# Freshening of the Mediterranean Salt Giant: controversies and certainties around the terminal (Upper Gypsum and Lago-Mare) phases of the Messinian Salinity Crisis

F. Andreetto, G. Aloisi, Fadl Raad, H. Heida, R. Flecker, K. Agiadi, J. Lofi, S. Blondel, F. Bulian, A. Camerlenghi, et al.

<figure><img src="image_1.png" /></figure>

## ► To cite this version:

F. Andreetto, G. Aloisi, Fadl Raad, H. Heida, R. Flecker, et al. Freshening of the Mediterranean Salt Giant: controversies and certainties around the terminal (Upper Gypsum and Lago-Mare) phases of the Messinian Salinity Crisis. Earth-Science Reviews, 2021, 216, pp.103577. 10.1016/j.earscirev.2021.103577. hal-03200312

HAL Id: hal-03200312

https://hal.science/hal-03200312

Submitted on 17 Apr 2021

HAL is a multi-disciplinary open access archive for the deposit and dissemination of scientific research documents, whether they are published or not. The documents may come from teaching and research institutions in France or abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est destinée au dépôt et à la diffusion de documents scientifiques de niveau recherche, publiés ou non, émanant des établissements d'enseignement et de recherche français ou étrangers, des laboratoires publics ou privés.

<figure><img src="image_2.png" /></figure>
<figure><img src="image_3.png" /></figure>

# Freshening of the Mediterranean Salt Giant: controversies and certainties around the terminal (Upper Gypsum and Lago-Mare) phases of the Messinian Salinity Crisis

F. Andreetto<sup>a,∗</sup>, G. Aloisi<sup>b</sup>, F. Raad<sup>c</sup>, H. Heida<sup>d</sup>, R. Flecker<sup>e</sup>, K. Agiadi<sup>f</sup>, J. Lofi<sup>c</sup>, S. Blondel<sup>g</sup>, F. Bulian<sup>h</sup>, A. Camerlenghi<sup>g</sup>, A. Caruso<sup>i</sup>, R. Ebner<sup>j</sup>, D. Garcia-Castellanos<sup>d</sup>, V. Gaullier<sup>k</sup>, L. Guibourdenche<sup>b</sup>, Z. Gvirtzman<sup>l,m</sup>, T.M. Hoyle<sup>a,n</sup>, P.T. Meijer<sup>j</sup>, J. Moneron<sup>l,m</sup>, F.J. Sierro<sup>h</sup>, G. Travan<sup>k</sup>, A. Tzevahirtzian<sup>i</sup>, I. Vasiliev<sup>o</sup>, W. Krijgsman<sup>a</sup>

## ARTICLE INFO

### Keywords:

Messinian Salinity Crisis
Mediterranean stratigraphy
Connectivity proxies
Paleogeography
Paratethys
Lago-Mare

## ABSTRACT

The late Miocene evolution of the Mediterranean Basin is characterized by major changes in connectivity, climate and tectonic activity resulting in unprecedented environmental and ecological disruptions. During the Messinian Salinity Crisis (MSC, 5.97-5.33 Ma) this culminated in most scenarios first in the precipitation of gypsum around the Mediterranean margins (Stage 1, 5.97-5.60 Ma) and subsequently > 2 km of halite on the basin floor, which formed the so-called Mediterranean Salt Giant (Stage 2, 5.60-5.55 Ma). The final MSC Stage 3, however, was characterized by a "low-salinity crisis", when a second calcium-sulfate unit (Upper Gypsum; substage 3.1, 5.55-5.42 Ma) showing (bio)geochemical evidence of substantial brine dilution and brackish biota-bearing terrigenous sediments (substage 3.2 or Lago-Mare phase, 5.42-5.33 Ma) deposited in a Mediterranean that received relatively large amounts of riverine and Paratethys-derived low-salinity waters. The transition from hypersaline evaporitic (halite) to brackish facies implies a major change in the Mediterranean's hydrological regime. However, even after nearly 50 years of research, causes and modalities are poorly understood and the original scientific debate between a largely isolated and (partly) desiccated Mediterranean or a fully connected and filled basin is still vibrant. Here we present a comprehensive overview that brings together (chrono)stratigraphic, sedimentological, paleontological, geochemical and seismic data from all over the Mediterranean. We summarize the paleo-environmental, paleohydrological and paleoconnectivity scenarios that arose from this cross-disciplinary dataset and we discuss arguments in favour of and against each scenario.

<footnote><sup>a</sup> Paleomagnetic Laboratory "Fort Hoofddijk", Dept. of Earth Sciences, Utrecht University, Budapestlaan 17, 3584 CD Utrecht, The Netherlands</footnote>

<footnote><sup>b</sup> Université de Paris, Institut de physique du globe de Paris, CNRS, F-75005 Paris, France</footnote>

<footnote><sup>c</sup> Géosciences Montpellier, CNRS, Université de Montpellier, Montpellier, France</footnote>

<footnote><sup>d</sup> Geosciences Barcelona, GEO3BCN, CSIC, Solè i Sabarís s/n, Barcelona, Spain</footnote>

<footnote><sup>e</sup> BRIDGE, School of Geographical Sciences and Cabot Institute, University of Bristol, University Road, Bristol BS8 1SS, United Kingdom</footnote>

<footnote><sup>f</sup> Department of Palaeontology, University of Vienna, Althanstraße 14 (UZA II), 1090 Vienna, Austria</footnote>

<footnote><sup>g</sup> Istituto Nazionale di Oceanografia e di Geofisica Sperimentale (OGS), Trieste, Italy</footnote>

<footnote><sup>h</sup> Department of Geology, University of Salamanca, Salamanca, Spain</footnote>

<footnote><sup>i</sup> Dipartimento di Scienze della Terra e del Mare, Università degli studi di Palermo, via Archirafi 20-22, 90123 Palermo, Italy</footnote>

<footnote><sup>j</sup> Department of Earth Sciences, Utrecht University, Utrecht, The Netherlands</footnote>

<footnote><sup>k</sup> Univ. Lille, CNRS, Univ. Littoral Côte d'Opale, UMR 8187, LOG, Laboratoire d'Océanologie et de Géosciences, F 59000, Lille, France</footnote>

<footnote><sup>l</sup> Geological Survey of Israel, Jerusalem 95501, Israel</footnote>

<footnote><sup>m</sup> Institute of Earth Sciences, The Hebrew University of Jerusalem, Israel</footnote>

<footnote><sup>n</sup> CASP, West Building, Madingley Rise, Madingley Road, Cambridge CB3 0UD, United Kingdom</footnote>

<footnote><sup>o</sup> Senckenberg Research Biodiversity and Climate Research Centre, Senckenberganlage 25, 60325 Frankfurt am Main, Germany</footnote>

<footnote>* Corresponding author.
E-mail address: f.andreetto@uu.nl (F. Andreetto).</footnote>
# 1. Introduction

At the end of the Miocene, orbital and tectonic drivers combined to alter the amount of water delivered to the Mediterranean Basin by the Atlantic Ocean from the west, the brackish Eastern Parathethys (i.e. Euxinic-Caspian Basin system) from the east and the major peri-Mediterranean freshwater drainage systems (e.g. African rivers and Rhône; Griffin, 2002; Gladstone et al., 2007; Van der Laan et al., 2006; Hilgen et al., 2007; Ryan, 2009; Flecker et al., 2015; Marzocchi et al., 2015, 2016, 2019; Simon et al., 2017; Krijgsman et al., 2018; Capella et al., 2020). The changes in extra and intrabasinal connectivity resulted in unprecedented paleoceanographic and paleohydrological budget changes that led to a relatively short-lived environmental and ecological crisis (approx. 660 kyr; 5.97-5.33 Ma), for which the term Messinian Salinity Crisis (MSC) was coined (Selli, 1954, 1960). Most conspicuous was the rapid accumulation of several kilometers of halite (i.e. ~1 million km³) on the Mediterranean abyssal plains (e.g. Hsü, 1972; Ryan, 1973; Montadert et al., 1978; Haq et al., 2020). This happened within 50 kyr, from 5.60-5.55 Ma, according to Roveri et al. (2014a) and Manzi et al. (2018), or in >300 kyr, when starting at 5.97 Ma, as put forward by Meilijson et al. (2018, 2019).

During the ~200 kyr lapse (i.e. MSC Stage 3 following Roveri et al., 2014a; Fig. 1a) between the end of salt precipitation (5.55 Ma) and the restoration of the still enduring marine conditions (5.33 Ma), the Mediterranean underwent a sequence of paleohydrological and base-level changes that are the topic of intense and long-standing debates. The initial and still widely endorsed hypothesis was that the Mediterranean Sea, following the major drawdown event that led to halite deposition (i. e. Stage 2), maintained the isolated, deeply-desiccated geography containing a series of hypersaline (substage 3.1; 5.55-5.42 Ma) and

hyposaline (substage 3.2; 5.42-5.33 Ma) ponds which only received water from local streams and were colonized by Black Sea organisms carried by aquatic migratory birds (Fig. 1b; e.g. Ruggieri, 1967; Decima and Sprovieri, 1973; Decima and Wezel, 1971, 1973; Cita et al., 1978; Müller et al., 1990; Benson and Rakic-El Bied, 1991; Benson et al., 1991; Müller and Mueller, 1991; Butler et al., 1995; Orszag-Sperber et al., 2000; Rouchy et al., 2001; Kartveit et al., 2019; Madof et al., 2019; Camerlenghi et al., 2019; Caruso et al., 2020; Raad et al., 2021). As morphological and seismic reflection studies at the Strait of Gibraltar documented a ~400 km long erosional trough connecting the Gulf of Cadiz (Atlantic Ocean) to the Mediterranean Sea, this scenario of a lowered Mediterranean Sea was promptly linked to the termination of the MSC (McKenzie, 1999; Blanc, 2002; Garcia-Castellanos et al., 2009, 2020). This conclusion has recently been reinforced by the discovery of vast chaotic deposits sitting at the claimed Miocene/Pliocene transition in the area of the Malta Escarpment-Ionian Abyssal Plain (Micallef et al., 2018, 2019; Spatola et al., 2020).

In more recent years, the desiccated basin model was challenged by the observation of deposits that are uniform in terms of sedimentology and stratigraphic architecture (Roveri et al., 2008a), ostracod content (Gliozzi et al., 2007; Stoica et al., 2016) and geochemistry (McCulloch and De Deckker, 1989; García-Veigas et al., 2018; Andreetto et al., 2021) throughout the Mediterranean marginal belt and of δD<sub>n-alkanes</sub> and δD<sub>alkenones</sub> sharing similarities with the coeval Atlantic Ocean and Black Sea, respectively (Vasiliev et al., 2017). A model of a (relatively) full Mediterranean Sea developed (Fig. 1c), where the debate mainly concerns the provenance of the hydrological fluxes and the resultant hydrochemical composition of the water mass. In this scenario, the Mediterranean was first, during substage 3.1, transformed into a new gypsum-precipitating basin filled with marine and continent-derived

<figure><img src="image_4.png" /><figcaption>**Fig. 1.** (a) Consensus chronostratigraphic model for the MSC events (Roveri et al., 2014a). Stage 3, here of interest, spans between 5.55 Ma and 5.332 Ma, the astronomical ages of the base of the Upper Gypsum Unit (following Manzi et al., 2009) and Trubi Formation (Van Couvering et al., 2000) in the Sicilian Eraclea Minoa section, respectively. CdB: Calcare di Base; PLG: Primary Lower Gypsum; RLG: Resedimented Lower Gypsum; UG: Upper Gypsum. (b), (c) Map of the Mediterranean region showing the two extreme and mutually exclusive paleoenvironmental scenarios proposed to have featured the Mediterranean during Stage 3 (see discussion in Chapter 7; modified after Krijgsman et al., 2018).</figcaption></figure>
<figure><img src="image_5.png" /></figure>

<figure><img src="image_6.png" /><figcaption>**Fig. 2.** Map of the Mediterranean Basin (modified from Lofi, 2018) showing: a) the location of the key intermediate and deep basins as well as physical thresholds that influenced the connectivity history of the Mediterranean; b) the onshore (i.e. basins and/or sections) and offshore (DSDP/ODP/Industrial drill sites) localities where deposits attributed to MSC Stage 3 have been studied. Mixed assemblages of Paratethyan-like ostracods and foraminifera are known from all mentioned onshore localities and some offshore locations (see text). The present-day spatial extent of the MSC seismic units, except for the Lower Unit, is also shown. The paleogeography of the (Eastern and Central) Paratethys and of the North Aegean domain is contoured after Van Baak et al. (2017) and Krijgsman et al. (2020a), respectively. W-E onshore localities: 1-6 Betic Cordillera (SE Spain); 1-Marbella and 2-Malaga basins (Guerra-Merchán et al., 2010); 3-Sorbas Basin (Roveri et al., 2009, 2019a); 4-Nijar Basin (Fortuin and Krijgsman, 2003); 5-Vera Basin (Fortuin et al., 1995); 6-Bajo Segura Basin (Soria et al., 2005, 2008a, 2008b); 7-Mallorca (Mas and Fornós, 2020); 8-Melilla Basin (Rouchy et al., 2003); 9-Boudinar Basin (Merzeraud et al., 2019); 10-Chelif Basin (Rouchy et al., 2007); 11-Sahel area (Frigui et al., 2016); 12-Aléria Basin and 13-Rhône Valley (Carbonnel, 1978); 14-Piedmont Basin (Dela Pierre et al., 2011, 2016); 15-Po Plain (Ghielmi et al., 2010, 2013; Amadori et al., 2018); 16-Fine Basin (Cava Serredi section; Carnevale et al., 2006a, 2008). 17-21 Apennine system: Romagna sections (17, Roveri et al., 1998), Trave section (18, Iaccarino et al., 2008), Maccarone section (19, Bertini, 2006, Grossi et al., 2008; Sampalmieri et al., 2010; Pellen et al., 2017), Colle di Votta (20)-Fonte dei Pulcini (21)-Stingeti (22) sections (Cosentino et al., 2005, 2012, 2013, 2018), Mondragone 1 well (23, Cosentino et al., 2006), Crotone Basin (24, Roveri et al., 2008a); 25-27 Sicily: Villafranca Tirrena (25) and Licodia Eubea (26) sections (Sciuto et al., 2018), Caltanissetta Basin (27, Manzi et al., 2009); 28-Corfu (Pierre et al., 2006); 29-Zakinthos (Karakitsios et al., 2017b); 30-Crete (Cosentino et al., 2007); 31-Cyprus (Rouchy et al., 2001; Manzi et al., 2016a); 32-Adana Basin (Radeff et al., 2016).</figcaption></figure>
waters (e.g. Manzi et al., 2009; Roveri et al., 2014c; Flecker et al., 2015; Vasiliev et al., 2017; García-Veigas et al., 2018; Grothe et al., 2020). Then, during substage 3.2, it became a brackish lake-sea comparable to the present-day Black Sea or Caspian Sea (Roveri et al., 2008a; Stoica et al., 2016; Andreetto et al., 2021), depending on whether a marine connection with the Atlantic was active (Manzi et al., 2009; Roveri et al., 2014b, 2014c; Flecker et al., 2015; Marzocchi et al., 2016; Vasiliev et al., 2017; García-Veigas et al., 2018) or not (e.g. McCulloch and De Deckker, 1989; Roveri et al., 2008a), and with a base-level fluctuating by hundreds of meters with precessional periodicity (Fortuin and Krijgsman, 2003; Ben Moshe et al., 2020; Andreetto et al., 2021). In the relatively full scenario, the revival of marine conditions is ascribed to either connectivity changes (Marzocchi et al., 2016) or to a moderate sea-level rise (Andreetto et al., 2021). In contrast, Carnevale et al. (2006a, 2006b, 2008, 2018) and Grunert et al. (2016), based on the recovery of fish remains ascribed to marine species, proposed that fully marine conditions were in force in the Mediterranean already at the end of substage 3.1.

After nearly 50 years of research on both onshore and offshore localities (Fig. 2), the observations backing up the competing desiccated and full-basin Mediterranean models remain extremely difficult to reconcile. Uncertainties regarding the chronostratigraphic framework of Stage 3 deposits, the origin and migration of its characteristic biota, the meaning of the data derived from the applied geochemical techniques and the relationship between the Mediterranean and its surrounding water bodies (i.e. Atlantic Ocean, Indian Ocean and Paratethys) all inhibit a clear understanding of the Mediterranean base-level and its hydrochemical structure.

In this paper we attempt to summarize all the existing, but heavily scattered, data resulting from ~50 years of cross-disciplinary studies with the aim of providing a comprehensive overview of the stratigraphic arrangement of Stage 3 onshore and offshore deposits, as well as of their sedimentological, paleontological, geochemical and seismic properties. Subsequently, we assemble the observations favoring both end-member scenarios of a relatively desiccated and relatively full Mediterranean. Finally, we focus on novel future analytical techniques and approaches that have the potential to constrain Mediterranean base-level during MSC Stage 3 as well as the changing hydrological fluxes and connectivity phases between the intra-Mediterranean basins and the neighboring Atlantic Ocean and Paratethyan domains as a mean of reconstructing the state of the art of the complex history of this enigmatic period of the Mediterranean history once and for all.

## 2. The terminal Stage 3 of the MSC

### 2.1. Historic overview of nomenclature and concepts

The final phase of the MSC (i.e. substage 3.2), also known as "Lago-Mare", finds its sedimentary expression in cyclically-arranged terrigenous and evaporitic sediments hosting unique faunal assemblages of ostracods, mollusks and dinoflagellate cysts (dinocysts). They are related, at species level, to those inhabiting, during the Miocene, the brackish basins of the Paratethys realm (e.g. Gliozzi et al., 2007; Stoica et al., 2016). But what exactly is the "Lago-Mare"? This widely employed expression in the MSC literature encompasses a variety of meanings that make its application doubtful and misleading. The root of the wording "Lago-Mare" is to be found in the Russian scientific literature of the late 1800s. Nikolai Andrusov (1890) used the corresponding Russian term with a geographical and chronological connotation in reference to the series of central-eastern European basins that during the Miocene turned from marine settings to desalinized semi-isolated lakes with an endemic fresh-brackish water biota association (e.g. Popov et al., 2006 and references therein). The original monograph of Andrusov (1890) was not widely available outside Russia, but his attendance of international conferences allowed his research to spread outside the Russian borders. From the publications of the French

geologists Suzette Gillet (Gillet, 1932, 1933) and Maurice Gignoux (Gignoux, 1936a) we can state with relative certainty that the original meaning of the word "Lago-Mare" (here reported with the French counterpart "Lac-Mer") had its provenance in the Russian literature:

"[...] An isolation of the basin, that became a brackish, isolated basin Then, a uniform fauna populated this immense lac-mer which was divided [...] into Pannonian basin, [...] Dacique Basin, and Euxin and Caspian basin [...]" (Gillet, 1932).

"[...] During the Volhynien (Sarmatique inferior) there was a lac-mer of uniform fauna which extended through all the eastern Europe. [...] and the fauna of the eastern regions of that huge lac-mer was completely differentiated [sic] from the one in the western regions. [...]" (Gillet, 1933).

"[...] The Pontien fauna is not anymore a fauna characteristic of an internal saline sea, as in the Sarmantien, but is a fauna of a "desalinated lagoon", a lac-mer, as the Russian geologists named it. [...]" (Gignoux, 1936b).

In the late 19th (Capellini, 1880) and 20th century (Ogniben, 1955; Ruggieri, 1962, 1967; Decima, 1964), late Messinian ostracod- and mollusk-bearing deposits in the Mediterranean were described at several Italian localities. Initially, the expressions "Congeria beds" (Capellini, 1880) and "Melanopsis beds" (Ruggieri, 1962) were used. Later on, Ruggieri (1967) pointed out the affinity of these faunal elements with those of the Pontian of the Paratethys. Consequently, he speculated on a feasible Paratethys-like paleoenvironmental configuration for the Mediterranean in the latest Messinian and he coined the Italian translation (i. e. "Lago-Mare") from the French "Lac-Mer" in reference to the shallow-water lakes claimed to be widely distributed across the Mediterranean. Progress in the 1970s in onshore and offshore exploration highlighted the temporally well-constrained distribution of the Paratethyan organisms in the Mediterranean (Carbonnel, 1978). On this premise, Hsü et al. (1978a) proposed to use "Lago-Mare" to "designate the latest Messinian oligohaline environment, postdating evaporite deposition and predating Pliocene marine sedimentation [...] in order to distinguish it from "lac mer" which, strictly speaking, was a Paratethyan environment". Notwithstanding the new definition, in various parts of the text they used "Lago-Mare" to refer to the Paratethyan lakes (pp. 1071-1072: "[...] The upper Messinian Mediterranean was floored by a series of desert basins, some with salt lakes, prior to inundation by the Lago-Mare."), thus giving rise to the confusion on how to use the term properly.

