pliocene samples, as well as for the broad absence of MSC deposits in the shelf domain; (4) a persistent Atlantic inflow without outflow seems to be a configuration that cannot be maintained stable for ~200 kyr. Indeed, models coupling the inflow of marine waters with the erosion of the gateway channel concluded that, if the Mediterranean level was lowered by at least several hundred meters below present sea level, any small overtopping of water from the Atlantic would inevitably trigger a fast refill of the basin that, if responsible for the erosion trough the Alborán Basin, should have involved an unprecedented water discharge and be completed within a few years or less (Garcia-Castellanos et al., 2020 and references therein). The scenario arising from Meilijson et al., 2019, Figs. 3b, 4a) is also problematic for a high base-level Mediterranean. In order to simultaneously reach precipitation of gypsum and halite in different basins sharing the same water, the water has to be of high salinity and stratified. Simon and Meijer (2017) demonstrated that this can be achieved with slow overturning circulation, but it is currently unclear how realistic this process is.

### 7.3. Demise of the MSC (5.33 Ma): the Zanclean marine replenishment

The conspicuous and abrupt transition to normal marine sediments in the Mediterranean is globally and historically important because it is the origin of the stratigraphic position of the Miocene/Pliocene boundary (Van Couvering et al., 2000). From an ocean perspective, it is not an ideal stratigraphic location being difficult to locate from biozone data even in the adjacent Atlantic (Hodell et al., 2001; Krijgsman et al., 2004; Van den Berg et al., 2015). However, from a Mediterranean perspective it provides a clear and unambiguous end to the MSC and the restoration of normal marine conditions. All evidence show that the onset of the Zanclean marine replenishment followed a period of relative lowstand that exposed all the Mediterranean margins (see subsection 3.12; Figs. 6a-b, f) and kept intermediate and deep basins underwater (see subsection 4.8). Yet again, the key dispute concerns the exact depth of the Mediterranean base level preceding the Miocene/Pliocene transition.

Building on the isolated Mediterranean scenario, base level immediately before the early Zanclean was more than thousand kilometers below eustatic sea level (Fig. 10e; e.g. Hsü et al., 1973a; Blanc, 2002; Loget et al., 2006; Garcia-Castellanos et al., 2009; Pérez-Asensio et al., 2012; García-Alix et al., 2016; Amadori et al., 2018; Micallef et al., 2018, 2019; Camerlenghi et al., 2019; Kartveit et al., 2019; Madof et al., 2019; Ben Moshe et al., 2020; Caruso et al., 2020; Cazzini et al., 2020; Mas and Fornós, 2020; Spatola et al., 2020). Hydrodynamic erosional models allowed a reinterpretation of the erosional features across the strait of Gibraltar (Campillo et al., 1992; Blanc, 2002) suggesting that a sudden breach of the Mediterranean-Atlantic divide at Gibraltar resulted in a vast cascade of Atlantic water that refilled the entire Mediterranean in less than 2 years (i.e. rates of ten meters per day) spilling first into the Western Mediterranean (see the extensive review in Garcia-Castellanos et al., 2020) and then, after reaching the level of the Sicily sill, pouring into the Eastern Mediterranean (Micallef et al., 2018, 2019; Ben Moshe et al., 2020; Spatola et al., 2020). This concept of catastrophic refilling has led to terms such as “Zanclean flood” or “deluge”. Evidence supporting the catastrophic flood mechanism mostly comes from the seismic reflection dataset and includes: 1) the presence of >250 m deep and 390-km-long incisions on both sides of the Gibraltar Strait (Garcia-Castellanos et al., 2020); 2) the detection of (allegedly) Pliocene-aged chaotic sedimentary bodies stretching for kilometers in the Alborán Basin (Garcia-Castellanos et al., 2020 and references therein) and Ionian Basin at the foot of the Malta Escarpment (Micallef et al., 2018, 2019; Spatola et al., 2020; Fig. 8c). A further argument is the bathymetric jump of several hundred meters between the late Messinian and the early

Pliocene sediments (e.g. Caruso et al., 2020; Fig. 6d).

Instantaneous sea level rise is not the only possible refilling model. Bache et al. (2012) suggested the reflooding occurred in two steps at ~5.60 Ma, accompanied by a moderate (≤ 500 m) rise, followed by a rapid rise of 600-900 m at around 5.46 Ma tracking the deposition of the deep basin evaporites and resulting from the collapse of the Gibraltar divide. There is also the reconnection model that follows from a Stage 3 Mediterranean that is already relatively full and with the base level possibly oscillating of 400 ± 100 m with precessional frequency (Fig. 10h; Fortuin and Krijgsman, 2003; Ben Moshe et al., 2020; Andreetto et al., 2021). In this case, only a sea level rise of a few hundred meters is required to restore the Mediterranean to the Atlantic level (Fig. 10h), which was hypothesized to have occurred in the last precessional cycle of the Messinian (Marzocchi et al., 2016; Fig. 3a).