In the most recent stratigraphic overview of the MSC (Fig. 1a; Roveri et al., 2014a), the terminal MSC stage is called Stage 3, which is in turn subdivided into substages 3.1 and 3.2 (also termed Lago-Mare). Beside such a chronostratigraphic definition, the term "Lago-Mare" has also been used for a typical biofacies of the late Messinian Mediterranean (e. g. Fortuin et al., 1995; Gliozzi, 1999; Gliozzi and Grossi, 2008; Sciuto et al., 2018), for the pelitic beds encasing the Paratethyan-related fauna (i.e. a lithofacies; e.g. Fortuin and Krijgsman, 2003; Sciuto et al., 2018), as the name of an informal lithostratigraphic unit (usually distinguished by its fossil content) sandwiched between the Sicilian Upper Gypsum and the Arenazzolo Fm. (Fig. 4b; Clauzon et al., 2005; Londeix et al., 2007; Popescu et al., 2009; Bache et al., 2012) and to denote multiple (3 to 4) spilling events of the Paratethys into the Mediterranean (Clauzon et al., 2005, 2015; Popescu et al., 2007, 2009, 2015; Suc et al., 2011; Bache et al., 2012; Do Couto et al., 2014; Frigui et al., 2016; Mas and Fornós, 2020).

This being a review, we use the widely employed definition of the model of Roveri et al., 2014a) (Fig. 1a) and regard the Lago-Mare as a "phase of massive biota migration from the Paratethys realm, cyclostratigraphically constrained between 5.42 Ma and 5.332 Ma (Roveri et al., 2008a; Grossi et al., 2011), during which the Mediterranean sedimentary environments underwent an impressive freshening". Nevertheless, we call for caution in the use of this definition of "Lago-Mare" in future studies, since 5.42 Ma as the (astronomical) age of the first entrance of Paratethyan organisms into the Mediterranean is likely
to be incorrect (see subsection 5.1) and evidence of ‘impressive fresh-
ening’ are already present much earlier (e.g. at Eraclea Minoa; Vasiliev
et al., 2017; García-Veigas et al., 2018).

## 2.2. Development of a chronostratigraphic framework

Issues of the timing and duration of the MSC only began to be tackled
in the 1990s, in parallel with discussion concerning the nature of its
extreme paleoenvironments (Schmalz, 1969; Hsü et al., 1973a, 1973b,
1973c, Hsü et al., 1978a, 1978b; Nesteroff, 1973; De Benedetti, 1982).
While published models (Butler et al., 1995; Clauzon et al., 1996;
Krijgsman et al., 1999a; Rouchy and Caruso, 2006) mostly converged on
the (astronomical) age of the marine replenishment at the beginning of
the Pliocene (5.332 Ma; Van Couvering et al., 2000), there were dis-
agreements about the age of the onset of the MSC (synchronous vs
diachronous) and of specific events within it (see discussion in Roveri
et al., 2014a). Among these, the work of Krijgsman et al. (1999a) has
obtained wide consensus. Their cyclostratigraphic tuning and correla-
tion of continuous and bio-magnetostratigraphically constrained pre-
evaporitic sections in Spain (Sorbas), Sicily (Gibliscemi/Falconara)
and Greece (Metochia) resulted in a synchronous age of 5.96±0.02 Ma
for the MSC onset (later refined to 5.97 Ma by Manzi et al., 2013). The
astronomical ages for the onset (Krijgsman et al., 1999a) and termina-
tion (Van Couvering et al., 2000) of the MSC are not contentious since
the characteristic sedimentary cyclicity and sediments’ properties (e.g.
color of the lithologies and biota content) of the pre- and post-MSC
successions fit robustly with the insolation curve (see also Van der
Laan et al., 2006 and Topper and Meijer, 2015).

The cyclic arrangement of the MSC sediments (Fig. 3a) led scientists
to interpret that the same cyclostratigraphic approach could be used to

gain precise dates for events within the MSC (e.g. Hilgen et al., 1995; Vai, 1997; Krijgsman et al., 1999b, 2001), bypassing the challenge posed by the unsuitability of the classic biomagnetostratigraphic tools for the MSC successions. Characteristic interference patterns of eccentricity and precession have been tentatively recognized in the Sicilian Eraclea Minoa section (see subsection 3.8; Van der Laan et al., 2006). However, clear orbital signals are typically poorly expressed in MSC records and, when they are present, like in Sicily, they are not (vertically) repeated with sufficient frequency to establish clear phase relations with the astronomical cyclicity. For this reason, the simple counting of cycles with no analysis of cyclostratigraphic pattern in proxy records has mostly been employed as a correlation method (Roveri et al., 2008a; Manzi et al., 2009; Manzi et al., 2016a; Cosentino et al., 2013).

The age of the base of Stage 3 is largely determined by correlating the sedimentary cycles of the Upper Gypsum unit (UG) at Eraclea Minoa (Sicily) with the astronomical curve La2004 (Laskar et al., 2004). The UG sedimentary cyclicity consists of alternating gypsum and mudstone beds of variable thickness (Figs. 5g-i; see subsection 3.8). Precessional variation of the Mediterranean freshwater budget tied tightly to the African monsoon and Atlantic storms are the drivers interpreted to lie behind the gypsum-mudstone cycles (e.g. Marzocchi et al., 2015, 2019; Simon et al., 2017). Variations of the freshwater discharge cause the pycnocline to shift vertically, resulting in brine concentration and gypsum precipitation during to the arid/dry phases of the precession cycles (precession maxima-insolation minima) and brine dilution and mudstone deposition during the humid/wet phases (precession minima-insolation maxima) (Van der Laan et al., 2006; Manzi et al., 2009). Two different tuning options exist in literature (Van der Laan et al., 2006 versus Manzi et al., 2009; Fig. 3a):

<figure><img src="image_7.png" /><figcaption>**Fig. 3.** (a), (b) Available astronomical tunings to astronomic curves of climatic precession (P), 100 kyr eccentricity (E) and 65°N insolation curve (I) of Laskar et al. (2004) of the lithological cyclicity of onshore Stage 3sections (a) and of the seismic cycles and/or well logs (gamma ray and resistivity) of the MU in the Levant Basin (b). Tunings of onshore sections in (a) are carried out downward from the M/P boundary (conformable in all sections). Astronomically-tuned glacial (even numbers) and interglacial (odd numbers) stages (i.e. TG) as defined by Hodell et al. (1994) are also indicated.</figcaption></figure>
1. Van der Laan et al. (2006) tentatively recognized sedimentary patterns that they correlated with the astronomical curves by using the same phase relationships between the sedimentary cycles and the astronomical cycles as are seen in Plio-Pleistocene sapropel-bearing marine successions of the Mediterranean (Hilgen, 1991). The four closely spaced gypsum beds III to VI were regarded as a cluster, i.e. the sedimentary expression of a 100 kyr eccentricity maximum (Hilgen, 1991; Strasser et al., 2006), whereas the preceding and following evaporite-free marly interval were attributed to a phase of low-amplitude precession oscillations caused by a 100 kyr eccentricity minimum (Fig. 3a). Tuning downward from the conformable Miocene/Pliocene boundary (Fig. 6d) and arguing that the precession peak at ~5.38 Ma, which has an extremely low amplitude, is not expressed in the sedimentary record, Van der Laan et al. (2006) correlated gypsums III to VI with the four successive precession/insolation peaks of the 100 kyr eccentricity maximum dated around 5.44 Ma and the overlying and underlying gypsum-free marly interval fell within 100 kyr eccentricity minimum cycles (Fig. 3a, right log). This tuning resulted in an astronomical age of ~5.51 Ma for the first gypsum bed in their log (i.e. gypsum II in the log of Manzi et al., 2009), and an approximate duration of 175 kyr for Stage 3 as whole.

2. An alternative tuning by Manzi et al., 2009; Fig. 3a, left log) argued that every precessional/insolation peak must have an expression in the rock record. Manzi et al. (2009) agreed with the solution of Van der Laan et al. (2006) on the sedimentary inexpressiveness of the (barely visible) insolation minima peak at ~5.38 Ma. However, these authors considered the insolation minima peaks immediately above and below of too low amplitude to promote the conditions required for gypsum precipitation, but also too high not to have some sedimentary expression. They therefore identified sandstone horizons VI' and VI'' as the sedimentary response to these weak insolation/precession signals. The addition of two precessional cycles (i.e. a total of 9) resulted in an astronomical solution that was adjusted one precessional cycle lower than that of Van der Laan et al. (2006), translating into an age of 5.53 Ma for the base of the UG and a total duration of ~200 kyr for Stage 3. But the more conspicuous difference between the two astronomical solutions discussed lies in the timing at which gypsum precipitation occurred, restricted to the 100 kyr eccentricity maxima according to Van der Laan et al. (2006), extended to the 100 kyr eccentricity minima by Manzi et al. (2009).

An age of 5.53 Ma for the first gypsum bed was also obtained by the astronomical tuning of the Upper Gypsum in Cyprus (Manzi et al., 2016a), but there the tuning is performed just by following the recognition, from the base up, of 6 gypsum beds just like in Sicily and therefore arguing for a bed-to-bed correlation with the Sicilian gypsums I-VI. In the consensus model of Roveri et al. (2014a) the base of Stage 3 coincides with the base of the Sicilian UG, placed by Manzi et al. (2009) at 5.55 Ma (Fig. 1A). However, in the model of Manzi et al. (2009) this age is attributed to a cumulate gypsum horizon interpreted as laterally equivalent of the Halite (i.e. Stage 2), and therefore implying the kickoff of Stage 3 at 5.53 Ma (Fig. 3a).

The post-evaporitic successions of the Romagna (Cussercoli and Sapigno sections; Roveri et al., 1998) and Marche (e.g. Trave and Maccarone sections; Iaccarino et al., 2008; Cosentino et al., 2013) areas provided evidence that led to the splitting of Stage 3 into substage 3.1 and 3.2. In the resulting composite section (Roveri et al., 2008a), a shift in the sedimentary facies and stacking pattern is observed (see description in subsection 3.7). Correlation of the sedimentary cyclicity in Romagna was from the (conformable) base of the Pliocene downwards (or from an U-Pb-dated ash layer upward; Cosentino et al., 2013) and linked three fluvial conglomerates and two black mudstone layers of unknown sedimentological significance to the arid phases of the precession cycles (Fig. 3a; Roveri et al., 2008a). The greater thickness of the oldest conglomerate was possibly assumed to be evocative of an oscillation of the amplitude of the corresponding precession minima peak

rather than the amplitude of the peaks responsible for the formation of the other facies. This approach resulted in an age of 5.42 Ma for the first conglomerate (i.e. the substage 3.1/3.2 transition; Fig. 5g) and an approximate duration of 90 kyr for substage 3.2 (the Lago-Mare phase). The same astronomical age is obtained by tuning the Upper Member of the Feos Formation in the Nijar Basin (Omodeo-Salé et al., 2012), where four pelite-conglomerate cycles plus one sandstone capped by the Miocene/Pliocene boundary mark the interval attributed to Stage 3.2 (Fortuin and Krijgsman, 2003).

Although the substage 3.1/3.2 transition is linked to a major Mediterranean-scale hydrological re-organization possibly coinciding with the migration of the Paratethyan biota (Roveri et al., 2008a; Grossi et al., 2011), the facies change used for its definition is hardly recognizable elsewhere (see Chapter 3). As such, other tools have been used to equip fragmentary and/or lithological cyclicity-lacking sections with an age model: the (highly controversial) ostracod biozonation (see sub-section 5.1; e.g. Stoica et al., 2016; Karakitsios et al., 2017a; Cosentino et al., 2018; Caruso et al., 2020) and the astronomical tuning of magnetic susceptibility records (e.g. Fonte dei Pulcini section, Central Apennines; Cosentino et al., 2012).

Comparison of Atlantic oxygen isotope records (Van der Laan et al., 2005, 2006) and the chronostratigraphy of Roveri et al. (2014a) revealed that Stage 3 sedimentation started during a prominent global eustatic lowstand associated with oxygen isotope (glacial) stage TG12, followed by a latest Messinian deglacial interval which comprised multiple obliquity- and possibly precession-forced global eustatic phases. As documented by Hodell et al. (2001) (later revised by Drury et al., 2018), Van der Laan et al. (2006) and Roveri et al. (2014a), the marine replenishment of the Mediterranean did not coincide with any major deglaciation, so non-eustatic causes of the Zanclean megaflood hypothesis are required.

## 3. Onshore domain: key sections, sedimentary expression and faunal content

### 3.1. *The Alborán region*

The westernmost outcrops of Stage 3 deposits in the Mediterranean are located in the Alborán region, close to the present-day Strait of Gibraltar (Fig. 2b). MSC deposits on the margins of this region are poorly developed, possibly because of a late Tortonian uplift that raised the margins above the Mediterranean water level (López-Garrido and Sanz de Galdeano, 1999). Near Malaga, however, two facies associations consisting of m-thick conglomerate-sandstone beds alternating with laminated pelites are documented in the Rio Mendelín section (informally referred to as “LM unit”; Guerra-Merchán et al., 2010) and attributed to (part of) the Lago-Mare phase (Fig. 4a) based on their paleontological content. These sediments are squeezed between the Paleozoic basement units, with an erosive contact and associated angular unconformity, and the Pliocene, from which they are separated by another erosional surface draped by conglomeratic accumulations (Fig. 6a). A well-preserved and diverse *in situ* Paratethyan-type ostracod and molluscan fauna (i.e. *Lymnocardiinae* and *Dreissenidae*) typical of shallow waterbodies (up to 100 m deep; Grossi et al., 2008; Gliozzi and Grossi, 2008) with low salinities (5-18‰) is reported from the pelitic units (Guerra-Merchán et al., 2010). The overlying Pliocene in the deeper depocenters starts with 30 m-thick littoral conglomerates with marine mollusks passing progressively upwards into deeper water facies, while fan deltas developed at the basin margins (López-Garrido and Sanz de Galdeano, 1999; Guerra-Merchán et al., 2010, 2014). Notably, the overall thickness of the Pliocene deposits reaches 600 m. The detailed regional studies by López-Garrido and Sanz de Galdeano (1999) and Guerra-Merchán et al. (2014) concluded that accommodation space was created during (Zanclean) sedimentation by local fault-driven subsidence, and that movement on these faults only reversed at the end of the Zanclean causing uplift.
<figure><img src="image_8.png" /><figcaption>**Fig. 4.** Schematic overview of different chronostratigraphic models for some of the Messinian successions presented in Chapter 3. Note the large controversies in timing, duration and chronostratigraphic position of the main erosion phase between models in (a) and (b). Models in (a) follow the recently established MSC chronostratigraphic model of Roveri et al. (2014a), according to which the Mediterranean base-level dropped and halite deposited on sea floor during Stage 2 and the Upper Gypsum/Upper Evaporites-Lago-Mare sequence followed. Models in (b) were proposed following the alternative scenario of Clauzon et al. (1996, 2005), which envisaged two Lago-Mare episodes (LM1 and LM3) that occurred before and after the main Mediterranean drawdown event, during which LM2 was deposited in the deep desiccated basins (Do Couto et al., 2014; Popescu et al., 2015; see Roveri et al., 2008c and Grothe et al., 2018 for further explanations). Note, in (b), the shifting of the position of the main erosional phase in Sicily through time as well as the time of the marine replenishment in the Apennines.</figcaption></figure>

An alternative scenario, based on the finding of (a few) specimens of the nannofossil *Ceratolithus acutus*, ascribed to the LM unit of *Guerra-Merchán et al. (2010)* to the earliest Zanclean (Fig. 4b; Do Couto et al., 2014).

On the southern Alborán margin in Morocco, latest Messinian deposits are reported from the Boudinar and Melilla basins (Fig. 2b). Up to 100 m-thick chaotic deposits containing selenite gypsum fragments, azoic conglomerates, sandstones yielding planktic foraminifera and nannofossils and lacustrine limestones are capped by early Pliocene marine marls (Rouchy et al., 2003; Azdimousa et al., 2006; Cornée et al., 2016; Merzeraud et al., 2019). Due to their stratigraphic position, these continental to lacustrine deposits are interpreted as the local expression of the Lago-Mare phase (Cornée et al., 2016) or alternatively as Zanclean successions (Azdimousa et al., 2006).

## 3.2. Algeria

The Chelif Basin in Algeria (Fig. 2b) displays the typical marginal Messinian succession comprising Tortonian to lower Messinian blue marls, diatomite-bearing sediments (Tripoli unit), cyclically-arranged primary evaporites (13 couplets), ostracod-rich post-evaporitic deposits and Zanclean foraminiferal marls (Rouchy et al., 2007). The post-evaporitic sediments show a great lateral variability in both thickness (from few meters up to 125 m) and facies. They are mainly dominated by terrigenous clastic lithologies, associated in the marginal areas with sandy carbonates and stromatolitic limestones. A mixed faunal assemblage of non-marine (Paratethyan-like ostracods) and marine (benthic and planktic foraminifera) organisms is present, showing an increase in ostracod species diversity from the bottom to the top (Rouchy et al., 2007).

## 3.3. Neogene basins of the Eastern Betics (Spain)

The external Neogene basins (Sorbas, Nijar, Vera and Bajo Segura) of the eastern Betic Cordillera (SE Spain; Fig. 2b) represent an important laboratory for understanding Messinian events. In particular, the Sorbas and Nijar basins preserve two allegedly continuous successions spanning the entire MSC (e.g. Roveri et al., 2009; Omodeo-Salé et al., 2012). The two basins are similar in many respects. Their stratigraphic organization, for example, suggests they were connected for much of the late

Miocene up until MSC Stage 1 (Fortuin and Krijgsman, 2003), which is represented by the gypiferous Yesares Member (e.g. Lu, 2006). However, facies differences are prominent in the Stage 3 formations according to the chronostratigraphic frameworks of Roveri et al. (2009) for the Sorbas Basin and Omodeo-Salé et al. (2012) for the Nijar Basin (Fig. 4a). Lithostratigraphically, two members are discerned between the Yesares Member and the basal Zanclean: the Sorbas and Zorreras members in the Sorbas Basin (Figs. 4a, 5a) and the lower and upper members of the Feos Fm. in Nijar (Figs. 3a, 4a; Roep et al., 1998; Krijgsman et al., 2001; Fortuin and Krijgsman, 2003; Braga et al., 2006; Roveri et al., 2009, 2019a; Omodeo-Salé et al., 2012).

The Sorbas Member (see Roep et al., 1998 and Aufgebauer and McCann, 2010 for a more detailed sedimentological description) consists of three overlapping coarsening-upward depositional sequences made of offshore clays and marls passing upward into shelf muds and coastal sandstone bodies. Still unclear is the chemistry of the subaqueous environment during the formation of the Sorbas Member and the provenance of the water fluxes. These shallow-water deposits are conformably replaced upward by the Zorreras Member that comprises alternations of reddish siltstones and sandstones (Fig. 5a) organized in five (or eight) lithological cycles expressing continental environments (Martín-Suárez et al., 2000; Aufgebauer and McCann, 2010). Up to four lenticular white limestone beds bearing brackish Paratethyan-like ostracods (*Cyprideis*, *Loxocorniculina djafarovi* and freshwater species of the family Limnocytheridae), bivalves and *Chara* oogonia (Roep and Harten, 1979; Aufgebauer and McCann, 2010) are found interrupting the fluvialite sequence (Fig. 5a) and are linked to either episodic flooding by local rivers (Braga et al., 2006; Aufgebauer and McCann, 2010) or episodic Mediterranean ingressions (Fortuin and Krijgsman, 2003; Andreetto et al., 2021). A correct interpretation of the paleo-depositional environment of these limestone beds is crucial for the discussion concerning the Mediterranean base-level position during the Lago-Mare phase. In fact, if the Sorbas Basin was relatively shallow during Zorreras deposition (50-100 m; Roveri et al., 2019a, 2020), repeated and sudden Mediterranean incursions would indicate that the Mediterranean Basin was relatively full and that its base level was oscillating, possibly with precessional periodicity (Andreetto et al., 2021). The contact between the Zorreras Mb. and the underlying near-shore Pliocene (<50 m depositional paleodepth; Roveri et al., 2019a) in the Sorbas Basin is conformable and expressed differently around the
<figure><img src="image_9.png" /><figcaption>**Fig. 5.** Sedimentary expression of Stage 3 from selected onshore Mediterranean localities. (a) Photograph from the Sorbas Basin showing the red continental sediments of the Zorreras member with intercalated white limestones (white arrows; from Andreetto et al., 2021). The conformable resting of the Zorreras Mb. above the Sorbas Mb. and underneath the Gochar Fm. of Pliocene age is also appreciable. Car for scale. (b) One typical lithological (and precessional) cycle of the Upper Mb. of the Feos Fm. in the Nijar Basin, here constituted by an ostracod-bearing, white and laminated mudstone bed overlain by an azoic fluvial sandstone (courtesy of Anne Fortuin). (c) Panoramic view of the Cuevas del Almanzora section (from Andreetto et al., 2021). Red rectangle indicates the position of the section straddling the Messinian (M)/Zanclean (Z) transition and studied by Fortuin et al. (1995), Stoica et al. (2016), Caruso et al. (2020) and Andreetto et al. (2021). Buildings for scale. (d) The sub-unit a of the Piedmont Basin composed of azoic grey mudstones grading into yellowish, mammal-rich overbank deposits. (e) WNW-ESE seismic profile in the Po Plain showing incised valleys filled during Stage 3 by suggested clastic deposits and sealed by deep-water turbidites in the Zanclean (modified from Amadori et al., 2018). (f) Typical aspect of the di Tetto/San Donato Formation in the Northern Apennines composed by grey mudstones (detail in the inset) with interbedded sandstone bodies (white arrows). The picture is taken from the Maccarone section. (g) The di Tetto Fm.-Colombacci Fm. transition in the Cusercoli area (Eastern Romagna, Fig. 2b), defined by the facies change underlined by the appearance of a fluvial conglomerate. This lithostratigraphic boundary also corresponds to substage 3.1/3.2 boundary of Roveri et al. (2014a). (h), (i), (l) Lithological cycles of the Upper Gypsum Unit in Eraclea Minoa (h), Siculiana Marina (i) and Polemi (l) sections. Cycles are several m-thick and primarily composed by beds of primary gypsum alternating with mudstones bearing Paratethyan ostracods (at least in Eraclea Minoa).</figcaption></figure>
<figure><img src="image_10.png" /><figcaption>**Fig. 6.** Photographs of the Miocene/Pliocene boundary (yellow lines) from selected onshore Mediterranean localities. (a) Erosive M/P transition in the Mendelín section (Malaga Basin). Note the conglomeratic lag draping the erosional surface and sharply overlain by foraminifera-rich marls. (b) Conformable stratigraphic contact between the uppermost Messinian sandstone of the Feos Fm. and the Zanclean biocалcarenites of the Cuevas Fm. in the Barranco de los Castellones section, Nijar Basin (hammer for scale; modified from Andreetto et al., 2021). (c) The Messinian/Zanclean boundary in the Pollenzo section (Piedmont Basin) marked by a characteristic black layer interbedded between Paratethyan ostracods-rich mudstones and marine foraminifera-rich marls (modified from Dela Pierre et al., 2016). (d) Uppermost segment of the Eraclea Minoa section (Caltanissetta Basin, Sicily) displaying the (non erosive) contact between the Pliocene Trubi Formation above and the sandy Arenazzolo Formation below. The inset is a close view of the transition, which occurs above a ~50 cm-thick burrowed mudstone horizon rich in Paratethyan ostracods and marine foraminifera. (e) Lago-Mare sediments in the Kalamaki section (Zakynthos) unconformable, through an erosional surface (i.e. the Messinian Erosional Surface, MES), over the PLG unit and also unconformable beneath the Trubi Fm. (modified from Karakitsios et al., 2017b). (f) Close view of the M/P boundary in the Pissouri Basin, where the foraminifera-rich Trubi marls lie above a black layer (paleosol according to Rouchy et al., 2001).</figcaption></figure>

basin, ranging from a bivalves-rich bed overlain by a yellow, fossiliferous calcarenite floored by a gravelly lag deposit (Mather et al., 2001) to a grey marl horizon with marine foraminifera assemblages followed by a second shell-rich bed (Roveri et al., 2019a). Similar to the situation in Malaga, the rare identification of *Ceratolithus acutus* in sediments of the continental Zorreras Mb. led Clauzon et al. (2015) to put forward an alternative chronostratigraphic and paleoenvironmental interpretation for the Sorbas MSC succession, shifting the Zorreras Mb. into the Pliocene (Fig. 4b) and thus associating the presence of brackish Paratethyan-like ostracods with exchanges between the Mediterranean and Paratethys following the Mediterranean re-filling, at high sea level.

In the Nijar Basin (Fig. 2b), the latest Messinian Feos Formation is bracketed at the base and top by an erosional surface along the basin margins and its correlative conformity in the deeper parts (Fig. 3a; Fortuin and Krijgsman, 2003; Aguirre and Sánchez-Almazo, 2004; Omodeo-Salé et al., 2012). The Lower Feos Member consists of azoic, graded and locally slumped siliciclastic-carbonate beds alternating with gyparenites and gypsiltites and including a laterally continuous Mn-rich bed (Fortuin and Krijgsman, 2003; Omodeo-Salé et al., 2012). In the basin center (e.g. Barranco de los Castellones section; Fig. 3a) the Upper Feos member comprises four complete lithological cycles of m-thick conglomerate to sandstone beds alternating with laminated pelites (Fig. 5b), and one incomplete cycle, which only consists of a sandstone horizon conformably capped by the Pliocene Cuevas Fm. (Fig. 6b; Fortuin and Krijgsman, 2003). A rich fauna of mixed brackish ostracods and marine foraminifera is found in all four pelitic beds (Bassetti et al., 2006). Its origin is questionable. These ostracods were regarded as endemic to the Mediterranean and inhabiting endorheic lakes by Bassetti et al. (2006). However, later they were shown to have been misidentified and were instead considered Paratethys-derived by Stoica et al. (2016; see subsection 5.1). Planktonic and deep-water benthic foraminifera are widely considered reworked by Fortuin and Krijgsman

(2003), Bassetti et al. (2006) and Omodeo-Salé et al. (2012), in place by
Aguirre and Sánchez-Almazo (2004).

In the Vera Basin (Fig. 2b), in situ gypsum deposits are missing because of widespread erosion or non-deposition and MSC deposits are only represented by ~12 m of laminated varicolored marly clays (Unit 2 Fig. 4a), which are best exposed in the Cuevas del Almanzora section (Fortuin et al., 1995; Fig. 5c). These clays contain a well-preserved and diversified in situ fauna of Paratethyan-like ostracod and shallow-water, benthic foraminifera mixed with physically reworked (mostly from the lower Messinian Abad marls) planktic and deep-water benthic foraminifera (Fortuin et al., 1995; Stoica et al., 2016; Caruso et al., 2020). The marly clays are assigned by Stoica et al. (2016) and Caruso et al. (2020) to (roughly) the whole late Messinian Lago-Mare phase (Fig. 4a) based on the ostracod biozonation of Grossi et al. (2011) and are considered to represent either sedimentation in an isolated lake subject to base-level and salinity fluctuations (Caruso et al., 2020) or deposition in a coastal lagoon that was connected to the water mass filling the open Mediterranean (Stoica et al., 2016; Andreetto et al., 2021). Similar to Malaga, these sediments are topped by an erosive surface draped by a conglomeratic accumulation which is overlain by the open marine fauna-rich sediments of the basal Zanclean (Fortuin et al., 1995; Caruso et al., 2020). This erosion feature likely indicates that the Miocene/Pliocene transition followed a base-level lowstand in the Vera Basin.

Stage 3 deposits (Garrucha Fm.) in the easternmost basin of the Betic
Cordillera, the Bajo Segura Basin (Fig. 2b), are bounded below and
above by two erosional surfaces related to lowered Mediterranean base-
levels and discontinuously present due to the widespread fluvial erosion
that occurred at the Miocene/Pliocene boundary (Soria et al., 2005,
2008a, 2008b). The Garrucha Fm. shows a maximum thickness of 100 m
in its type section (Soria et al., 2007, 2008b). It consists of 20-50 cm
thick sandstone bodies interrupting a dominantly marly succession
deposited in a subaqueous environment inhabited by Cyprideis sp. and
euryhaline, shallow-water benthic foraminifera (Ammonia beccarii, Elphidium granosum, Elphidium macellum, Haynesina germanica and Quinqueloculina laevigata). Planktic foraminifera are also observed and for a long time were considered to be physically reworked (Soria et al., 2005, 2008b). However, some stratigraphic levels contain dwarf tests of long-ranging taxa such as *Globoturborotalita decoraperta*, *Globigerina bulloides*, and *Neogloboquadrina spp.* which recently have been interpreted as being in-situ mostly due to the absence of notable signs of reworking (Corbí and Soria, 2016). Among these dwarf taxa is *Neogloboquadrina acostaensis* (dextral; Corbí and Soria, 2016). Since this group is mainly dextral in the latest Messinian Atlantic successions (e.g. Sierro et al., 1993; Bassetti et al., 2006), this may indicate that Atlantic inflow to the Mediterranean occurred during the late Messinian and the base level of the Mediterranean was high enough to reach the marginal Bajo Segura Basin. The Miocene/Pliocene boundary is, once again, marked by an erosional surface which outlines up to 200 m deep paleovalleys engraved down into the pre-MSC sediments and filled with conglomerates and sandstones of claimed coastal and shallow marine environments (Soria et al., 2005, 2008b; García-García et al., 2011; Corbí et al., 2016).

### 3.4. Mallorca

Mallorca, which constitutes an emerged segment of the Balearic Promontory (Fig. 2), does not expose the classical MSC evaporite sequence. Instead, two main MSC-related units are found above late Tortonian-Messinian reefal carbonates (Reef Complex Unit) and beneath the Pliocene: the Santanyí limestones and the Ses Olles Formation (Mas and Fornós, 2020 and references therein). The Santanyí limestones are microbialites and oolite-dominated sediments in which a baleen whale neurocranium has been found (Mas et al., 2018a). This unit was interpreted either as a Terminal Carbonate Complex (TCC) laterally equivalent to the Primary Lower Gypsum (PLG) which has been drilled in the deeper parts of the bay of Palma (Mas and Fornós, 2020) or as time-equivalent to the Reef Complex Unit (e.g. Arenas and Pomar, 2010; Suárez-González et al., 2019). The Ses Olles Formation consists of marls, sandy-marls and marly-calcareous lacustrine deposits rich in in-situ freshwater *Chara* spp., brackish water Paratethyan-like mollusks and ostracods and littoral benthic foraminifera (*Elphidium* sp., *Ammonia* sp.).

The upper contact of the Ses Olles Formation with the Pliocene corresponds to an erosional ravinement surface draped by a transgressive lag of coastal deposits usually containing coquinas and/or conglomerates (Mas, 2013, 2015; Mas and Fornós, 2020). The lower contact of the Ses Olles Formation with the Santanyí limestones is sporadically marked by a well-developed reddish paleosol (Mas, 2013, 2015; Mas and Fornós, 2020), which indicates that a (unquantified) period of subaerial exposure occurred before the emplacement of the Ses Olles Fm. However, in their more recent study, Mas and Fornós (2020) surprisingly conclude that the Ses Olles Formation has a conformable contact with the Santanyí limestones, ascribed to part of Stage 1. This led Mas and Fornós (2020) to conclude that the emplacement of the Ses Olles Fm. pre-dated the MSC peak and that the erosional surface marking the Miocene/Pliocene boundary is associated with a 270 kyr hiatus linked to the main MSC base-level drawdown (Fig. 4b). This conclusion is, however, in disagreement with the unconformity at the base of the Ses Olles Fm., which instead points to the deposition of the Ses Olles Fm. (and therefore to the arrival of the Paratethyan fauna in Mallorca) at some point during Stage 3 of Roveri et al. (2014a).

### 3.5. Piedmont Basin

The Piedmont Basin (NW Italy) contains the northernmost record of the MSC (Fig. 2b). The terminal MSC sediments (i.e. the Cassano Spinola Conglomerates Fm.) overlay pre-MSC units, the PLG deposits (Gessoso Solifera Fm.) or reworked evaporites (Valle Versa chaotic complex, VVC) and underlie the Zanclean marls of the Argille Azzurre Fm. (Dela

Pierre et al., 2011).

The Cassano Spinola Conglomerates is split in two sub-units by Dela Pierre et al. (2016). Sub-unit a consists of azoic grey mudstones turning to yellowish silty mudstones (Fig. 5d) typified by in situ root traces, paleosols and mud cracks and including three/four intercalated lens-shaped, cross-bedded conglomeratic layers (Ghibaudo et al., 1985; Dela Pierre et al., 2011, 2016). Abundant land plant leaves and a diverse terrestrial vertebrate fauna are found in the yellowish siltstones, which have been interpreted as overbank deposits (Harzhauser et al., 2015; Colombero et al., 2017 and references therein). In this continental interval, a low-diversity fish fauna consisting of otoliths of marine and Paratethyan species is found (Grunert et al., 2016; Carnevale et al., 2018; Schwarzhans et al., 2020). These otoliths were Sr-dated to the early-middle Miocene (Grunert et al., 2016). Nevertheless, they were concluded not to be physically reworked, but rather to have been transported by large marine predators, therefore implying a Piedmont Basin-(marine) Mediterranean connection was in force (Grunert et al., 2016; see subsection 5.6). Sub-unit b (i.e. Strati a Congeria sensu Sturani, 1973) is made of grey mudstones bearing a mixture of in-situ brackish water mollusks (Sturani, 1973; Esu, 2007) and ostracods (Trenkwalder et al., 2008) of Paratethyan affinity along with physically reworked foraminifera and calcareous nannofossils (Trenkwalder et al., 2008; Violanti et al., 2009). The transition to the Pliocene Argille Azzurre Fm. is sharp above a characteristic black and azoic sandy layer (Fig. 6c) rich in terrigenous and intrabasinal (i.e., glaucony and phosphates) grains and disarticulated valves of both brackish-water and continental bi-valves, but barren of in-situ fossils (Trenkwalder et al., 2008). The occurrence, at its top and directly below the Argille Azzurre Fm., of abundant *Thalassinoides* trace fossils filled with Pliocene sediments led Trenkwalder et al. (2008) and Dela Pierre et al. (2016) to interpret the top surface of this layer as an omission surface. This surface indicates a period of basin starvation (and therefore a hiatus) due to a sudden increase in water-depth, ascribed by Trenkwalder et al. (2008) to the Zanclean reflooding. This hiatus may have lasted for only part of the late Messinian (Violanti et al., 2009; Dela Pierre et al., 2016) or may have endured into the Pliocene (Trenkwalder et al., 2008).

### 3.6. Po Plain

To the east, the Messinian sediments in the Piedmont Basin disappear beneath the km-thick Plio-Quaternary succession of the Po Plain-Adriatic Foredeep (PPAF; Fig. 2a). By definition of Ghielmi et al. (2010) and Amadori et al. (2018), the PPAF includes two main elongated depocenters enclosed within the northern Apennines to the South and the Southern Alps to the North: the easternmost portion of the Po Plain and the whole present-day northern Adriatic Sea. Here, for simplicity, we include in the definition of PPAF also its westernmost depocenters of the Western Po Plain Foredeep.

The Messinian-Pleistocene sedimentary sequence, studied through the integration of seismic and borehole observations, is mostly represented by thick sequences of turbidite deposits in the foreland depo-center passing towards the margins, to fluvial and deltaic systems related to the proximity of the marginal thrust-fold-belts (Cipollari et al., 1999; Ghielmi et al., 2010, 2013; Rossi et al., 2015a; Rossi, 2017). During MSC Stage 1, primary evaporites and dolomiticites were deposited in some shallow-water settings, while evaporitic deposition was inhibited in the deep-water settings, where it was replaced by deposition of anoxic mudstones (Ghielmi et al., 2010). Instead, the post-evaporitic deposits consist of large thicknesses (up to 1 km) and volumes of coarse-grained clastics (LM1 and LM2 of Rossi and Rogledi, 1988; ME3 or Fusignano Fm. of Ghielmi et al., 2010; ME4 of Ghielmi et al., 2013; ME3b and possibly ME3a of Rossi et al., 2015a). Several authors (Ghielmi et al., 2010, 2013; Rossi et al., 2015a; Amadori et al., 2018; Cazzini et al., 2020) showed that these post-evaporitic sediments are the infilling of ca. N-S and NW-SE trending, V-shaped valleys (Fig. 5e). These valleys were carved at least as far as 50 km into the Alps, to a
depth up to 1 km into the pre- and syn-evaporitic Messinian deposits and nicely shape the present-day river network of the southern Alps (Amadori et al., 2018).

Different mechanisms for the incision have been proposed, with major implications for the desiccated vs full Mediterranean controversy (Figs. 1b-c). Ghielmi et al. (2010, 2013), Rossi et al. (2015a), Amadori et al. (2018) and Cazzini et al. (2020) ascribed the valley incision along the PPAF northern margin to fluvial erosion, whose basinward shifting was triggered by the Stage 2 Mediterranean drawdown, estimated to have been around 800-900 m (Ghielmi et al., 2013; Amadori et al., 2018). In this case, Stage 3 deposition in the PPAF occurred in endorheic lakes fed by the Alpine rivers and kept isolated until the Zanclean, when the sudden sea-level rise following the Zanclean reflooding was enough to bypass morphological highs (e.g. Gargano-Pelagosa and/or Otranto paleosills) located in the southern Adriatic foredeep (Fig. 2a; see Pellen et al., 2017; Amadori et al., 2018; Manzi et al., 2020). Conversely, Winterberg et al. (2020) suggested that the over-deepened valleys on the southern slope of the Alps are related to Pleistocene glacial erosion. Although Winterberg et al. (2020) do not address the paleoenvironment during the Messinian, this interpretation does not rule out the possibility that (at least part of) Stage 3 sedimentation occurred in a PPAF connected to the Mediterranean water mass and that no catastrophic reflooding occurred at the Miocene/Pliocene boundary. The conclusion of a non-catastrophic refilling was also drawn by Pellen et al. (2017) on the basis of the onshore Adriatic record (see subsection 3.7).

### 3.7. Apennine system

The Messinian deposits resurface to the south of the PPAF sector and extensive sections are found in several basins on both the foreland domain (Adriatic side of the partially uplifted Apennine chain), subjected to compressional tectonics during the late Messinian, and the back-arc domain (Tyrrhenian side), contemporaneously affected by extension (Fig. 2b; Cipollari et al., 1999; Schildgen et al., 2014; Cosentino et al., 2018). Overall, the MSC record of the Apennines is subdivided into an evaporitic and post-evaporitic interval squeezed in between two marine units (Messinian Euxinic Shales Fm. at the base and Zanclean Argille Azzurre Fm. atop; Fig. 4a). Different vertical motions related to ongoing Apenninic tectonics resulted in the deposition of Stage 3 sediments with highly variable sedimentary expression and stratigraphic resolution from basin to basin. The post-evaporitic deposits are alternatively found resting unconformably, with an erosional contact associated to an angular unconformity, above the alternations of the Gessoso Solifera Fm./PLG, or conformably above evaporitic-free cycles lateral equivalent of the marginal PLG (Fig. 4a; e.g. Roveri et al., 1998, 2008a). This led to the conclusion that both shallow and deep-water successions are present in the Apennine foredeep system (Roveri et al., 2001).