In detail, the re-establishment of a fully marine faunal diversity and oceanic geochemistry (e.g. $^{87}$Sr/$^{86}$Sr ratios and $\delta^{18}$O) occurred more gradually over one or more precessional cycles in the earliest Zanclean (e.g. Iaccarino et al., 1999; Pierre et al., 1998, 2006; Cipollari et al., 2013; Roveri et al., 2019a; Bulian et al., 2021). This suggests that stressed ecological conditions at first only suitable for opportunistic organisms to survive (e.g. Bulian et al., 2021) developed (or persisted) in the Mediterranean as marine replenishment occurred (e.g. Rouchy et al., 2003). One possible mechanism for achieving this may be the physicochemical turnover in the water column triggered by the re-established two-way exchange with the Atlantic which, for reasons that are largely unknown, took time (at least oneprecession cycle; Pierre et al., 2006) to displace surficial Paratethyan water and restore normal marine conditions (Marzocchi et al., 2016).

## 8. Methods and proxies to better reconstruct base level and connectivity changes

Chronological uncertainty and spatial variability limit the use of both sedimentological and paleontological information to achieve a comprehensive and coherent basin-wide interpretation of the conditions and drivers of Stage 3 environments and water levels. Alternative methods are therefore required to clarify connectivity relationships and constrain base-level conditions. This section explores the principles and potential of geochemical, backstripping and numerical modelling techniques that could be used to further test existing hypotheses and enhance understanding of the complex environmental conditions experienced by the Mediterranean during the latest Messinian.

### 8.1. Geochemical proxies

**Radiogenic strontium isotopes.** Radiogenic strontium isotope ratio ($^{87}$Sr/$^{86}$Sr) is a widely applied geochemical tool in provenance studies, including the reconstruction of the hydrological circuit and connectivity of basins with little or null oceanic entries. Its potential to detect the provenance of the hydrological fluxes derives from the unique $^{87}$Sr/$^{86}$Sr ratio that typifies each water source and from the negligible effects of isotopic fractionation during the liquid-solid transition (see Hajj et al., 2017).

Mineral phases precipitating in endorheic lakes uptake Sr with $^{87}$Sr/$^{86}$Sr ratio that reflects the mixing of all feeding surficial and underground streams and whose $^{87}$Sr/$^{86}$Sr fingerprint hinges on the composition and age of watershed bedrock (see Peucker-Ehrenbrink and Fiske, 2019; Andreetto et al., 2021 and references therein). When river water mixes with seawater such as in the oceans, semi-enclosed basins or estuaries, mineral phases uptake Sr with oceanic $^{87}$Sr/$^{86}$Sr ratios because the high oceanic Sr concentration (~7.8 mg/l today; Veizer, 1989) masks the impact of the ~100 times lower concentrated continental Sr-sources (~0.0780 mg/l; Palmer and Edmond, 1992). This is valid as long as a certain ratio of continental-marine water mixing is fulfilled, beyond which $^{87}$Sr/$^{86}$Sr ratios deviate towards the $^{87}$Sr/$^{86}$Sr ratios of the non-marine source(s) (Ingram and Sloan, 1992). For the
Mediterranean to attain non-marine $^{87}$Sr/$^{86}$Sr ratios (like during the MSC), Topper et al. (2014) calculated a mixing of at least 1:4 (Atlantic: freshwater) to be required.

If Mediterranean subbasins hosted endorheic lakes (Figs. 10c, e), the $^{87}$Sr/$^{86}$Sr isotope ratios measured on ostracod valves or gypsum crystals of each lake are expected to generate a scattered distribution by virtue of the different geology in the hinterland of each basin. By contrast, some degree of connection between different basins and the Mediterranean water mass (Figs. 10d, f) is expected to result in more homogeneous $^{87}$Sr/$^{86}$Sr ratios because, although isotopically-different, local rivers mix with a water mass that has the same $^{87}$Sr/$^{86}$Sr value and (much higher) Sr concentration for each basin (Andreetto et al., 2021). In this scenario, differences in the $^{87}$Sr/$^{86}$Sr ratios between basins are likely the reflection of the different $^{87}$Sr/$^{86}$Sr ratio of the local input in each basin (Andreetto et al., 2021). The application of numerical models assists to identify and quantify the different water sources feeding the basin(s) in question and (e.g. Placzek et al., 2011; Topper et al., 2011, 2014; Doebber et al., 2014; Rossi et al., 2015b; Modestou et al., 2017; Grothe et al., 2020; Andreetto et al., 2021).