The physical-stratigraphic model developed for the post-evaporitic interval in the Romagna area (i.e. Northern Apennines) and applied to the whole Apennine domain was subdivided into two allounits (named p-ev₁ and p-ev₂) based on a basin-wide shift in facies, overall stacking patterns and depositional trends (i.e. progradational and retrogradational, respectively; Roveri et al., 1998, 2001, 2005, 2008a; Manzi et al., 2005, 2007, 2020). Allounit p-ev₁ only accumulated in deep-water settings (e.g. Cusercoli, Sapigno, Maccarone and Trave sections; Roveri et al., 1998; Iaccarino et al., 2008; Cosentino et al., 2013) during the subaerial exposure of the basin margins (e.g. Vena del Gesso Basin, Monticino quarry, Pellen et al., 2017). It starts with resedimented clastic evaporites (i.e. Sapigno Fm.) followed by a coarsening- and shallowing-upward succession (i.e. di Tetto or San Donato Fm.) of mudstones with intercalated turbiditic sandstones (Fig. 5f) and a volcaniclastic marker bed dated initially by $^{40}$Ar-$^{39}$Ar at ~5.5 Ma (Odin et al., 1997) and later by $^{238}$U-$^{206}$Pb at 5.5320±0.0046/0.0074 Ma (Cosentino et al., 2013; Fig. 3a). Allounit p-ev₂ (i.e. Colombacci Fm.) occurs in the deeper depocenters in 4/5 sedimentary cycles consisting of three > 5 m-thick coarse-grained bodies (conglomerates and sandstones) and two black-

colored mudstone beds alternating with fine-grained mudstones/clays with intercalated three micritic limestones (known in literature as Colombacci limestones; Figs. 3a, 5g; Bassetti et al., 2004). By contrast, an incomplete Colombacci Fm. deposited in the shallower thrust-top basins (e.g. Vena del Gesso Basin and Molise sections; Pellen et al., 2017; Cosentino et al., 2018). The p-ev₂ cycles have been interpreted as reflecting the alternation of wet (mudstones and Colombacci limestones in Eastern Romagna) and dry (coarse-grained facies and Colombacci limestones in the Maccarone section) phases controlled by Milankovitch-driven climatic factors (Fig. 3a; Roveri et al., 2008a; Cosentino et al., 2013) and, as such, they have been used for the astronomical tuning of the Colombacci Fm. to the Lago-Mare phase (Figs. 3a, 4a; see subsection 2.2). By contrast, Clauzon et al. (2005) and Popescu et al. (2007) moved the Colombacci Fm. into the Pliocene (Fig. 4b). However, this conclusion has been proven to rely on wrong stratigraphic and paleontological arguments (see Roveri et al., 2008c, Grothe et al., 2018 and subsection 5.5). Substage 3.2 records in the Apennines do not always contain the three prominent conglomeratic facies as in Romagna, but only laminated to massive clays with sandy intercalations equivalent to the ones typifying substage 3.1 records (e.g. Maccarone section; Sampalmieri et al., 2010; Cosentino et al., 2013; Fig. 5f). The absence of a lithological cyclicity that clearly mimics an orbital signal largely hampered the astronomical tuning of these clay-dominated sections, although an attempt has been made with the Maccarone section (Cosentino et al., 2013). The only exception is represented by the Fonte dei Pulcini section, which has been equipped with an age framework by astronomical tuning of the magnetic susceptibility record (Cosentino et al., 2012). Despite the lack of outstanding lithological changes these sections are often provided with a lithostratigraphic subdivision using the same nomenclature as in the Romagna area. When applied, the di Tetto Fm.-Colombacci Fm. boundary is placed high in the sections, i.e. few tens of meters underneath the Miocene/Pliocene boundary, resulting in a much different thickness of the formations compared to the Romagna area.

Stage 3 sediments are poorly exposed on the Tyrrhenian Sea onshore side of Italy (Fig. 2b). The best known succession crops out in the Cava Serredi quarry in the Fine Basin (Tuscany; Bossio et al., 1978, 1993; Carnevale et al., 2006b, 2008). Here the MSC has a thickness of ~150 m, of which only the uppermost ~100 m are attributed, without clear arguments, to Stage 3 by Carnevale et al. (2006b). The lowermost ~40 m of the Stage 3 unit consists of mudstone with alternating sandstone bodies which have been attributed to Roveri et al. (1998)'s p-ev₁ allounit, while the uppermost ~60 m form the p-ev₂ allounit and include two prominent conglomerate bodies alternating with mudstones interbedded with sandstone horizons and black, organic-rich layers (Carnevale et al., 2006b). A few and more fragmented sections are also described on the Tyrrhenian Sea side of Italy by Cipollari et al. (1999).

The Miocene/Pliocene boundary is variably expressed through the Apennine system: unconformable above the ostracod-bearing clays and highlighted by erosional surfaces draped by conglomeratic accumulations (e.g. Stingeti section in Molise; Cosentino et al., 2018), conformable above 0.5-1 m-thick black mudstones similar to how it is observed in Piedmont and of equally unknown paleoenvironmental significance (e.g. Romagna area and Maccarone section; Roveri et al., 1998; Gennari et al., 2008) or conformable above the ostracod-rich mudstones (e.g. Maccarone and Fonte dei Pulcini sections; Cosentino et al., 2005, 2012, 2013; Sampalmieri et al., 2010).

All p-ev₁ deposits studied are almost devoid of in-situ biota, except for fish otoliths and three fish skeletons found in the upper substage 3.1 part of Cava Serredi (Carnevale et al., 2006b). The p-ev₂/Colombacci deposits, instead, host typical Paratethyan assemblages of brackish-water mollusks, ostracods, dinocysts and fish (Bassetti et al., 2003; Bertini, 2006; Popescu et al., 2007; Grossi et al., 2008; Iaccarino et al., 2008; Cosentino et al., 2012, 2018; Schwarzhans et al., 2020). A diverse array of marine fossils (benthic and planktic foraminifera, calcareous nannofossils, dinocysts and fish otoliths and skeletons) has also been
reported from the horizons containing these Paratethyan taxa (Bertini,
2006; Carnevale et al., 2006a; Popescu et al., 2007; Pellen et al., 2017).
While the autochthony of ostracods, when considered, is unquestioned,
the allochthonous vs autochthonous character of the other mentioned
fossils is disputed and still unclear (see Chapter 5).

3.8. Sicily

The MSC record is widely exposed on Sicily, mainly in the Caltanissetta Basin and in scattered locations on the Hyblean Plateau (i.e. Ragusa-Siracusa area) and the Messina area (Fig. 2b; Butler et al., 1995; Manzi et al., 2009; Sciuto et al., 2018). Like the Northern Apennines, it shows a complex distribution and variable stratigraphy that mirrors the structuring of Sicily into basins with different characters, geometries and depocenters which subsided at different times and rates (Butler et al., 1995; Catalano et al., 2013). This structural setting permitted the simultaneous deposition of shallow and intermediate-water sediments (Roveri et al., 2008b). Mostly found in the Caltanissetta Basin, these intermediate-water successions have for decades been considered the onshore counterpart of the offshore evaporitic trilogy seen in seismic data from the Western Mediterranean Basin (Decima and Wezel, 1973). More recently, Raad et al. (2021) attempted a similar onshore-offshore correlation but with the intermediate Central Mallorca Depression. The currently endorsed stratigraphic model (Fig. 4a), refined over the years by Decima and Wezel (1971, 1973), Decima et al. (1988), Butler et al. (1995), García-Veigas et al. (1995), Rouchy and Caruso (2006), Roveri et al. (2008b) and Manzi et al. (2009), envisages two ‘evaporitic cycles’. The ‘First cycle’, overlying both alluvial and deep-water sediments (Tripoli Fm., Licata Fm. and Terravecchia Fm.; see Maniscalco et al., 2019 and references therein), comprises the disputed Calcare di Base (Manzi et al., 2011, 2016b vs Caruso et al., 2015), PLG or Gessi di Cattolica Fm. (Decima and Wezel, 1973; Lugli et al., 2010) and the Halite Unit (Lugli et al., 1999). The ‘Second cycle’ comprises the Upper Gypsum (UG) or Gessi di Pasquasia Fm., which is only present in depocenters of the Caltanissetta Basin (see Manzi et al., 2009 for a detailed overview), sporadically overlain by the siliciclastic Arenazzolo Fm. (Decima and Wezel, 1973; Cita and Colombo, 1979). The whole succession is sealed by the Pliocene marine Trubi Fm. (Fig. 4a). The two evaporite cycles are separated by an erosional surface (MES) associated with an angular discordance broadly linked to the main Mediterranean drawdown event (e.g. Butler et al., 1995; Roveri et al., 2008b). Clauzon et al. (1996), however, placed the MES at the Arenazzolo Fm.-Trubi Fm. transition, implying that the entire evaporitic deposition in the Caltanissetta Basin pre-dated the offshore one, but they do not provide evidence of erosion at that level. In more recent publications from the same research group, the MES is shifted towards the base of the Arenazzolo Fm. (e.g. Bache et al., 2012), again without evidence of major erosion, and different ages are assigned (see Fig. 4b and Grothe et al., 2018 for details).

The Upper Gypsum successions are commonly incomplete in many of
the Caltanissetta Basin sections (Pasquasia-Capodarso, Casteltermini,
Alimena, Nicosia, Siculiana-Marina; Decima and Sprovieri, 1973; Rou-
chy and Caruso, 2006; Manzi et al., 2009; Fig. 5i). In the most complete
section, Eraclea Minoa (Fig. 3a), the Upper Gypsum Unit consists of 6
(Van der Laan et al., 2006) to 7 (Manzi et al., 2009) primary gypsum
beds with a repetitive internal organization of facies (see Schreiber,
1997 and Manzi et al., 2009 for facies description) interbedded with
marls and lenticular terrigenous sandstone bodies, gypsarenites and
gyprudites (Fig. 5h). Two of the terrigenous sandstone bodies are
highlighted by Manzi et al. (2009) in the thick (~60 m), Cyprideis
agrigentina-rich (Grossi et al., 2015), marly interval dividing gypsum VI
and VII for its alleged astronomical significance (Fig. 3a; see subsection
2.2). A mixed (physically reworked) marine (foraminifera and dino-
cysts) and (in-situ) brackish biota (ostracods and dinocysts) of Para-
tethyan origin characterizes the marly interbeds from at least gypsum III
upwards (following the investigations carried on the Eraclea Minoa

section; Bonaduce and Sgarrella, 1999; Rouchy and Caruso, 2006;
Londeix et al., 2007; Grossi et al., 2015; Fig. 3a). Calcareous nannofossils
have been found along with the above organisms in a more northerly
location by Maniscalco et al. (2019) and considered reworked. Above
the last gypsum, the ~6-7 m-thick Arenazzolo Fm. is found, represented
by reddish arkosic cross-laminated and poorly consolidated sand
(Bonaduce and Sgarrella, 1999; Roveri et al., 2008b) and interpreted as
the expression of a shallow-water delta, albeit without a sedimento-
logical investigation (e.g. Decima and Wezel, 1973; Cita and Colombo,
1979). The whole Stage 3 sequence is conformably overlain by the
Zanclean marine Trubi Fm. in the basin center (e.g. at Eraclea Minoa and
Capo Rossello; Fig. 6d; Brolsma, 1975; Cita and Colombo, 1979; Van
Couvering et al., 2000; Rouchy and Caruso, 2006; Manzi et al., 2009;
Fig. 6d) and unconformably in the shallower, marginal areas (Manzi
et al., 2009; Roveri et al., 2019b). Only Decima and Wezel (1973) and
Raad et al. (2021) report the Miocene/Pliocene transition in the key,
intermediate water-representative section of Eraclea Minoa as erosive.
However, they do not provide evidence (e.g. photographic documenta-
tion) for the presence of an erosional unconformity and, moreover, Raad
et al. (2021) erroneously refer to Cita and Colombo (1979), where no
erosion is mentioned at the M/P boundary in Eraclea Minoa.

The bathymetric jump between the <100 m of water depth during the late Messinian and the >200 m at the base of the Trubi Fm. is often regarded as a key onshore evidence of the sudden and catastrophic Mediterranean-Atlantic re-connection at the Miocene/Pliocene boundary (e.g. Caruso et al., 2020). However, the real depth of the base of the Trubi is all but obvious. In fact, variable estimates have been proposed based on the observed benthic foraminifera and/or psychrospheric ostracods at Capo Rossello and Eraclea Minoa: 200-500 m (Decima and Wezel, 1973), 600-800 m (Sgarrella et al., 1997, 1999; Barra et al., 1998), 1400-2400 m (Cita and Colombo, 1979).

3.9. Greece

Several MSC localities are reported from the Greek Ionian Islands
(Corfu, Cephalonia and Zakynthos) and from Crete (Fig. 2b).

On the NW coast of Corfu (Aghios Stefanos section), the PLG unit is missing and only a 32 m-thick cyclically-arranged terrigenous succession is present comprising three m-thick conglomerate beds alternating with fine-grained deposits rich in unspecified species of brackish water ostracods (Pierre et al., 2006).

In the southern part of Zakynthos, an evaporitic succession composed of eight gypsum cycles (Kalamaki section) occurs above marine marly deposits (Karakitsios et al., 2017b). These gypsum beds were initially ascribed to the UG unit (Pierre et al., 2006) and later to the PLG (Karakitsios et al., 2017b). The gypsum unit is overlain by approximately ~13 m of siltstones and marls with scattered, cm-thick beds of sandstones, conglomerates and carbonates with nodular texture (Pierre et al., 2006; Karakitsios et al., 2017b). Although no ostracods are reported from this interval, due to its stratigraphic position the post-evaporitic unit is correlated to the Lago-Mare phase (Karakitsios et al., 2017b). Except for the rare presence of marine nannofossils (Ceratolithus acutus together with Reticulofenestra zancleana) just below the Miocene/Pliocene boundary, only reworked marine fauna has been reported from the post-evaporitic package (Karakitsios et al., 2017b). This dominantly terrigenous succession is unconformably overlain by the Zanclean Trubi Formation (Fig. 6e; Karakitsios et al., 2017b).

MSC deposits on Crete (e.g. Meulenkamp et al., 1979; Delrieu et al., 1993; Cosentino et al., 2007; Roveri et al., 2008a; Zachariasse et al., 2008, 2011) were deposited in Miocene extensional, fault-bound basins driven by tectonic subsidence that ceased in the late Pliocene and Pleistocene (Van Hinsbergen and Meulenkamp, 2006). Because of the strong tectonic and eustatic sea-level-related fragmentation of the stratigraphic record, reconstructing the late Miocene stratigraphy of Crete has not been straightforward (Zachariasse et al., 2008, 2011). Several primary and clastic gypsum facies are recognized, but their
correlation with the MSC stratigraphy is disputed (see Cosentino et al., 2007; Roveri et al., 2008a, 2014a; Zachariasse et al., 2008). Coarse-grained, mammal-bearing terrigenous facies irregularly alternating with marls are in places found unconformably overlying the gypsum and separated from the Pliocene facies by an erosion surface (see Meulenkamp et al., 1979; Delrieu et al., 1993; Cosentino et al., 2007). In two localities on the Messarà Plain, Cosentino et al. (2007) described a highly diversified ostracod fauna with Paratethyan affinity in some marly intervals.

Messinian evaporites and/or Lago-Mare deposits are also reported from the North Aegean region onshore in the Strymon Basin (Snel et al., 2006; Suc et al., 2015; Karakitsios et al., 2017a) and Dardanelles region (Melinte-Dobrinescu et al., 2009) and offshore (Prinos-Nestos Basin; Karakitsios et al., 2017a), but recent integrated studies suggested that the sections studied by the above listed authors are older than the MSC (see Krijgsman et al., 2020a, 2020b). In particular, Krijgsman et al. (2020a) proposed that for most, if not all, of the MSC the North Aegean was a brackish water, mostly Paratethyan-fed basin restricted by the Cyclades sill to the south (Fig. 2a) and forming a passageway for Paratethyan overspill waters towards the Mediterranean.

### 3.10. Cyprus

MSC deposits on Cyprus outcrop in the Pissouri, Psematismenos, Mesaoria and Polemi basins on the southerly fringe of the Troodos massif (Fig. 2b; Rouchy et al., 2001; Manzi et al., 2016a). According to Rouchy et al. (2001) and Orszag-Sperber et al. (2009), sediments belonging to all MSC stages of Roveri et al. (2014a) are preserved in the Cypriot basins. By contrast, Robertson et al. (1995) and Manzi et al. (2016a) considered that PLG evaporites on Cyprus are only present as fragments reworked within a chaotic unit (the Lower Gypsum and Intermediate breccia units of Orszag-Sperber et al., 2009) and that the only *in situ* evaporites belong to the overlying Upper Gypsum Unit, which encompasses the whole of Stage 3 (Figs. 3a, 4a). A continuous, Eraclea Minoa-like section is not known in Cyprus (Manzi et al., 2016a). The best exposure of the lower 60 m of this unit is found in the Polemi Basin (Manzi et al., 2016a). It comprises up to six gypsum beds (the lower three of which are mainly selenitic, while the upper three are predominantly laminated; Fig. 3a). Gypsum beds range in thickness from 1 to 6 m and are separated by laminated marls (Fig. 5j) occasionally interbedded with conglomerates and sandstones (e.g. between the 5th and 6th gypsum layers; Rouchy, 1982; Rouchy et al., 2001; Manzi et al., 2016a). The sixth gypsum bed is reported by Rouchy et al. (2001) to be hollowed in the upper part with the cavities filled with overlying sediments. The similarity of the cyclicity and facies association of this Cyprus succession with the substage 3.1 interval of the Sicilian UG led Manzi et al. (2016a) to propose a bed-to-bed correlation and to recognize the substage 3.1/3.2 boundary at the top of the last gypsum bed (Fig. 3a). According to Orszag-Sperber et al. (2000) and Rouchy et al. (2001), this chronostratigraphic boundary coincides with a Mediterranean-scale sea-level drop, a conclusion that arises from the interpretation of the cavities in the uppermost gypsum as the product of karstic dissolution following a prolonged period of subaerial exposure.

The sedimentary sequence overlying the last gypsum bed and assigned by Manzi et al. (2016a) to the Lago-Mare phase lacks a clear and rhythmic sedimentary cyclicity. In the Pissouri Basin this interval (up to 25-30 m-thick) mostly consists of conglomerates, sandstones, limestones, paleosols (which appear as dm to m-thick dark marly horizons, in one case with pulmonated gastropods) and subordinated clay-marly horizons (Rouchy et al., 2001). By contrast, in the Polemi sections the clay-marly facies dominates this interval (Rouchy et al., 2001). *In situ* fresh-brackish water species of articulated mollusks (*Limno cardiidae*, *Melanopsis*), Paratethyan (*Loxocorniculina djafarovi*, *Euxynocythere praebaquana*) and Mediterranean (*Cyprideis agrigentina*) ostracods and foraminifera (*Ammonia beccarii*), *Characeae*, abundant fragments of the marine euryhaline fish *Clupeidae* and a fish skeleton of

the euryhaline *Aphanius crassicaudus* are described from some of the substage 3.1 and 3.2 fine-grained facies and within the terrigenous laminae of some *balatino* gypsum (Orszag-Sperber et al., 2000; Rouchy et al., 2001; Orszag-Sperber, 2006; Manzi et al., 2016a). The upward change in diversity of the ostracod fauna seen elsewhere (e.g. Malaga, Nijar, Vera, Apennines and Eraclea Minoa) is not reported in Cyprus but this may be because no detailed study of ostracod assemblages in Stage 3 sediments has been published. The Miocene/Pliocene boundary, near Polemi village is described by Manzi et al. (2016a) as a sharp contact above a dark, organic-rich layer (Fig. 6f). It appears to be similar to the boundary reported from Piedmont (Fig. 6c; Trenkwalder et al., 2008; Dela Pierre et al., 2016) and Northern Apennines (Gennari et al., 2008; Grossi et al., 2008), if not for the presence, in Cyprus, of (possibly) *in-situ* *Cyprideis agrigentina* (Manzi et al., 2016a). A layer with the same field appearance, thickness (~ 1 m) and stratigraphic position is reported in Pissouri by Rouchy et al. (2001), which they interpreted as a paleosol based on mottling, oxidized roots, carbonate concretions and plant fragments.

### 3.11. Southern Turkey

The tectonically active, during the Miocene, thrust-top basin of Adana in southern Turkey (Radeff et al., 2017) retains the most complete and better exposed easternmost successions of the MSC (Fig. 2b), whose deposits were attributed to the Handere Fm. (Cosentino et al., 2010; Radeff et al., 2016).

MSC Stage 3 finds expression in a >1 km thick continental unit unconformable, through an erosional surface, above the pre-evaporitic, Stage 1 anhydrite-shale alternations (Radeff et al., 2016) and resedimented gypsum-bearing Stage 2 deposits (Cosentino et al., 2010; Cipollari et al., 2013). This unit mainly consists of fluvial coarse- and fine-grained deposits representing channel fill and overbank deposits. Sporadically, some fine-grained intercalations are found containing a mixed brackish (ostracod) and marine (foraminifera and calcareous nannofossils) fauna. The ostracod fauna has affinity with the Paratethyan fauna but, unlike to many other Mediterranean onshore localities, is poorly diversified, with monospecific assemblages of *Cyprideis agrigentina* (Avadan section and T-191 borehole; Cipollari et al., 2013) or with *Cyprideis agrigentina* accompanied by rare to abundant specimens of *Loxoconcha muelleri*, *Euxinocyhere* (*Maeotocythere*) *praebaquana*, and *Loxoconcha* sp. (Adana section; Faranda et al., 2013). Ostracods are often associated with *Ammonia beccarii* and rare *Elphidium* and *Cribroelphidium*, which are the only foraminifera considered as autochthonous. Conversely, the entire nannoflora is interpreted as physically reworked (Cipollari et al., 2013; Faranda et al., 2013).

The Handere Fm. is followed by early Zanclean marine sediments (Avadan Fm.) deposited, according to the paleoecology of the benthic foraminifera species recognized, at bathymetries ranging from 200 to 500 m (Cipollari et al., 2013). The lithological nature of the Miocene/Pliocene boundary in the Adana Basin is not clear, but it occurs either above the continental or subaqueous, ostracod-bearing facies.

A similar stratigraphic sequence is present in the subsurface. Here, however, chaotic gypsum-bearing deposits are not found and two halite bodies ~20 and ~170 m-thick are present, separated and followed by fluvial gravels, sands and silts (Cipollari et al., 2013).

### 3.12. Summary of the onshore Stage 3 record

Most of the onshore Stage 3 records formed in shallow marginal Mediterranean basins, which underwent substantial uplift from the Messinian till nowadays and are assumed to have had their depocenter at ~200 to 50 m below the Atlantic level during the late Messinian (Roveri et al., 2014a, 2019a; Radeff et al., 2016, 2017). The Calta-nissetta Basin (Sicily), some basins along the Apennines and (possibly) Cyprus represent, instead, possible onshore representative of intermediate basins. The nature and duration of these records is quite variable,
and there are only six sections that may record an entire Stage 3 sequence (i.e. Sorbas, Nijar, Northern Apennines, Eraclea Minoa and Cyprus; Fig. 3a). Reasons for the fragmentary nature of the Stage 3 sedimentary record include different durations of subaerial exposure following the Stage 2 drawdown, local tectonics and associated depositional erosion and deposition. One of the consequences of this is that any sedimentary cyclicity that resulted from orbital fluctuations is typically either less well developed or poorly preserved, making the chronology of Stage 3 rather uncertain or controversial in places.

Despite this variability, several fairly consistent characteristics are widely expressed. These are:

1) Stage 3 sedimentation follows a period of intensive tectonic and/or eustatic-driven erosion of the margins, as demonstrated by the frequent presence of erosional unconformities and/or chaotic Stage 2 deposits (RLG unit);

2) Stage 3 lithologies are mostly terrigenous (conglomerates, sandstones and mudstones) and deposited in a variety of continental (fan delta, fluvial channels and alluvial plains) and shallow water environments (endorheic lakes or water bodies connected with the Mediterranean water mass is the riddle). Carbonate intercalations are sometimes present (e.g. Sorbas Basin and *Colombacci* limestones in the Apennines). Stage 3 gypsum is only found in deeper-water intermediate basins of Caltanissetta in Sicily and Cyprus.

3) A diversified fossil assemblage with Paratethyan affiliation (ostracods, dinocysts, mollusks) is commonly found in the shallow-water sediments of Lago-Mare successions. Only in the intermediate Caltanissetta Basin (Sicily) do these diversified Paratethyan forms (only ostracods) occur earlier, in the sediments from substage 3.1. Where these have been studied in detail, these assemblages typically show an increase in diversity with time (possibly every wet phase of the precession cycles). Some of these sediments also contain marine fossils and there is controversy over whether these are *in situ* and contemporaneous or reworked.

4) In outcrop, the Miocene/Pliocene boundary has four main sedimentary expressions: erosive and followed by a conglomeratic lag (e. g. Malaga, Vera, Mallorca; Fig. 6a); conformable above continental facies (e.g. Nijar Basin; Fig. 6b); conformable above ostracod-rich mudstones (e.g. Eraclea Minoa; Fig. 6d); sharp contact above a black layer of still largely unknown paleoenvironmental significance (Piedmont, Apennines and Cyprus; Fig. 6c, f).

For a better understanding of how Stage 3 developed across the Mediterranean these marginal records now need to be considered alongside the evidence from intermediate to deep offshore settings.

We note that alternative chronostratigraphic frameworks have been proposed for several onshore (Malaga, Sorbas, Mallorca, Apennines, Sicily) and offshore (Sites 134B, 976B, 978A) locations (see Fig. 4b for references), but we have omitted them as they are shown to rely on incorrect (bio)stratigraphic arguments (see Roveri et al., 2008c, Grothe et al., 2018 and subsection 5.5).

## 4. Offshore domain

The offshore Mediterranean is a complex array of variable-depth and morphologically complex subbasins framed by morphological highs or sills. Traditionally it is divided into two main domains (Fig. 2a), the Western and Eastern Mediterranean, with the intervening divide (or Sicily sill) situated in the Sicily channel at present with a maximum depth of 316 m. The Alborán Basin, the depressions on the Balearic Promontory, the Corsica, Valencia, Algero-Balearic, Liguro (or Sardo)-Provençal and Tyrrhenian basins belong to the "Western Mediterranean" (Fig. 2a). The Adriatic foredeep, the Ionian, Sirte, Aegean and Levant basins belong in the "Eastern Mediterranean" (Fig. 2a). Smaller-sized depressions, again surrounded by sills of variable depth, are identified and labelled within each of these subbasins.

Although the exact topography and hypsometry of the Messinian Mediterranean is difficult to reconstruct, this present-day geography is generally assumed to have been in place, with five main differences: (1) the Alborán Basin was split into a Western (WAB) and Eastern Alborán (EAB) by a volcanic arc (Booth-Rea et al., 2018); (2) the Tyrrhenian Basin was only partly opened (Lymer et al., 2018); (3) the precise depth and width of the ancient Sicily Sill are difficult to estimate, but may have been much deeper than today (~300 m; Meijer and Krijgsman, 2005; Jolivet et al., 2006). Paleodepth estimations for the Messinian configuration range from 380 m (Just et al., 2011) to 430 m (García-Castellanos et al., 2009); (4) one or two sills were present at the southern termination of the Adriatic foredeep (Pellen et al., 2017; Amadori et al., 2018; Manzi et al., 2020); (5) the North Aegean was partially isolated from the Mediterranean by the Cyclades Sill and with high Paratethys affinity (Krijgsman et al., 2020a). Following the schematic classification of the Messinian sub-basins by Roveri et al. (2014a), all these subbasins are regarded as either intermediate (i.e. relatively deep-water, 200–1000 m) or deep (water depth > 1000 m).

Compared with the onshore domain, the offshore basins hold a far greater percentage of the total volume of MSC sediments (Ryan, 1973; Haq et al., 2020). The architecture, geometry and main lithologies of the MSC and younger deposits are well known thanks to the high density of seismic data and the fact that evaporites (halite particularly) are easily identified on seismic profiles due to their unusual seismic properties, especially compared to those of terrigenous and carbonate sediments (seeofi et al., 2011a, 2011b; Lofi, 2018; Haq et al., 2020). However, the detailed lithological, sedimentological, paleontological and geochemical nature and their chronostratigraphy are still poorly constrained offshore because these cannot be univocally defined on the basis of seismic data alone (Roveri et al., 2019b) and direct information about these deep MSC successions is limited to a small number of cores (16) drilled during the DSDP (Ryan et al., 1973; Hsü et al., 1978b) and ODP (Kastens et al., 1987; Comas et al., 1996; Emeis et al., 1996) drilling campaigns that penetrated exclusively the uppermost tens of meters of the deep MSC suite in very scattered localities (Fig. 2b). Only recently, access to industrial boreholes crossing the base of the halite in the deep Levant Basin has been granted, providing a rare glimpse of the deep MSC deposits in the easternmost part of the Mediterranean (Gvirtzman et al., 2017; Manzi et al., 2018; Meilijson et al., 2018, 2019).

The MSC is commonly described as bipartite ('Messinian trilogy' after Montadert et al., 1978) in the Western Mediterranean (Lower-Mobile-Upper units: LU-MU-UU, respectively). However, in the Ionian Basin is described as bipartite (MU-UU) by Camerlenghi et al. (2019) while according to Lofi et al. (2011a), Gvirtzman et al. (2013, 2017), Lofi (2018) and Camerlenghi et al. (2019), the Levant Basin consists of the MU and the UU is only present locally and possibly represented by evaporite-free terrigenous accumulations (Kartveit et al., 2019; Madof et al., 2019). The lack of many age constraints within the offshore MSC successions hampers unambiguous correlation with onshore sequences (Fig. 1a; Roveri et al., 2014a). Nevertheless, different authors have proposed onshore-offshore correlation of specific events (e.g. onset, Ochoa et al., 2015; and termination of the MSC, Biscaye et al., 1972, Iaccarino et al., 1999) and stratigraphic schemes (Decima and Wezel, 1971; Raad et al., 2021) based on and biostratigraphic evidence (Cosentino et al., 2006), $^{87}$Sr/$^{86}$Sr isotope ratios (Roveri et al., 2014b; Gvirtzman et al., 2017; Manzi et al., 2018) and astronomical tuning of the deep seismic record (Ochoa et al., 2015, 2018; Manzi et al., 2018; Meilijson et al., 2018, 2019). Here we focus on the seismic and
<figure><img src="image_11.png" /><figcaption>**Fig. 7.** Seismic profiles from intermediate-deep Western Mediterranean basins containing MSC markers/units. (a) Seismic reflection line CAB01-104 from the WAB (modified from Booth-Rea et al., 2018). The line shows the variable geometry of the inferred M/P boundary, erosive in proximity of mud diapirs, (para)conformable in tectonically undisturbed sectors. Chaotic reflectors are occasionally imaged below the inferred M/P boundary. (b) Seismic profile SF12-09 imaging the lower slope of the south Algero-Balearic margin and part of the Algero-Balearic abyssal plain (modified from Mocnik et al., 2014). Here a high reflecting and horizontally stratified UU covers a thin layer of MU evidenced by salt diapirism. Note the concordant deformation of the UU and MU. (c) Line MS-39 from the abyssal plain of the Liguro-Provençal Basin showing the Messinian trilogy and non-erosive bottom and top surfaces (BS and TS; Dal Cin et al., 2016). Halokinesis of MU gives rise to domes that deform the UU and PQ units. (d) Interpreted seismic profile from the lower-middle slope of the west Sardinian margin (modified from Dal Cin et al., 2016). Thin MU and UU are present on the lower slope, while on the middle slope (and upper slope here not shown) they converge in the margin erosion surface MES. (e) Line drawing of seismic line imaging from the Catalan margin (or Ebro Basin) to the abyss of the Liguro-Provençal basin (modified from Maillard et al., 2011b). Note the pinch out of the MU in the Valencia Basin and of the UU in the Ebro Basin, which is MSC free. (f) Interpreted seismic profile Simbad 15 crossing the depocenter of the CMD showing all the MSC units and erosional surfaces (modified from Raad et al., 2021). (g) Interpreted seismic profile MYS40 illustrating the MU-UU-PQ units in the East-Sardinia Basin and Cornaglia Terrace, separated by the MSC deposits-free Quirra Seamounts (modified from Lymer et al., 2018).</figcaption></figure>
geological (core-derived)<fnref n="1" /> properties of the Upper Unit (and laterally grading/interfingering sediments when present), stratigraphically below the Plio-Quaternary deposits (PQ) suggesting that it belongs to (at least part of) Stage 3.

## 4.1. Western Alborán Basin and westernmost East Alborán Basin

The Alborán Basin has received particular attention because of its proximity to the Gibraltar Corridor (Estrada et al., 2011; Popescu et al., 2015 and references therein). Evaporites only occur on the eastern side of the EAB (which is treated in subsection 4.2; Fig. 2a). To the west of the volcanic archipelago (Booth-Rea et al., 2018, i.e. the WAB) and immediately to the east on the western side of the EAB only terrigenous sediments occur in the MSC interval (Booth-Rea et al., 2018; de la Peña et al., 2020). Sediments at the Miocene/Pliocene boundary appear in the seismic reflection data as parallel reflectors with increasing reflectivity (Comas et al., 1996, 1999; Booth-Rea et al., 2018). Locally, just below the M discontinuity, some of the reflectors suggest a chaotic seismic facies (Fig. 7a; Booth-Rea et al., 2018; Bulian et al., 2021). Miocene sediments with a maximum thickness of 100 m have been recovered from two out of nine holes drilled in the region (ODP 976B, 978A; Comas et al., 1996, 1999). These sediments mostly consist of claystones, siltstones and sandstones with *Chondrites* and *Zoophycos* ichnofacies at site 976B and include a conglomerate close to the Miocene/Pliocene boundary at Site 978A. The lack of age-diagnostic fossils hampers their correlation with the Geologic Time Scale (GTS). However, the presence, high in the Miocene section, of an oligotypic association of ostracods (*Candona* sp., *Loxoconcha* muelleri, and *Cyprideis* sp.) with different stages of growth (Site 978A; Iaccarino and Bossio, 1999) and Paratethyan dinocysts (including *Galeacysta etrusca*; see subsection 5.2; Popescu et al., 2015) indicates a latest Messinian (substage 3.2) age and brackish paleodepositional conditions. Foraminifera and nannofossils are also present, but all species recognized are of no help in narrowing down the paleoenvironmental interpretation because they are considered either definitely or likely to be reworked (Iaccarino and Bossio, 1999). By contrast, Popescu et al. (2015) interpreted some species of calcareous nannofossils (*Reticulofenestra pseudoumbilicus*, *Discoaster quinqueramus*, *Ceratolithus acutus*, *Triquetrorhabdulus rugosus*, *Amaurolithus primus*) and marine dinocysts as autochthonous.

The nature of the Miocene/Pliocene boundary is also uncertain. According to some authors, the “M” discontinuity is a high-amplitude reflector with evidence of erosion attributed to subaerial processes (Estrada et al., 2011; Urgeles et al., 2011) and locally (e.g. close to Site 121; Ryan et al., 1973) associated with an angular unconformity that abruptly truncates the upper Miocene deposits and morphological highs (Comas et al., 1999; Estrada et al., 2011; Garcia-Castellanos et al., 2020). Although the M-reflector was drilled at Sites 976B, 977A and 978A, a lithological contact was only recovered at Site 976B coinciding with a major erosional surface between the early Messinian and the base of the Pliocene (Bulian et al., 2021). Only at Site 978A (and possibly 977A) was a few meters of what may be the contact interval recovered (Comas et al., 1996). This comprises a 25 m-thick cemented succession containing pebbles of volcanic and sedimentary rocks likely to derive from the Alborán substrate (46R, 620.9-630.67 mbsf, between the Pliocene-bearing core 45R and the Messinian-bearing core 47R; Comas et al., 1996). In contrast, Booth-Rea et al. (2018) concluded that the M-reflector is an unconformity only close to the mud diapirs and owes its erosive shape and angular discordance to the activity of these structures

(Fig. 7a). In more undisturbed sectors these authors argue that the boundary is a paraconformity with no evidence of erosion (Fig. 7a). The lack of Messinian erosion in the shallow regions of the WAB margins has prompted the hypothesis that this area did not desiccate during the MSC (Booth-Rea et al., 2018; de la Peña et al., 2020). This contradicts much of the interpretation made of the DSDP and ODP cores of this interval in the Alborán Sea. The succession recovered by drilling from beneath the Pliocene comprises gravels that contain a mixed Miocene fauna. These sediments and their faunal content are thought to have been reworked from older sediments exposed as Alborán substrate with no evidence of an extensive wet Lago Mare interval immediately before the Zanclean (Comas et al., 1996).

Two W-E-aligned erosional channels straddling the Strait of Gibraltar and stretching for 390 km from the easternmost Gulf of Cádiz (Atlantic Ocean) into the Alborán Basin are clearly observed in seismic profiles (Garcia-Castellanos et al., 2009; Estrada et al., 2011). There is disagreement, however, concerning the timing and nature of their formation. These incisions are classically considered to occur at the very top of the MSC suite (when present) and to be the consequence of the Zanclean megaflood (Garcia-Castellanos et al., 2009, 2020; Estrada et al., 2011 among others). More recently, Krijgsman et al. (2018) highlighted that an accurate age determination for these channel incisions is lacking and that they might have been formed earlier during the MSC as a result of the Mediterranean-Atlantic gateway currents. Interpretation of E-W seismic profiles across the Alborán Basin combined with mammal records in Africa and Iberia led Booth-Rea et al. (2018) to suggest the existence of an emergent volcanic archipelago that temporarily connected southeastern Iberia with northern Africa, separating an open marine, Atlantic-influenced West Alborán Basin realm from a restricted, hydrologically complex Mediterranean realm to the east.

## 4.2. Eastern Alborán, Algero-Balearic and Liguro-Provençal basins

From the eastern margin of the EAB as far east as the Tyrrenian coast of Italy, the Messinian (evaporites-bearing) trilogy LU-MU-UU is found and sealed by the PQ. The MU and UU are interpreted to fill the deepest depocenters (Algero-Balearic, Valencia and Liguro-Provençal basins; with minor interruptions due to seamounts) and the lower slope domain, where they comprise ~500 to ~800 m of UU and ~1000 m of MU/halite (Figs. 7b-d; Camerlenghi et al., 2009; Lofi et al., 2011a, 2011b; Geletti et al., 2014; Mocnik et al., 2014; Dal Cin et al., 2016; Lofi, 2018). Upslope, a thinner, possibly incomplete UU is locally described on the middle-upper continental slopes of Western Corsica (Guennoc et al., 2011) and Sardinia (Mocnik et al., 2014; Dal Cin et al., 2016) and the northern (Maillard et al., 2006) and southern (Maillard and Maufret, 2013; Mocnik et al., 2014; Dal Cin et al., 2016) flanks of the Balearic Promontory, even though the structural settings of these locations are mostly dominated by erosion (Fig. 7d). MSC evaporites are absent on the continental shelves bordering the deep Algero-Balearic and Liguro-Provençal basins, where the PQ directly overlies the MES which, in turn, cuts through the middle Miocene deposits (Gorini et al., 2005; Lofi et al., 2005). The only late Messinian sediments are present as Complex Units (Gulf of Lion, Bessis, 1986; Gorini et al., 2005; Lofi et al., 2005; Algerian Basin, Medaouri et al., 2014; Arab et al., 2016; Fig. 7e). CUs can have various origin (Lofi et al., 2011a, 2011b), but when identified at the outlet of drainage systems, they are commonly interpreted as Messinian clastics supplied by rivers (Lofi et al., 2005). In the Gulf of Lion, the MES is a high angle unconformity with substantial erosion along highly rugged relief thought to have been generated by fluvial incision (Lofi et al., 2005). In contrast, Roveri et al. (2014c) suggested that the drainage networks visible on the seismic could be of subaqueous origin.

When not involved in MU-related deformation processes, the UU appears as a highly reflective series of flat reflectors alternating with less reflective, but concordant, reflectors (Figs. 7b-c; Lofi et al., 2011a, 2011b) aggrading in the basin center and onlapping the margins

<footnote n="1">¹ Lithostratigraphic and biostratigraphic information from DSDP and ODP cores are primarily extracted from the Scientific Shipboard Party documents, accessible from https://www.marum.de/en/Research/Cores-at-BCR.html. These documents are referenced in the text as follows: Ryan et al. (1973): DSDP 120-134; Hsü et al. (1978b): DSDP 371-378; Kastens et al. (1987): ODP 650-656; Comas et al. (1996): ODP 974-979; Emeis et al. (1996): ODP 963-973.</footnote>
(Fig. 7b; Camerlenghi et al., 2009; Lofi et al., 2011a, 2011b; Geletti et al., 2014; Mocnik et al., 2014; Dal Cin et al., 2016). The aggravating nature, shelf-ward thinning and the onlap terminations of the UU are interpreted as evidence of sedimentation in a lake with fluctuating base level (e.g. Lofi et al., 2005; Lofi et al., 2011a). In the abyssal plains (Figs. 2a, 7c), nine to ten cycles have been interpreted on high resolution seismic profiles as corresponding to gypsum-marl alternations (Geletti et al., 2014; Mocnik et al., 2014). At Sites 124 and 372, ~40-50 m of the UU have been drilled at the feet of the east Menorca continental rise and the northern Menorca slope, where 3-4 gypsum-marl cycles are recognized (Fig. 2b; Ryan et al., 1973; Hsü et al., 1978a). Primary gypsum facies are widely overprinted by post-depositional diagenetic processes, but the still recognizable laminated and clastic primary textures indicate precipitation at the water-air interface and emplacement by gravity flows, respectively (Lugli et al., 2015). The marl interbeds are made from stiff to firm dolomitic mud containing substantial quantities of detrital material intercalated with current-bedded sandstones and, at Site 124, with diatomites (Ryan et al., 1973). Cyprideis sp. specimens are reported from some mudstone interbeds at Site 372, while dwarf planktonic foraminifera are present just below the Miocene/Pliocene boundary at Site 124 (Ryan et al., 1973).