### Sulfate isotopes.
When sulfate-bearing minerals precipitate in a basin they uptake dissolved S and O with $\delta^{34}\text{SSO}_4$ and $\delta^{18}\text{O SO}_4$ isotopic composition that, once corrected for the fractionation effects during liquid-solid transition (see subsection 6.2), can be assimilated to that of the mother brine. The higher concentrated source of sulfate is seawater (with present-day $\delta^{34}\text{SSO}_4=21.15\pm0.15\%$ and $\delta^{18}\text{O SO}_4=8.67\pm0.21\%$, Johnston et al., 2014; with Messinian values of ~22±0.2‰ for the $\delta^{34}\text{SSO}_4$ and ~9±2‰ for the $\delta^{18}\text{O SO}_4$; Turchyn and Schrag, 2004; Markovic et al., 2016; Masterson et al., 2016). Significantly higher inputs from the ~1000 times less concentrated riverine freshwater (both superficial and underground) with respect to the ocean water (more than 1:5 according to Lu et al., 2001) can modify the marine $\delta^{34}\text{SSO}_4$ and $\delta^{18}\text{O SO}_4$ isotopic composition of the mother brine (Utrilla et al., 1992; Lu et al., 2001) and have it deviated from that of the ocean (Lu et al., 2001). This deviation is normally towards lower values, because river-derived dissolved sulfate is generally depleted in heavy isotopes $^{34}$S and $^{18}$O compared to oceanic sulfate because these isotopes mainly come from the oxidation of $^{34}$S-depleted pyrite ($\text{FeS}_2$) on the continents and to a lesser extent from the dissolution of older sulfate-bearing minerals (Claypool et al., 1980; Turchyn and Schrag, 2004; Burke et al., 2018). However, when marine sulfate is preferentially leached in the catchment, $^{34}$S of the freshwater-dissolved sulfate and $[^{2}O_4]^{\text{S}}$ likely increase, therefore reducing the continental-marine mixing ratio necessary to deviate the resulting sulfate isotopic signature away from marine values. Unfortunately, the sulfate isotopic composition is not provided for a number of major Mediterranean rivers (Burke et al., 2018) nor for the Eastern Paratethys and it is hardly assessed with the catchment-forming lithologies (Liu et al., 2017; Burke et al., 2018), making sulfate isotopes still an unsuitable tracer of non-marine water provenance in Mediterranean subbasins.

Deviation of $\delta^{34}\text{SSO}_4$ and $\delta^{18}\text{O SO}_4$ from the marine average can also be the result of isotopic fractionation during microbial sulfate reduction (MSR; Fritz et al., 1989; Berner, 1999). MSR produces $^{34}$S-depleted hydrogen sulfide (~0 to 70‰ lighter than initial sulfate; Brunner and Bernasconi, 2005; Sim et al., 2011; Leavitt et al., 2013) and induces the enrichment in $^{34}$S and $^{18}$O of the residual sulfate pool (Kaplan and Rittenberg, 1964; Thode and Monster, 1965; Turchyn et al., 2006; Wortmann et al., 2007). Therefore, if isotopically light H$_2$S produced by MSR leaves the system as a sulfide mineral (most likely pyrite), the resulting dissolved sulfate would have $\delta^{34}\text{SSO}_4$ and $\delta^{18}\text{O SO}_4$ isotopic signatures higher than the oceanic one (Brunner et al., 2005). However, if the MSR-produced H$_2$S is re-oxidized back to sulfate through abiotic or microbial sulfide oxidation, isotopically light sulfate will be brought back to the $^{34}$S-enriched sulfate pool, producing little or no enrichment in $^{34}$S observed in the resulting sulfate (Gomes and Johnston, 2017 and references therein; Pellerin et al., 2019). Slight deviations from marine $\delta^{18}\text{O SO}_4$ and $\delta^{34}\text{SSO}_4$ values of sulfate reflect both biological sulfur

cycling and/or freshwater riverine inputs (e.g. Utrilla et al., 1992; Lu et al., 2001; Turchyn et al., 2009) (Fig. 10b). Untangling the relative importance of these processes is key to understanding the Mediterranean sulfur isotope record and gleaning paleoenvironmental insights into Stage 3.