The Miocene/Pliocene boundary coincides with the top of the UU where present (labelled TES when erosional and TS when conformable; Lofi et al., 2011a, 2011b). In the abyssal plain-lower slope domain it appears to be undulating, although this geometry is related to the deformation of the underlying salt (Figs. 7b-c), and it actually corresponds to a sharp surface lacking signs of erosion (Lofi et al., 2011a, 2011b; Geletti et al., 2014; Mocnik et al., 2014). By contrast, the UU-PQ boundary commonly appears strongly erosional in the middle-upper slope and shelf domain, where it coincides with the MES (Fig. 7d; Lofi et al., 2005; Maillard et al., 2006; Geletti et al., 2014; Mocnik et al., 2014). Among the six DSDP-ODP Sites drilled in this region (Fig. 2b), only Hole 975B recovered the Miocene/Pliocene boundary (Iaccarino and Bossio, 1999; Marsaglia and Tribble, 1999). Here the Messinian is a few centimeters thick and consists of banded micritic silt clays with minor calcareous siltstones to sandstones typified by a diverse faunal assemblage consisting of dwarf planktonic foraminifera, Ammonia tepida tests and brackish Paratethyan ostracods (Loxocorniculina djafarovi, Euxinocythere praebaquana, Amnicythere idonea, Leptocythere limbate, Candona sp., and Cyprideis sp.; Iaccarino and Bossio, 1999).

Halite is present at the bottom of Hole 134 drilled within the UU (Ryan et al., 1973; Sage et al., 2005; Lugli et al., 2015). High-resolution seismic profiles from both the Algero-Balearic and Ligurian-Provençal basins confirm the presence of a halite layer high in the UU sequence (Geletti et al., 2014; Mocnik et al., 2014). This layer is up to 50 m thick (Dal Cin et al., 2016) and is correlated with an erosional surface (called IES: Intermediate Erosional Surface by Lofi et al., 2011a, 2011b) associated with an angular unconformity which is better developed on the lower slope (Fig. 7d). Geletti et al. (2014) and Mocnik et al. (2014) interpreted this layer as autochthonous and indicative of at least one important sea level drop during UU deposition. However, this intra UU halite layer is always described from areas strongly affected by salt diapirism (just like in the Ionian Abyssal Plain; see subsection 4.6.1) and is never found in adjacent, undisturbed areas (see Camerlenghi et al., 2009; Geletti et al., 2014; Mocnik et al., 2014; Dal Cin et al., 2016), two features that may suggest it has an allochthonous origin.

Site 134 shows evidence of a "desiccation crack" cutting through a sandy silt layer interbedded with unaffected laminated halite (Hsü et al., 1973c). Unfortunately, the core photograph of this crack has been published in two different orientations (Hsü et al., 1973a, 1973b), leading both Hardie and Lowenstein (2004) and Lugli et al. (2015) to question the evidence for subaerial desiccation. Because of these ambiguities, we suggest to dismiss this example from the book of evidence.

### 4.3. Valencia Basin

The Valencia Basin (VB; Fig. 2a) is an aborted rift formed during the late Oligocene-early Miocene opening of the back-arc Liguro-Provençal Basin (e.g. Jolivet et al., 2006). It is located between the Spanish Ebro Margin to the west and the Balearic Promontory to the east, while it grades into the Liguro-Provençal Basin to the E/NE (Fig. 7e; Maillard and Mauffret, 2006; Maillard et al., 2006).

Numerous exploratory boreholes exist on the western Ebro margin. These boreholes, tied to seismic data, confirm that MSC-related sediments on the northwestern shelf are missing (Fig. 7e; Urgeles et al., 2011; Pellen et al., 2019). The only MSC feature present is a prominent erosional surface (the MES) incising Serravallian-early Messinian sediments (Urgeles et al., 2011). By contrast, on the southwestern and southern part of the margin, well data show the presence of evaporitic sediments (e.g. Delta L and Golfo de Valencia D1 boreholes; Del Olmo, 2011; Del Olmo and Martín, 2016; Lozano, 2016). Del Olmo and Martín (2016) described these evaporites as primary selenites and ascribed them to MSC Stage 1 (their unit M7). Lozano (2016) described the same evaporitic deposits in the same boreholes as 'white anhydrites', leaving open the question as to whether the anhydrite is primary (sabhka's) or secondary at the expense of a primary gypsum facies. On the eastern margin of the VB boreholes and seismic studies suggest there are no MSC units with only a prominent MES cutting pre-MSC sediments (Driussi et al., 2015; Raad et al., 2021). All authors conclude that the shelves of VB were exposed to subaerial erosion during and following the main drawdown.

MSC deposits are also absent along the slopes and, where present, consist of coarse- and fine-grained terrigenous facies filling valleys largely related to fluvial incision (Fig. 7e; Stampfli and Höcker, 1989).

A different situation features in the depocenter. Despite its present-day depth of > 2000 m, no MU is observed in the depocenter, as the salt pinches-out at the edge of the deeper Provençal Basin (Fig. 7e). Only the seismic properties of UU suggest that it is roughly continuous from the Provençal Basin into the VB (Fig. 7e; see subsection 4.2), although it thins from ~1000 m to < 500 m. The UU is characterized by aggrading and onlapping geometries towards the slopes, where it also thins out until it disappears along the middle-upper slope (Fig. 7e; Maillard et al., 2006; Cameselle and Urgeles, 2017). Maillard et al. (2006), Urgeles et al. (2011), Cameselle et al. (2014) and Cameselle and Urgeles (2017) all stated that the UU formed during an important Mediterranean-level lowstand (~1000 m). Several Complex Units (CU), with different origin, have been observed and described as belonging to the MSC (Cameselle and Urgeles, 2017).

DSDP Site 122, drilled at the mouth of a valley incision, recovered a few meters of sand-gravels made of well-rounded basalt, marine limestones, nodules of crystalline gypsum and mollusk fragments in a clay-sility matrix rich in deep-water benthonic foraminifera and early Pliocene nannofossils, all interpreted as erosional products of the VB seabed (Ryan et al., 1973). The uppermost Messinian in two industrial wells (Ibiza Marino and Gabriel boreholes; see Lozano, 2016) is represented by intercalations of clastic gypsum/anhydrite and marls (unit M8-P1 of Del Olmo and Martín, 2016). These are interpreted as turbidites sourced from the shelf and/or slope during a lowstand phase of the Mediterranean base level (Del Olmo, 2011; Del Olmo and Martín, 2016; Cameselle and Urgeles, 2017).

In seismic profiles, the UU/PQ transition (M-reflector or TES) is locally both sharp and smooth (in more distal settings) and erosional (in more proximal settings; Fig. 7e). Maillard and Mauffret (2006) indicate that the smooth parts have been caused by increasing fresh water influx during the Lago-Mare phase, leading to dissolution of the evaporites, and the rough erosional segments are of subaerial origin. For Cameselle and Urgeles (2017), the top of the UU is a smooth and conformable downlap surface, representing the rapid inundation of the basin with only local minor erosional features.
## 4.4. Balearic Promontory

Sticking out from the surrounding deep-water locations, the Balearic Promontory (BP; Fig. 2a) is a continental high that has undergone tectonic extension since the late Serravallian (Roca and Guimera, 1992; Sabat et al., 2011). During the Messinian, it comprised in topographic lows/subbasins at different water depths and separated by structural highs/sills (Maillard et al., 2014; Driussi et al., 2015; Roveri et al., 2019b; Raad et al., 2021). The area has been the subject of multiple studies (Maillard et al., 2014; Driussi et al., 2015; Ochoa et al., 2015; Roveri et al., 2019b; Raad et al., 2021) and several controversies arose after the publication of Roveri et al. (2019b).

The first controversy concerns the Messinian paleodepth of the BP's subbasins. According to Roveri et al. (2019b) the subbasins were shallow during the Messinian and acquired today's paleodepths following a strong post-MSC subsidence; Maillard and Mauffret (2011), Maillard et al. (2014) and Raad et al. (2021), instead, consider the tectonic movements in the BP to have been minor since the late Miocene (Messinian) and the region to have been already structured as it is today during the MSC. Well-to-seismic ties in the shallower basins closer to the Spanish coast (i.e. Bajo Segura, San Pedro and Elche basins) comprise up to 14 Stage 1 primary gypsum-marl cycles similar to the onshore PLG unit (Lugli et al., 2010) truncated at the top by the MES (Soria et al., 2008a, 2008b; Ochoa et al., 2015). At first, Ochoa et al. (2015) concluded that all sediments overlying the MES are Pliocene in age. A later re-appraisal of the same downhole logging data and cuttings led Ochoa et al. (2018) to attribute the first ~13 m-thick micritic and evaporite-free sediments to the late Messinian (stage 2 or 3 of the MSC according to the authors). The MSC stratigraphy of the shallowest offshore basins of the BP closely resembles that described from cores and outcrops onshore Mallorca (see subsection 3.4; Roveri et al., 2019b).

High resolution seismic reflection data in the Central Mallorca Depression (CMD) highlighted up to 500 m of MSC deposits made of a Bedded Unit (BU) and a thin salt layer (Maillard et al., 2014; Driussi et al., 2015). This BU has never been drilled and, therefore, lacks lithological and chronostratigraphic constraints. Two contrasting chronostratigraphic and lithological interpretations are proposed: Roveri et al. (2019b) ascribed these sediments to Stage 2 and 3 and suggested that only reworked evaporites and halite are present. By contrast, following the seismostratigraphic description of Maillard et al. (2014), Ochoa et al. (2015) and Raad et al. (2021) inferred the presence of Stage 1 gypsum also in the CMD.

Raad et al. (2021) made a step forward by disclosing a possible tripartition of the BU unit (Fig. 7f). In their seismostratigraphic framework, Raad et al. (2021) noticed that the uppermost evaporite-bearing unit (called BU3), ~120 m-thick in the CMD, has geometric, stratigraphic and facies analogies with the astronomically-tuned UG unit of the Caltanissetta Basin (Fig. 3a) that endorse its attribution to Stage 3. Similar to the UU in the deepest basins (see subsection 4.2), BU3 consists of up to 9 low- and medium-amplitude reflectors that are interpreted as alternating terrigenous and gypsum beds (Maillard et al., 2014; Raad et al., 2021). Reflectors are parallel and continuous in the more distal areas, while they appear more chaotic in the more proximal sectors (Raad et al., 2021). The base of BU3 coincides with the erosional top of the salt, interpreted as created by salt exposure, dissolution and locally salt gliding towards the depocenter (Fig. 7f; Raad et al., 2021). By contrast, the top of BU3, which corresponds to the Miocene/Pliocene boundary, is largely flat without signs of erosion (Fig. 7f; Maillard et al., 2014; Raad et al., 2021). An irregular geometry is sometimes visible, but is likely to be related to deformation of the underlying salt (Fig. 7f; Raad et al., 2021).

## 4.5. Tyrrhenian Basin

The Tyrrhenian Basin to the east of Sardinia is a back-arc basin that opened by continental rifting and oceanic spreading related to the

eastward migration of the Apennine subduction system from middle Miocene to Pliocene times (Gaullier et al., 2014; Lymer et al., 2018; Loreto et al., 2020 and references therein). Three main domains are traditionally identified (Lymer et al., 2018 and references therein): 1) the East Sardinia Basin, with present-day water depths between 200-2000 m consisting of a system of seamounts and depressions that do not contain MSC sediments (Lymer et al., 2018); 2) the Cornaglia Terrace (2000-3000 m deep), a wide, flat area with dispersed structural highs; 3) the Tyrrhenian Basin s.s., with water depths varying from 3000-3600 m. Whether the Tyrrhenian Basin acquired the present-day topography and hypsometry before the MSC (Lymer et al., 2018) or at least part of it (e.g. Eastern Sardinia margin, where Site 654 is located, and Northern Tyrrhenian) was much shallower (possibly comparable to the Caltanissetta Basin; Roveri et al., 2014b) and underwent extension and subsidence during the Messinian-Pliocene (e.g. Kastens and Mascle, 1990; Loreto et al., 2020) is still unresolved.

The MSC units in seismic profiles from the Tyrrhenian Basin (Fig. 7g) are very similar to the ones described in the Algero-Balearic and Liguro-Provençal basins (Fig. 7b-c; Gaullier et al., 2014; Lymer et al., 2018). ODP Sites 652, 653 and 654 confirmed the seismic-inferred lithological nature of UU as consisting, of gypsum-mudstone alternations (8 are counted at Site 654; Kastens et al., 1987; Borsetti et al., 1990; Roveri et al., 2014b). Intercalations of ripple-cross-laminated, fine-grained, azoic sandstones occur within the mudstone intervals in places (Cita et al., 1990; Iaccarino and Bossio, 1999). In some mudstone samples, the ostracod Cyprideis sp. (Site 654) and Candona sp. (DSDP Hole 974B) and the foraminifera Ammonia beccarii and Ammonia tepida have been found, indicating a shallow-water (< 50 m) brackish environment (see subsections 5.1 and 5.4; Cita et al., 1990; Iaccarino and Bossio, 1999). $^{87}$Sr/$^{86}$Sr isotope ratios of UU gypsum and planktic foraminifera of the overlying Pliocene (Unit 1 at Site 654) show values much lower (from 0.708627 to 0.708745) and roughly equivalent (from 0.708935 to 0.709112) to coeval ocean water (~0.709020-30; McArthur et al., 2012), respectively (Müller et al., 1990; Müller and Mueller, 1991). Similar $^{87}$Sr/$^{86}$Sr values were obtained from the gypsum cored at Site 652 (0.708626-0.708773; Müller and Mueller, 1991).

The Miocene/Pliocene boundary at DSDP Site 132 is placed above a cross-bedded sand rich in quartz, mica, pyrite, rounded fragments of gypsum and specimens of Ammonia beccarii and Elphidium macellus (Ryan et al., 1973). In the adjacent ODP Site 653 a similar sandstone is found slightly below the biostratigraphically-defined Messinian/Pliocene boundary and ~70 cm of grey mudstones with foraminifera and nannofossils of undisclosed provenance are squeezed in between (Cita et al., 1990). These mudstones also contain rare dwarf planktic foraminifera (Globorotalia acostaensis, Orbulina universa, and Globigerina bulbosa; Cita et al., 1990).

Overall, the uppermost Messinian sediments of the Tyrrhenian Basin are interpreted as having been deposited in lakes with periodic episodes of increased salinity and dilution under the strong influence of high energy rivers and, perhaps occasionally, of the Atlantic (Cita et al., 1990; Müller et al., 1990; Müller and Mueller, 1991).

## 4.6. Ionian Basin

The deepest basin in the Mediterranean today is the Ionian Basin, with its lowest point at 5,267 meters. The so-called Ionian Abyssal Plain (IAP) is bounded on all sides by pre-MSC structural highs (Fig. 2a; Camerlenghi et al., 2019): the Malta Escarpment to the west; the Medina Escarpment to the south separating it from the Gulf of Sirt (Fig. 8a); the Gargano-Pelagosa and/or Otranto sill to the north dividing it from the Adriatic Foreland and, finally, an unnamed sill to the east, separating the IAP from the Levant Basin. Evidence from recent high-resolution seismic studies across the region have been used to support Stage 3 desiccation models (e.g. Hsü et al., 1978a, 1978b; Bowman, 2012; Micallef et al., 2018, 2019; Camerlenghi et al., 2019; Spatola et al., 2020).
<figure><img src="image_12.png" /></figure>

Site 374

<figure><img src="image_13.png" /></figure>

<figure><img src="image_14.png" /></figure>

<figure><img src="image_15.png" /></figure>

<figure><img src="image_16.png" /></figure>

<figure><img src="image_17.png" /></figure>

<figure><img src="image_18.png" /></figure>

<figure><img src="image_19.png" /></figure>

<figure><img src="image_20.png" /></figure>

<figure><img src="image_21.png" /></figure>

(caption on next page)
**Fig. 8.** Seismic profiles from intermediate-deep Eastern Mediterranean basins containing MSC markers/units (see Fig. 7 for abbreviations). (a) High-resolution seismic line MS27 imaging the PQ and the uppermost MSC's UU and MU in the Ionian Abyssal Plain and Gulf of Sirt (modified from Camerlenghi et al., 2019). Note how the MSC units are thinner, more difficult to distinguish and more deformed in the Sirt Abyssal Plain than in the IAP. PQ, UU and MU all onlap the structural highs of the Medina Ridge and VHS-2 sill. (b) High-resolution imaging of the lower part of the Plio-Quaternary (PQc unit) and upper part of the Messinian (UU) in the IAP (Camerlenghi et al., 2019). The MSC-PQ boundary is a highly irregular surface, describing apparent V-shaped incisions (symbol V) of controversial origin (see subsection 4.6.1 for insights). Note the coherent deformation of PQc with the underlying MSC sequence and the absence of fluvial facies within the incisions (Unit II is made of lower Pliocene dolomitic marls recovered in Site 374 drilled nearby the seismic line; see text). (c) Multichannel seismic reflection profile MEM-07-203 running approximately parallel to the Malta Escarpment and showing the relationship between Unit 2 of Micallef et al. (2018) with the overlying and underlying Zanclean and Messinian sediments, respectively (modified from Spatola et al., 2020). (d) Uninterpreted (left) and interpreted (right) seismic profiles showing the cyclic and channelized nature of the uppermost Messinian observed in the offshore Sirt Basin (modified from Bowman, 2012). (e) Interpreted 2D regional WNW-ESE seismic profile crossing the continental shelf and offshore Levant Basin and the Herodotus Abyssal Plain (Jagger et al., 2020). Note the lateral continuity of the Messinian MU. (f) Seismic profile from the Levant Basin showing the 6 sub-units distinguished inside the MU as well as its lower (N-reflector) and upper (M-reflector) boundaries (modified from Gvirtzman et al., 2013). (g) High-resolution seismic reflection image with wireline logs from Aphrodite-2 well illustrating that M-reflector previously considered as top evaporitic sequence and M/P boundary here consists of a ~100-m-thick unit (i.e. Unit 7 of Gvirtzman et al., 2017) in which different layers are distinguished (modified from Gvirtzman et al., 2017). (h) Interpreted and uninterpreted seismic profiles imaging the Mavqi'im and Afiq formations described in the canyons on the Levant continental margin (modified from Ben Moshe et al., 2020).

### 4.6.1. Ionian Abyssal Plain

The typical "trilogy" of seismic units representing the MSC deposition in the Western Mediterranean is recognized also in the IAP by Gallais et al. (2018) and Mocnik et al. (2018). However, Camerlenghi et al. (2019) states the MSC sequence in the IAP is ~1,300 m-thick and composed of only two units (Fig. 8a). The lowermost 150-700 m-thick Mobile Unit (MU) is seismically transparent without discernible bedding and with diapiric structures, all features diagnostic of halite. By contrast, the 350-1,000 m-thick Upper Unit (UU) alternates highly reflective with acoustically transparent reflectors (Figs. 8a-b), similar to those described of the UU sequences of the Western Mediterranean (Figs. 7b-c). These are therefore assumed to represent gypsum-mudstone cycles (Camerlenghi et al., 2019). The uppermost 80 m of UU has been recovered from DSDP Site 374 (Hsü et al., 1978b), confirming the presence of gypsum (both primary cumulate and clastic; Lugli et al., 2015) alternating with mudstones (Unit III; Hsü et al., 1978b). These mudstones are largely barren of *in situ* fossils. However, the presence of some foraminifera and siliceous microfossils led Cita et al. (1978) and Hsü et al. (1978a) to suggest that sporadic marine incursions, possibly from the Indian Ocean, took place during Stage 3. Similar to Site 372, the basal part of Hole 374 intercepted one thin (~3.5 m) halite layer within the UU (Hsü et al., 1978b).

The UU and the overlying Zanclean (subunit PQc of Camerlenghi et al., 2019) reflectors are conformably folded throughout most of the abyssal plain, locally showing chaotic internal structure, irregular folding mimicking V-shaped incisions and truncations (Fig. 8b; Camerlenghi et al., 2019). These features are interpreted by Camerlenghi et al. (2019) as fluvial valleys carved in subaerially exposed evaporites by the Eso-Sahabi fluvial system, the closest fluvial drainage network to the area (see Micallef et al., 2018) that drained Libya in the late Miocene (Griffin, 2002) and has been traced across the Gulf of Sirt offshore (Sabato Ceraldi et al., 2010; Bowman, 2012). Several arguments counteract this interpretation: 1) the coherent, deformation, mostly of post-Messinian age, of both the UU and the lower Zanclean; 2) the absence of fluvial facies above the bottom of the "valleys", which instead correspond to a mudstone interval that underwent post-depositional dolomitization (Unit II; see below; Fig. 8b); 3) the unlikelihood that the Eso-Sahabi fluvial system managed to bypass the Medina Ridge divide (Fig. 8a). Alternatively, the apparent incisions at the M/P boundary in the IAP may represent post-sedimentary folds resulting from post-Messinian tectonic and/or salt movements-driven deformation (e.g. Mocnik et al., 2018). At Site 374 the Miocene/Pliocene boundary has been recovered (Unit II), but its primary nature (likely a mudstone) is obscured by diagenesis (cementation by dolomite; Hsü et al., 1978b). A lithified dolostone at the (seismic) predicted depth of the M-Horizon is a characteristic of several sites. Sometimes this interval has been recovered (e.g. Sites 125 and 374; Ryan et al., 1973; Hsü et al., 1978b; Comas et al., 1996); at others the hard lithology is inferred by the torqueing of the drill string (resistance to turning) accompanied by

bouncing of the drill bit at the (e.g. Sites 122, 124, 125, 132, 133 and 134; Ryan et al., 1973). Dolomitization was (Hsü et al., 1973a, 1973b; Ryan et al., 1973) and still is (e.g. Ryan, 2009) largely considered a "diagnostic feature of tidal (sabkha) sediments" (Friedman, 1973, pp. 705). Its occurrence at locations with present-day water depth exceeding 2000 m was therefore considered strong evidence that the Mediterranean floor was subaerially exposed prior to the Zanclean marine replenishment (e.g. Ryan et al., 1973). It is, however, now widely accepted that dolomite precipitation is not exclusive of sabkha environments, but rather is a common process also in bottom sediments under a relatively deep water column (see Dela Pierre et al., 2012, 2014 and references therein). In the specific case of the offshore Mediterranean's M/P boundary on the Ionian Abyssal Plain, already in the '70s dolomitization was thought to have occurred after burial (Hsü et al., 1978b), a conclusion recently reinforced by McKenzie et al. (2017).

### 4.6.2. Malta Escarpment

At the foot of the Malta Escarpment, Micallef et al. (2018, 2019) and Spatola et al. (2020) amalgamated the MU and UU into one seismic unit, Unit 3, which is separated from the Plio-Quaternary marine sediments (Unit 1) by Unit 2, a chaotic transparent seismic package (Fig. 8c). Unit 2 has a maximum thickness of 760-860 m, a volume of 1430–1620 km³ and a wedge-shaped geometry that thins eastwards, disappearing before reaching the IAP (Micallef et al., 2018). Micallef et al. (2018) and Spatola et al. (2020) proposed a lithological/sedimentological interpretation of this chaotic body, suggesting it is composed of well-sorted sediments of the Pelagian Platform to the west of the Malta Escarpment, with coarser material at the mouth grading into more distal finer sediments. This chaotic body has recently been attributed to the Zanclean megaflood during its passage from the western to the eastern Mediterranean via a gateway located in south-eastern Sicily (Micallef et al., 2018, 2019; Spatola et al., 2020). Given the rapidity of the reflooding (≤ 2 years, Garcia-Castellanos et al., 2009, 2020), this interpretation implies rapid mass deposition. Other interpretations of this Unit 2 include being the result of extensive marine mass movement (Polonia et al., 2011), being folded UU (Butler et al., 2015) or being a complex unit built during lower sea level phases (Lofi et al., 2011a, 2011b).

### 4.6.3. Gulf of Sirt

The Gulf of Sirt (or Sirt embayment; Figs. 2a, 8a), the offshore extension of the Sirt Basin onshore Libya (Griffin, 2002 and references therein), is cross-cut by high-density seismic and well datasets as a result of oil potential of the region (Fiduk, 2009). However, only few studies have tackled the MSC (e.g. Hallett, 2002; Fiduk, 2009; Bowman, 2012).

In the Sirt embayment the MSC unit(s) is unevenly distributed in sub-basins controlled by a pre-existing topography, there is little distinction between the MU and UU, the overall thickness of the MSC unit(s) is lower and the degree of deformation is higher than in the adjacent IAP (Fig. 8a; Camerlenghi et al., 2019). The presence of halite in the Sirt
embayment is debated, but most authors think it is absent (see Fiduk, 2009; Sabato Ceraldi et al., 2010; Lofi, 2018; Jagger et al., 2020; Fig. 2b). Bowman (2012) distinguishes seven seismic units within the MSC-related sequence (Fig. 8d). On the basis of high-resolution 3D and 2D data, each seismic unit has been interpreted consisting of clastics filling erosional channels cutting up to 100 m deep and wide (Fig. 8d) and evaporites (gypsum and anhydrite) alternating with precessional cyclicity (Bowman, 2012). The presence of anhydrite in the topmost part of the sequence is confirmed by the B1 NC 35A well (Hallett, 2002). Sabato Ceraldi et al. (2010) and Bowman (2012) envisaged a three-step evolution of each unit: 1) evaporitic deposition during precession maxima in a dried out Sirt embayment; 2) erosion of the valleys during the arid-wet transition fed by the Eso-Sahabi paleofluvial system; 3) filling of the valleys with the fluvial sediments during the wet phase. Based on this chronostratigraphic interpretation, the evaporite cycles should be time equivalent to most of Stage 3, with the upper four seismic units reflecting the Lago-Mare phase (Fig. 1a).

## 4.7. Levant Basin

### 4.7.1. Abyssal plain

Old seismic data in the Levant Basin show an up to 2 km-thick, high velocity, acoustically transparent sequence bounded by the N-reflector at the base and the M-reflector at the top (Figs. 8e-f; Ryan, 1978; Netzeband et al., 2006). This sequence thickens and extends for tens of kilometers towards WNW and thins eastward along the continental margin (Fig. 8e), where the N and M-reflectors converge forming the condensed MSC section of the Mavqiim and Afiq formations (described in subsection 4.7.2; Gvirtzman et al., 2017; Manzi et al., 2018). High resolution 2D and 3D industrial seismic and well data from the southern Levant Basin revealed that this transparent sequence is largely made of pure halite with internal stratification picked out by diatomite, clay- and clastic-rich layers that allowed the division of the sequence into six sub-units, basinward-tilted and truncated at the top by the flat TES (Fig. 8f; Gvirtzman et al., 2013, 2015, 2017; Feng et al., 2016, 2017; Manzi et al., 2018; Meilijson et al., 2018, 2019). Clastic beds ~3 to 20 m-thick are abundant in the highly reflective and chaotic Unit 5 (i.e. Interbedded evaporites of Meilijson et al., 2019; MC2 unit of Feng et al., 2016; Figs. 3b, 4a), where they are interbedded with evaporites (probably halite with minor occurrences of anhydrite) varying in thickness from ~6 to 30 m (Manzi et al., 2018; Meilijson et al., 2019). Paleontological analyses of cuttings probably belonging to one of the clastic beds revealed the presence of a few mollusk fragments, ostracods, echinoid spines and a relatively rich assemblage of benthic and planktic foraminifera which Meilijson et al. (2019) concluded to be reworked. Based on seismic and well-log data, clastic intercalations (probably of clays, silts and sands) within a halite-dominated sequence are thought to persist in the overlying Unit 6, although they diminish in thickness and frequency (Gvirtzman et al., 2013; Feng et al., 2016; Meilijson et al., 2019). The expression of the end of the MSC is highly controversial. Until recently, the M-reflector of Ryan (1978) (later renamed as the Top Erosion Surface, TES; Lofi et al., 2011a, 2011b) bounding Unit 6 at the top was considered to be the Miocene/Pliocene boundary (Fig. 8f; Ryan, 1978; Gvirtzman et al., 2013; Feng et al., 2016). Instead, Gvirtzman et al. (2017) showed that in higher resolution seismic data the M-reflector/ TES is a bundle of reflectors forming a distinct layer (Unit 7) overlying a truncation surface (i.e. Unit 6/7 boundary) that they re-labelled intra-Messinian truncation surface (IMTS; Fig. 8g) and ascribed to subaqueous dissolution rather than subaerial incision (e.g. Bertoni and Cartwright, 2007; Lofi et al., 2011a, 2011b; Kartveit et al., 2019; Madof et al., 2019). This conclusion was recently corroborated by the independent study of Kirkham et al. (2020). Analysis of gamma-ray and resistivity logs in three deep basin wells (Aphrodite-2, Myra-1, Sara-1; Fig. 2b) and correlation with the Or-South-1 well (located between the deep basin and the shelf) showed that Unit 7 maintains a constant thickness of ~100 m-thick and consists of clastic-rich anhydrite of undisclosed provenance.

Meilijson et al. (2019)'s lithological interpretation of industrial boreholes slightly farther to the NE (Fig. 2b) give Unit 7 a significantly smaller thickness (5 m; Fig. 3b). Independent studies offshore Lebanon and Syria (Kartveit et al., 2019; Madof et al., 2019) describe a unit (Nahr Menashe complex) interpreted as a thicker (up to 300 m; Madof et al., 2019), but lateral equivalent of Gvirtzman et al. (2017)'s Unit 7. Based on its channelized morphology identified upslope near the Lebanese coast, Kartveit et al. (2019) and Madof et al. (2019) interpreted the Nahr Menashe Unit and the IMTS underneath as fluvial in origin, deposited/ formed on a subaerially exposed floor of the Levant Basin. Six (Madof et al., 2019) to eight (Madof and Connell, 2018) lobes were identified and are proposed to have stacked with precessional frequency. The Nahr Menashe sequence has been correlated by the same authors with the Abu Madi Fm. located within the Messinian canyons offshore Egypt (Abdel Aal et al., 2000; Loncke et al., 2006; Abdel-Fattah, 2014), the Handere Formation offshore Turkey (Radeff et al., 2017) and with the Eosahabi deposits offshore Libya (Bowman, 2012). This interpretation implies a low base-level during the final stage of the MSC.

Manzi et al. (2018) and Meilijson et al. (2018, 2019) attempted astronomical dating of the abyssal MSC succession of the Levant Basin by integrating biostratigraphy on the pre-MU succession, reflector counting within the MU (only Meilijson et al., 2019) and well-log data (Fig. 3b). They achieved two contrasting results that gave rise to an outstanding controversy (Figs. 3b, 4a). Manzi et al. (2018) proposed that Stage 1 in the deep Levant is represented by a foraminifera-barren, evaporite-free shales interval labelled FBI (foraminifer barren interval) observed in the deep Aprodite-2 well and in the more proximal Myra-1 well. In this interpretation Unit 7 comprises the whole of Stage 3 (with the IMTS corresponding to the Stage 2/3 transition) and all halite deposition took place in ~50 kyr estimated during Stage 2 of the MSC (Fig. 1a; Roveri et al., 2014a). By contrast, the FBI is not present in the Dolphin well targeted by Meilijson et al. (2019), which is located in an intermediate position between the Aprodite-2 and Myra-1 wells studied by Manzi et al., 2018; Fig. 2b). Instead, in the Dolphin well a relatively open-marine, foraminifera-rich sequence extends below the (conformable) base of the MU, placed in correspondence to a ~2 to 5.5 m-thick anhydrite/shale (Unit 0; Manzi et al., 2018 and Meilijson et al., 2018, respectively). Astronomical tuning of the ~33 cycles counted in the MU in the Dolphin well, which are assumed to be precessional, results in the Main Halite body (i.e. Unit 0-4 of Gvirtzman et al., 2013 and Manzi et al., 2018) spanning MSC Stage 1 and 2, the Interbedded Evaporites/Unit 5 covering substage 3.1 and the Argillaceous Evaporites/Unit 6-7 to encompass the Lago-Mare phase (Figs. 3b, 4a). In this interpretation from Meilijson et al. (2019), halite deposition in the Levant Basin started in Stage 1 and persisted throughout the entire MSC, including Stage 3, during which more allochthonous material was delivered to the basin (Fig. 3b). Madof and Connell (2018) and Madof et al. (2019) also attempted an astronomical tuning of the Nahr Menashe Unit, concluding that it spans throughout substage 3.2 and part of substage 3.1. Feng et al. (2016) claim, however, that the impressive thickness of clastics found in the Levantine MU is more indicative of distinct short-term events (shorter than the precession cycle) associated with transport of extraordinary power and magnitude.

Late Messinian sediments have also been recovered at several DSDP (129, 375, 376; Ryan et al., 1973; Hsü et al., 1978b) and ODP Sites (965, 967, 968; Emeis et al., 1996), but the assignment of the retrieved sediments to seismostratigraphic units is problematic. Nevertheless, they provide several key nuggets of precious information about the Stage 3 paleoenvironment:

- Sites 965 and 966, located roughly on the crest of the Eratosthenes Seamount, just south of Cyprus (Fig. 2b), recovered soil structures above the evaporites indicating exposure above sea level (Robertson, 1998a, 1998b; Maillard et al., 2011a; Reiche et al., 2016).

- ODP Sites 967 and 968, located at the base of the northern and southern slope of Eratosthenes Seamount (Fig. 2b), respectively,
revealed the presence, within the MSC interval, of calcareous turbi-
dites with Cyprus-derived clasts and clays containing *Ammonia tep-
ida*, *Cyprideis pannonica* and *pulmonate gastropods* (Blanc-Valleron
et al., 1998; Robertson, 1998a,b; Reiche et al., 2016).

- Abundant *Cyprideis pannonica* specimens were also recovered from DSDP Sites 375 and 376 drilled on the crest of the Florence Rise, west of Cyprus (Fig. 2b; Hsü et al., 1978b).

- Abundant, well-preserved *Ammonia* tests and *Cyprideis* specimens are also known from Site 129A (Fig. 2b), occurring with dwarf planktonic foraminifera (Ryan et al., 1973).

All the evidence listed above suggest that a base-level fall leading to
subaerial exposure occurred at some point(s) during Stage 3 in the
Eastern Mediterranean (Ryan, 2009). However, it must be kept in mind
that both the Florence Rise and Eratosthenes Seamount are likely to have
been much more elevated during the Messinian than they are today
because of Pliocene-Quaternary subsidence related to the Cyprus sub-
duction zone (Robertson, 1998a, 1998b; Maillard et al., 2011a; Reiche
et al., 2016).

Sites 375 and 376 display several discrete layers of primary and clastic gypsum interbedded in the *Cyprideis*-rich mudstones (McCulloch and De Deckker, 1989; Lugli et al., 2015). This succession shares several similarities with sites drilled in the Western Mediterranean (e.g. ODP 654A) and Ionian Basin (DSDP 374), where they have been correlated with the seismic Upper Unit (see subsections 4.5 and 4.6). This may suggest that a Western Mediterranean-like gypsum-bearing UU was also locally deposited in the easternmost abyss of the Mediterranean (see Güneş et al., 2018).

4.7.2. Levant continental margin

Evaporitic and non-evaporitic deposits are buried beneath PQ de-
posits (Yafo Formation) along the Levant continental margin, where
they are mostly preserved within canyons carved underwater in pre-
Messinian time (Druckman et al., 1995; Lugli et al., 2013). Within the
Afiq canyon, Druckman et al. (1995) distinguished three formations in
the Messinian sequence: the evaporitic Mavqi'im Formation, 115 m-
thick and mostly composed of anhydrite in places interbedded with
micritic limestones; the Be'eri Formation, comprising gypsum; the Afiq
Formation, varying in thickness from 30 to 90 m and consisting of
conglomerates, sandstones and marls interpreted as representing fluvial
and lacustrine-marsh environments (Druckman et al., 1995). The Afiq
Fm. is only present in the deepest portions of the canyon where it
 overlies the Mavqi'im Fm. By contrast, the Be'eri gypsum is only found
 along the canyon shoulders covered by the Pliocene, at elevations > 600
 m with respect to the Mavqi'im Fm. Based on Sr values, Druckman et al.
 (1995) attributed the Mavqi'im Fm. to MSC Stage 1, the Be'eri Fm. to
 substage 3.1 and the Afiq Fm. to the Lago-Mare phase. These authors
 also suggested that gypsum precipitation occurred under subaqueous
 conditions, with the water level ~600 m (i.e. the difference in altitude
 between the Mavqi'im and Be'eri fms.) higher during the deposition of
 the Be'eri Fm. Two base-level falls of approximately the same magnitude
 are thought to have occurred between the evaporitic phases and after
 Mavqi'im deposition. A lowstand phase was therefore in force during
 Afiq deposition (Druckman et al., 1995).

However, combining stratigraphic, sedimentological and geochem-
ical (Sr isotopes) considerations, Lugli et al. (2013) revealed the clastic
nature of both the Mavqi'im and Be'eri evaporites and suggested the
persistent drowning of the canyon(s), filled with turbidites (Lugli et al.,
2013). Due to the presence of clastic evaporites, Gvirtzman et al. (2017)
suggested that the Mavqim Formation is a condensed section encom-
passing MSC Stage 2 and early Stage 3, while the evaporite-free Afiq
Formation represents the Lago-Mare phase.

Ben Moshe et al. (2020) ascribed (at least part of) the Afiq Fm. to the
whole of Stage 3 as a wedge-shape clastic complex lying on top of the
Mavqi'im Fm. and with the basal surface corresponding to the correla-
tive conformity of the MES developed landward, at the expense of the

Mavqi'im Fm (Fig. 8h). Ben Moshe et al. (2020) distinguished a lower sub-unit composed of clastic gypsum and with fore-stepping and down-stepping internal geometry typical of progradational wedges, and an upper sub-unit containing anhydrite fragments and marls with Lago-Mare fauna (e.g. *Cyprideis torosa*; Rosenfeld, 1977) and with seismic characteristics typical of a transgressive systems tract. Incisions are reported throughout the Afiq Fm. at different depths, while erosional surfaces bound both sub-units (Ben Moshe et al., 2020). In particular, the surface capping the upper subunit and correlated to the M horizon or TES basinward (see subsection 4.7.1) shows dendritic drainage patterns of gullies and channels (Ben Moshe et al., 2020).

Ben Moshe et al. (2020) identify the variation of base level specif-
ically during Stage 3 by analyzing the morphology of truncation surfaces
bounding the Afiq Formation on the continental margin of the Levant
Basin. This suggests high amplitude fluctuations of base-level in the
order of one hundred meters, with development of subaerial erosion
surfaces and the deposition of clastics and incision by fluvial drainage
systems that occurred during the lowstand phases, while aggradational
units (of unknown lithological nature) accumulated during the high-
stand phases. According to their analysis, base level dropped down to a
maximum 535 m during Afiq deposition (i.e. below the maximum 430 m
estimated paleodepth of the Sicily Sill; Garcia-Castellanos et al., 2009),
implying hydrological disconnection between the Eastern and Western
basins at various times during Stage 3. A regression to 615-885 m is
proposed to have occurred at the top of the Afiq Fm., pre-dating the
abrupt refilling at the base of the Zanclean (e.g. Micallef et al., 2018,
2019; Garcia-Castellanos et al., 2020; Spatola et al., 2020).

4.8. Summary of the offshore Stage 3 record

Knowledge of the Stage 3 sequence offshore is mainly based on the
integration of seismic interpretations and analysis of material recovered
from fragmentary and unevenly distributed DSDP/ODP/industrial
cores.

- MSC sediments are absent on the eroded continental shelves bordering the deep basins, except in the Messinianthalwegs and at their mouth. Here the PQ lies directly above the MES which, in turn, cuts through the pre-MSC deposits (Fig. 7e). A similar stratigraphic arrangement is found along the middle-upper slopes (Fig. 7d), although the presence of a thin, possibly incomplete UU in morphological depressions is sometimes postulated. Seamounts also lack MSC Stage 3 sediments and are strongly incised by the MES (Fig. 7g).

- The thick UU is widespread and roughly present everywhere in the abyssal plains from west of the Alborán volcanic arc to the eastern edge of the Ionian Basin (Fig. 2b). In the abyssal plains its seismic facies appears homogeneous, comprising parallel and relatively continuous high amplitude reflections (Figs. 7b-c). The UU pinches out towards the foot of continental slopes and seamounts (Figs. 7b, d-g), where it can be irregularly bedded or relatively well bedded (Lofi et al., 2011a, 2011b). The uppermost part of the Bedded Units (defined in depressions physically disconnected from the abyssal plains and, therefore, from the UU; e.g. CMD and Corsica Basin; Maillard et al., 2014; Thinin et al., 2016; Raad et al., 2021) shows seismic features comparable to those of the UU.