Hydrogen isotopes. Organic geochemistry biomarker-based tools can be used as independent proxies for reconstructing sea surface temperatures, relative changes in the basin hydrology and, indirectly, salinity. Basin water properties are reflected in a variety of life forms. Different types of organisms produce specific organic compounds that serve as molecular biomarkers. These large biomolecules record the changes in the hydrogen isotopic composition of the water used by different groups of biomarker producers (i.e. different organisms). The principle behind the method is to measure $\delta\text{D}$ on biomarkers produced in Mediterranean Sea waters (e.g. alkenones, produced by a few species of haptophyte coccolithophores algae) during the MSC and compare the results with the $\delta\text{D}$ signals retrieved from biomarkers produced in the open ocean ideally at the same time intervals. The influence of precipitation on the changes in hydrological budget can be monitored by measuring the $\delta\text{D}$ of long chain *n*-alkanes (Sachse et al., 2006), biomarkers predominantly produced by higher terrestrial plants that rely on precipitation for plant growth, therefore reflecting the changes in the $\delta\text{D}$ of the precipitation. The extreme base level drop(s) suggested for the Mediterranean during the MSC would, in principle, indicate a negative precipitation (P) + runoff (R) – evaporation (E) ratio. Such a negative water budget ($E>P+R$) results in waters increasingly enriched in $\delta\text{D}$ whereas, a positive water balance ($E<P+R$) results instead in a negative shift of $\delta\text{D}$ values. The analysis of compound specific $\delta\text{D}$ of alkenones, long and short chain *n*-alkanes can be used to constrain E/(P+E) relationships.

## 8.2. Backstripping analyses

Backstripping uses paleobathymetry, sea level and sediment thickness to quantify the tectonic and isostatic components of subsidence. If tectonic subsidence or uplift history are known relative to the current position and depth of paleoshoreline markers, an inverse approach allows base level to be estimated. A number of approaches have been applied to the MSC, using erosional surfaces (e.g. Amadori et al., 2018), terraces (Micallef et al., 2018) or fluvial network characteristics (Urgeles et al., 2011) as paleoshoreline indicators. The relief on erosional features has also been used to estimate minimum base-level variation (Frey-Martinez et al., 2004).

Apart from the quantitative constraints on base level that backstripping provides, consideration of the regional implications of isostatic subsidence and the gravitational impact of redistributing water masses (such as in the cascading model of Roveri et al., 2014c; Fig. 10b) and evaporite precipitation is important in gateway regions like Gibraltar, which due to their shallow and restricted nature are exceptionally sensitive to vertical motions. Here, both flexural effects and gravitational effects on local sea level on the Atlantic side of the strait has the potential to influence Mediterranean-Atlantic connectivity driving paleoenvironmental changes in the basin itself (Coulson et al., 2019).

## 8.3. Modelling

Numerical models can be used complementary to lab- and field-based studies, or to find answers to open questions by testing the physical plausibility of hypotheses and their compatibility with the available sedimentological/stratigraphic/paleontological/geochemical data, which have to constrain model results and not adjust to it. For example, whether gypsum beds in marginal/intermediate basins can precipitate at the same time as the halite in deep basins is an intriguing question that circulates in the MSC literature (e.g. Van Couvering et al., 1976), but whether this is physically and geochemically possible is yet to be answered. In a first model analysis, Simon and Meijer (2017) found that the required stratification can indeed be achieved for specific
conditions including a slow overturning circulation. A different approach is needed to determine whether such slow circulation is to be expected or if other scenarios should be considered. A thermo-haline stratification that enables coeval precipitation of two evaporites for a considerable time span might also influence the degree of heterogeneity of other parameters, such as strontium concentration. Previous models showed the influence of the Atlantic Ocean and major rivers (Topper et al., 2014) and of evaporation (Flecker et al., 2002) on the Sr value of a basin with restricted oceanic inflow and assuming a homogeneous distribution of strontium throughout the water column (Flecker et al., 2002; Topper et al., 2011, 2014; Modestou et al., 2017), but it is uncertain if this holds true in conditions of water stratification. New insights into this behavior would have consequences for the way the strontium dataset (Fig. 10a) must be interpreted. Another loose end where the model approach can provide insight relates to the question whether a high water level could have been reached without an inflow from the Atlantic. Climate simulations conducted by Gladstone et al. (2007), Simon et al. (2017) and Marzocchi et al. (2019) indicate that this is not possible with today's bathymetry. A quantitative analysis exploring the Mediterranean water level reached in different situations (i.e. with or without an Atlantic or Paratethys in and outflow) and with information on the Mediterranean hypsometry that may be provided by isostatic restoration of the seafloor topography (flexural backstripping) could help understanding how the Messinian Salinity Crisis ended.

## 9. Certainties, open problems and future directions

Our understanding of the nature of MSC Stage 3 has evolved considerably over the last fifty years. However, there are still such disparate models for the paleoenvironmental conditions and basin connectivity that led to Stage 3 deposition and that express the challenges that the study of this interval presents: it is a relatively short interval and its sedimentary expression varies spatially. It is no surprise that the main point of contention lies in reconciling the observations from seismic profiles and well data, which are largely interpreted as favoring the desiccated scenario, with the sedimentological, paleontological and geochemical data from the marginal basins record, largely discontinuous and unaddressed from seismic-based and computational-based studies and mainly supporting the full-Mediterranean hypothesis.

To a large extent this mismatch is the result of the lack of intersection of the two datasets. Some Stage 3 onshore localities are meticulously studied from the stratigraphic, sedimentological, paleontological and geochemical point of view, showing remarkably consistent and homogeneous trends and patterns (as previously highlighted by Roveri et al., 2008a). However, changes at precessional and subprecessional scale are difficult to trace from one exposure to another and are well below the level of seismic resolution, making onshore-offshore correlation at this scale impossible. Even correlation between onshore sections is problematic since most of the stratigraphic sections are incomplete, with erosion surfaces at the bottom and/or top (i.e. the Miocene/Pliocene boundary), and this lack of stratigraphic continuity frustrates attempts to constrain ages by cyclostratigraphy. A future focus on strengthening the available chronostatigraphic framework (Fig. 3) and making it inclusive of the fragmented outcrops is required to better understand the paleoenvironmental and paleohydrological changes suffered by the Mediterranean marginal belt through time. The successful drilling of the three IODP proposals currently in the scheduling pool (see Camerlenghi and Aloisi, 2020), all of which propose to recover Stage 3 sediments, will also provide transformative information enabling far better offshore-onshore correlation and interpretation of currently enigmatic seismic data. In the meantime, re-evaluation of existing DSDP and ODP material, particularly through the application of more novel geochemical techniques and, where possible, access to material collected during industrial drilling would be helpful for understanding the deep Mediterranean Basin during this interval.

Extensive paleontological studies have established that Stage 3

contains in situ biota assemblages of Paratethyan provenance implying brackish water conditions. More problematic is the differentiation of in situ and reworked marine microfauna and flora and the paleoecological significance of dwarfism in marine calcareous microfossils/algae. These have important repercussions for the Mediterranean connectivity and base-level reconstruction.

The geochemical dataset for Stage 3, particularly strontium isotopes and hydrogen isotopes on biomarkers, is both demonstrably valuable in providing key constraints on connectivity and environmental conditions, and frustratingly inadequate in terms of data distribution. It has great potential as a constraint on quantitative sensitivity analysis of the different hydrochemistry scenarios that follow from the endmember Stage 3 hypotheses, but substantially more data is required.

An approach which combines geological, geochemical, geophysical and paleontological data with numerical modelling (e.g. climate simulations, backstripping analyses and paleoceanographic models) will provide a more accurate reconstruction of Mediterranean paleogeography and the processes that occurred during the final phase of the Messinian Salinity Crisis.

## Data availability

The compilation of strontium, sulfate and hydrogen isotope data plotted in figure 9, as well as some of the available, and here not (graphically) shown, oxygen isotope values is accessible in separate excel spreadsheets (Supplementary material 1).

## Declaration of Competing Interest

The authors whose names are listed immediately below certify that they have NO affiliations with or involvement in any organization or entity with any financial interest (such as honoraria; educational grants; participation in speakers' bureaus; membership, employment, consulancies, stock ownership, or other equity interest; and expert testimony or patent-licensing arrangements), or non-financial interest (such as personal or professional relationships, affiliations, knowledge or beliefs) in the subject matter or materials discussed in this manuscript.

## Acknowledgments

We thank the entire SALTGIANT community for many profitable workshops that inspired the development of this manuscript. This research was supported by the project SALTGIANT-Understanding the Mediterranean Salt Giant, a European project which has received funding from the European Union's Horizon 2020 research and innovation program, under the Marie Sklodowska-Curie [grant agreement No 765256]. We greatly thank the two reviewers Domenico Cosentino and William Ryan and the editor Alessandra Negri for the fruitful comments provided that led to a substantial improvement of the manuscript.

## Appendix A. Supplementary data

Supplementary data to this article can be found online at https://doi.org/10.1016/j.earscirev.2021.103577.

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