- Drill Sites revealed that the reflections of relatively high amplitude in seismic profiles correspond to gypsum and mudstone alternations with sporadic intercalations of massive to cross-bedded sandstones. Some mudstone interbeds contain low-diversity assemblages of benthic organisms (ostracods, foraminifera and diatoms) indicative of shallow to neritic environments. Except for dwarfed forms of planktonic foraminifera and the monospecific nannofossil assemblages described
<figure><img src="image_22.png" /><figcaption>**Fig. 9.** Photomicrographs of the key micro-and macrofossils featuring Stage 3 sediments. (a) Scanning electron microscope (SEM) photographs of the more common Paratethyan ostracod species in substage 3.1 and 3.2 sediments (compiled from Stoica et al., 2016, Cosentino et al., 2018 and Sciuto et al., 2018). (b) Photomicrographs of the Paratethyan dinoflagellate cyst *Galea-cysta etrusca* under the optical microscope (left) and SEM (right) (modified from Do Couto et al., 2014 and Grothe et al., 2018). Scale=20 µm. (c) SEM microphotographs of the euryhaline, shallow-water benthic foraminifera *Ammonia beccarii* (1-spiral side, 2-umbilical side) and *Ammonia tepida* (3-spiral side; Carnevale et al., 2019) and of the dwarf fauna of planktonic foraminifera from the Bajo Segura Basin (4; Corbí and Soria, 2016). (d) Photographs in polarized light (crossed nicols) of some specimens of *Ceratolithus acutus* (1-3) described in the Lago-Mare unit of Malaga (1-Do Couto et al., 2014), the Zorreras Mb. of Sorbas (2-Clauzon et al., 2015) and the Colombacci Fm. of the Northern Apennines (3-Popescu et al., 2017) and of destroyed (4) and intact (5) ascidian spicules of *Micrascidiscus* sp. (Golovina et al., 2019). Note that *C. acutus* specimens closely resemble ascidian spicules of *Micrascidiscus* sp., which may lead to misinterpretation of the *C. acutus* (see Golovina et al., 2019), and that pictures 1 and 2 are identical, despite they are attributed to samples taken from two different localities. (e) Articulated skeletons of marine fish from substage 3.1 mudstone horizons in Cyprus (1-Aphanius crassicaudus; Manzi et al., 2016a) and substage 3.2 deposits in Cava Serredi (2-Mugil cf. cephalus; Carnevale et al., 2018).</figcaption></figure>
by Castradori (1998), the rest of planktonic foraminifera and nannofossils are largely regarded as reworked.

* The deep Levant Basin contains a ~1.8-2.0 km-thick MU (Figs. 8e-f), consisting of 6 to 7 seismic units depending on the resolution of the seismic employed. In high resolution seismic data, the lateral equivalent of part of the UU is identified as a ~100-m-thick, clastic-rich, anhydrite layer (Unit 7 of Gvirtzman 1207) offshore Israel, thickening to 300 m offshore Lebanon (Nahr Menashe complex, Madof et al., 2019). The Levant Basin still has major controversies concerning the timing of halite deposition (~50 kyr vs ~550 kyr; Manzi et al., 2018 vs Meilijson et al., 2019), the origin of the clastic accumulations overlying the halite (fluvial vs subaqueous) and the presence or absence of gypsum-mudstone cycles.

* Apart from the halite flow-related deformation, the Miocene/ Pliocene boundary (i.e. UU/PQ transition) is conformable in intermediate (e.g. Balearic Promontory) and deep (WAB, EAB, Algero-Balearic, Liguro-Provençal, Tyrrhenian, Ionian and Levant) depocenters, while it shows signs of erosion on the shelf domain and along the upper-middle continental slopes and seamounts. Clear arguments of floor exposure at the M/P boundary are absent in all drill sites but 978A.

## 5. The paleontological perspective

Paleontological data have been used to define and identify Stage 3 sediments, but have also been a source of profound contention over the interpretation of its paleoenvironmental and paleohydrological nature. Biotic groups impacted by the evolution of the gateways linking the Mediterranean with the Atlantic, Indian Ocean and Paratethys include marine species (e.g. foraminifera, calcareous nannofossils, fish) and brackish water-species (ostracods, fish, mollusks and dinocysts endemic or with affinity to species of the Paratethys region) that were unable to migrate when these corridors were closed, and terrestrial species (e.g. mammals) that, conversely, got across the gateway during periods of exposure (see Colombero et al., 2017; Booth-Rea et al., 2018; Mas et al., 2018b). Analysis of these faunal datasets provides key insights into likely gateway dimensions and the timing of their opening, restriction and closure (e.g. Palcu et al., 2017). Furthermore, they are a key constraint on the water sources likely to have been affecting the Mediterranean during MSC Stage 3.

### 5.1. Ostracods

Ostracods are by far the most prolific faunal group during Stage 3. Brackish species are known from both land sections and deep-sea cores across the whole Mediterranean (see Fig. 2b for sites and references; Fig. 9a). Two characteristic biofacies are commonly distinguished: Biofacies 1 (Bonaduce and Sgarrella, 1999) or *Cyprideis* assemblage (Iaccarino and Bossio, 1999) consists of an monospecific population of *Cyprideis* species or of an oligotypic population dominated by *Cyprideis* species alongside rare specimens of *Tyrrhenocythere ruggierii*, *Loxoconcha kochi*, *Loxoconcha muelleri* and *Caspio cypris alta*); Biofacies 2 (Bonaduce and Sgarrella, 1999) or *Loxocorniculina djafarovi* assemblage (Iaccarino and Bossio, 1999) has a higher species diversity characterized by the abundant occurrence of truly Paratethyan species belonging to the genera *Amnicythere*, *Loxoconcha*, *Loxocauda*, *Cytheromorpha*, *Cyprinotus* and *Tyrrenhocythere* (see species names in Fig. 9a). The number of species reported in the onshore sections is variable, ranging from a dozen (e. g. Caruso et al., 2020) to more than sixty (e.g. Gliozzi et al., 2007; Grossi et al., 2008). This variability is not explained, but it may result from the application of different taxonomic concepts that resulted in the recognition of more or fewer species (Stoica et al., 2016) or from local environmental conditions that differed from basin to basin and resulted in different patterns of colonization.

Compared to the onshore domain, the ostracod fauna offshore is

impoverished. Monospecific assemblages of *Cyprideis* sp. (Sites 372, 129A, 376, 654A, 967, 968; Ryan et al., 1973; Cita et al., 1990) or oligospecific assemblages dominated by *Cyprideis* and rare specimens of *Candona* sp. (Hole 974B, Iaccarino and Bossio, 1999) and L. *muelleri* (Hole 978, Iaccarino and Bossio, 1999) are the more widely reported. Interestingly, these assemblages are always associated with *Ammonia* sp. tests and, in some cases, with other species of shallow-water, euryhaline benthic foraminifera (see subsection 5.4). Only in Hole 975, close to the M/P boundary is a more heterotypic ostracod assemblage found (*Euxinocythere praebaquana*, *Amnicythere idonea*, *Leptocythere limbate*, *Loxocorniculina djafarovi*, *Candona* sp., and *Cyprideis* sp.; Iaccarino and Bossio, 1999) and lacking of euryhaline benthic foraminifera. The likely cause of the widespread barrenness of ostracods in most of the offshore samples is perhaps because environmental conditions in the deep basins (depth and/or salinity) where not suitable to permit population by this benthic fauna (see below for the ecological requirements; e.g. Hsü et al., 1978b in reference to Site 374). Finally, one must bear in mind that studying these organisms require much more material (some hundreds of grams) than the quantity of core sediments usually processed (i.e. ~10 cm³; Iaccarino and Bossio, 1999).

The paleoecology (salinity and depth ranges) of Stage 3 ostracods has been based on both observations of few species that still live in the Caspian and Black seas today and have affinities with the Stage 3 species and on the interpretation of sedimentological, geochemical and mineralogical data of the surrounding sediments (see Gliozzi and Grossi, 2008 and Grossi et al., 2008 for insights). Biofacies 1 is thought to represent very shallow water environments (i.e. <15 m) with salinity fluctuating between mesohaline and hypersaline when the euryhaline *Cyprideis* is dominant. Instead, more stable oligo-mesohaline water is inferred when the other species are more abundant in Biofacies 1. The variegated Biofacies 2, on the other hand, is thought to represent somewhat deeper environments (up to 100 m) and less salty conditions (oligo-low mesohaline; Gliozzi and Grossi, 2008; Grossi et al., 2008; Caruso et al., 2020).

Some authors consider the time when the Paratethyan ostracods arrived in the Mediterranean to be well constrained (e.g. Roveri et al., 2008a; Grossi et al., 2011; Cosentino et al., 2018) by the scarce occurrence of the first Paratethyan immigrant *Loxoconcha muelleri* 20 cm below the ash layer in the Colla di Votta section, which has a $^{238}$U-$^{206}$Pb age of 5.5320±0.0074 Ma (Cosentino et al., 2013), and in the chaotic deposits of the Adana Basin, ascribed to Stage 2 (Faranda et al., 2013). Instead, the first appearance of *Loxocorniculina djafarovi* has been considered to coincide with the biofacies 1-2 shift and to have occurred Mediterranean-wide synchronously at 5.40 Ma (Roveri et al., 2008a; Grossi et al., 2011; Cosentino et al., 2013). Roveri et al. (2008a) also showed Biofacies 2 diversity as increasing linearly through the Lago-Mare phase, reaching its maximum diversity just beneath the Miocene/Pliocene boundary and before disappearing in the Pliocene. Following the claimed synchronicity of the FO of both *Loxoconcha muelleri* and *Loxocorniculina djafarovi*, Roveri et al. (2008a) and Grossi et al. (2011) recognized one biozone in each biofacies: the *Loxoconcha muelleri* Biozone, spanning from 5.59 to 5.40 Ma, and the *Loxocorniculina djafarovi* Biozone, whose boundaries correspond respectively to the first (5.40 Ma) and last occurrence (5.33 Ma) of L. *djafarovi* in the Mediterranean. This biozonation, erected by Grossi et al. (2011), is often used for dating incomplete successions (e.g. Vera Basin; Stoica et al., 2016; Caruso et al., 2020). However, the first appearance of a diversified ostracod assemblage (including *Loxocorniculina djafarovi*) occurred in already cycle 3 of the Sicilian Upper Gypsum at Eraclea Minoa (Fig. 3a; Grossi et al., 2015), which has an astronomical age of 5.45 Ma (Van der Laan et al., 2006) or 5.47 Ma (Manzi et al., 2009). Furthermore, the sudden appearance of Biofacies 2 and its linear, upward increase in diversity have not been recognized in localities like Nijar and Malaga, where biofacies 1 and 2 are found stacked in more than one lithological (possibly precession-controlled) cycle in the Lago-Mare succession (Bassetti et al., 2006; Guerra-Merchán et al., 2010). These findings argue that the appearance of Paratethyan ostracods in the Mediterranean may
not have been synchronous, therefore casting serious doubts upon the
biostratigraphic relevance of the Mediterranean ostracods.

Except for *Cyprideis* specimens, where species attribution is debated (see discussion in Stoica et al., 2016), the affinity of all other ostracod species observed in Mediterranean Stage 3 sediments (Fig. 9a) with those of the Eastern Paratethys basins (i.e. Dacian, Euxinic and Caspian) has been demonstrated in several publications (e.g. Ruggieri, 1967; Gliozzi et al., 2007; Stoica et al., 2016; Sciuto et al., 2018). Only Bassetti et al. (2003, 2006) have questioned the Paratethyan affinity by suggesting that species from the Northern Apennines and Nijar Basin have ambiguous affinities with Paratethyan fauna as described in the mainly Russian literature from the '60-'70s. However, these differences between the late Messinian Mediterranean and Paratethyan ostracods resulted from misidentifications and/or a different use of species nomenclature (Stoica et al., 2016). Recently acquired knowledge of the Pontian assemblages of the Dacian, Euxinic and Caspian basins now permit to trace the provenance of Mediterranean Stage 3 ostracod species from the entire Black Sea region (Stoica et al., 2016) and, for a few species, from the Dacian (Stoica et al., 2013; Lazarev et al., 2020), Caspian (Van Baak et al., 2016) and North Aegean (see references in Krijgsman et al., 2020a) basins.

The means by which the ostracods travelled from the Paratethys to and across the Mediterranean during Stage 3 is as crucial for reconstructing the Stage 3 paleoenvironment as it is poorly addressed in onshore studies or overlooked in seismic and computational studies. Two migratory mechanisms have been suggested:

1) the aerial dispersion of ostracods through the migration of aquatic birds (Benson, 1978; Caruso et al., 2020); this hypothesis was proposed because, in a Mediterranean concluded to have been desiccated, it was the only possible migration mechanism.

2) direct aqueous migration by the ostracods themselves (which are
planktonic in the larval stage) through the establishment of similar
paleoenvironmental conditions; by this mechanism, the dispersion of
Paratethyan ostracod fauna from right across the Mediterranean re-
quires E-W intraconnection and a Mediterranean water-level high
enough to reach the marginal basins (Gliozzi et al., 2007; Stoica et al.,
2016; Sciuto et al., 2018; Sciuto and Baldanza, 2020).

Finally, Carnevale et al. (2006a, 2006b, 2008, 2018) recognized the Paratethyan affinity of the Mediterranean Stage 3 species but, in view of their occurrence with *in-situ* species of marine fish, they suggested that Stage 3 ostracods descended from a Paratethyan stock that migrated into the Mediterranean well before the MSC and survived the extreme salinity conditions of Stage 1 and 2 in marginal, fresher water refugia. In this scenario the brackish water ostracod assemblages found in Stage 3 have no paleoecological significance for Stage 3 paleoenvironment (Carnevale et al., 2006a, 2006b, 2008, 2018). However, there are two, unflagged problems with this hypothesis: 1) the Mediterranean-Central Paratethys connection through the Trans-Tethyan gateway in Slovenia already closed in the early Tortonian (Kováč et al., 2007; Sant et al., 2017; Palcu et al., 2017); 2) No Paratethyan ostracod species have been found in the Mediterranean before the MSC (see Gliozzi et al., 2007).

5.2. *Dinoflagellate cysts*

Dinoflagellate cysts (dinocysts) are the fossil remains of unicellular
protists that live in the upper water column of many water bodies (e.g.
Zonneveld et al., 2013; Mudie et al., 2017). They can be used as pale-
oenvironmental indicators and for biostratigraphy, providing the ages of
speciation and extinction events, as well as supplying evidence of age
diagnostic dispersals of characteristic taxa/assemblages. Influxes of
these microorganisms into a basin may occur as the result of intercon-
nection with another basin and dinocysts can therefore be useful in-
dicators of the open gateways between adjacent basins and the resultant
changes in conditions (e.g. Grothe et al., 2018). In the case of the MSC,
presence of *in situ* marine and/or Paratethys dinocyst assemblages in a
marginal basin are likely to indicate the presence of Atlantic and/or

Eastern Paratethys water (respectively) in the Mediterranean and
(relatively) high water level conditions (e.g. Pellen et al., 2017).

Palynological studies on the late Messinian Mediterranean dinocysts record are rather scarce, confined to a limited number of outcrops (Malaga Basin, Do Couto et al., 2014; Northern Apennines, Bertini, 2006; Popescu et al., 2007; Iaccarino et al., 2008; Cosentino et al., 2012; Pellen et al., 2017; Caltanissetta Basin, Londeix et al., 2007) and deep wells (976B, 977A, 978A and 134B, Popescu et al., 2015). These studies describe substage 3.1 as being barren of dinocysts. By contrast, substage 3.2 dinocyst assemblages are diverse particularly a few meters/tens of meters below the Miocene/Pliocene boundary and show recurrent vertical variation in abundance between brackish, Paratethyan-type taxa and marine stenohaline and euryhaline species. Taxa with Paratethyan affinities are largely considered to be autochthonous by all aforementioned authors. The extent to which reworking may have affected the marine assemblages is more controversial and debated between none (in Malaga and in the Apennines; Popescu et al., 2007; Do Couto et al., 2014; Pellen et al., 2017), partial (in the uppermost part of the Sicilian Upper Gypsum; Londeix et al., 2007) and total (in the Apennines; e.g. Bertini, 2006; Iaccarino et al., 2008; Cosentino et al., 2012). Given the extent of the implications (i.e. re-establishment of a Mediterranean-Atlantic flow or connection earlier than the Zanclean; e.g. Pellen et al., 2017), this is an issue that will require further clarification.

A key dinocyst influencing our understanding of the late Miocene Lago-Mare phase is *Galeacysta etrusca* (Fig. 9b; see Bertini and Corradini, 1998; Popescu et al., 2009 and Grothe et al., 2018 for more insights). This species was originally described from sediments in the Mediterranean (Corradini and Biffi, 1988), but has since been discovered in much older deposits in Paratethys (Magyar et al., 1999a, 1999b). The earliest recorded occurrence of *Galeacysta etrusca* is in sediments from the Pannonian Basin dated at ~8 Ma (Magyar et al., 1999a, 1999b). It subsequently dispersed throughout Paratethys at ~6 Ma and was present in the Black Sea throughout the MSC interval (Grothe et al., 2014, 2018). Despite a Mediterranean-Eastern Paratethys connection that is thought to have been established at ~6.1 Ma (Krijgsman et al., 2010; Van Baak et al., 2016; Grothe et al., 2020), *G. etrusca* is not found in the Mediterranean during MSC Stages 1, 2 and 3.1 (5.97-5.42 Ma; Bertini, 2006, Londeix et al., 2007, Manzi et al., 2007, Iaccarino et al., 2008, Gennari et al., 2013) and is only reported in the uppermost part of the Lago-Mare phase, very close to the transition to the Pliocene (e.g. Bertini, 2006; Londeix et al., 2007; Popescu et al., 2007; Iaccarino et al., 2008; Cosentino et al., 2012; Pellen et al., 2017). This implies that *Galeacysta etrusca* may have migrated from Paratethys into the Mediterranean after 5.42 Ma or that environmental conditions in the Mediterranean and in its marginal basins were only suitable for this species (and more generally the whole dinocysts Paratethyan contingent) to proliferate in the uppermost Messinian. Several authors report multiple occurrences of *Galeacysta etrusca* within the Zanclean (e.g. Clauzon et al., 2005; Londeix et al., 2007; Popescu et al., 2007, 2015; Do Couto et al., 2014; Clauzon et al., 2015), but these interpretations are based on the use of an alternative stratigraphic model for the MSC sections (Fig. 4b; see Grothe et al., 2018 for details).

5.3. *Diatoms*

Among the fresh-brackish organisms found in Stage 3 sediments are also species of diatoms. To date (and to our knowledge), there are no onshore studies that have ever looked for these organisms. By contrast, two samples from DSDP Site 124 in the Algero-Balearic Basin (Fig. 2b) revealed the presence of littoral planktonic forms accompanied by brackish water, and even freshwater, euryhaline, benthonic, and epiphytic species in considerable numbers (Hajós, 1973). Diatoms of undisclosed paleoecological significance are also reported from the ~60 cm-thick mudstone bed between an anhydrite and halite bed found in the last core of Site 134 (Ryan et al., 1973). According to Hajós (1973) and Ryan (2009), the diatoms found in these drill cores attest to an
extremely low salinity and a base level in the Balearic and Valencia basins below wave action. Further study of these indicative species and a wider distribution is required to apply this interpretation more generally.

## 5.4. Foraminifera

A reasonably diverse benthic and planktic foraminiferal assemblage containing no age-diagnostic taxa have been found co-occurring with the brackish Paratethyan fauna in both the onshore and offshore record throughout the Mediterranean (Fig. 2b for localities and references).

The benthic foraminifera assemblage is dominated by euryhaline representatives of the genus *Ammonia*, which today dwell in marginal marine (lagoons, estuaries, fjords and deltas) and lacustrine environments at depths < 50 m and tolerate salinities of up to 50% (Milker and Schmiedl, 2012; Consorti et al., 2020). *Ammonia tepida* and *Ammonia beccari* (Fig. 9c) are by far the most abundant species in both onshore (see Fig. 2b for localities and references) and offshore (e.g. Site 968A, Blanc-Valleron et al., 1998; Sites 375, 376, 965-968, Orszag-Sperber, 2006) localities, where they co-occur with ostracods belonging to Biofacies 1. Other commonly occurring benthic euryhaline taxa are *Elphiidium* sp., *Cribroelphidium* excavatum, *Haynesina* sp., *Nonion* sp., *Quinqueloculina* sp., *Discorbis* sp. and *Trichohyalus* sp., *Brizalina* dentellata, *Bulimina* echinate and *Bolivina* spp. (Ryan et al., 1973; Hsü et al., 1978a, 1978b; Rouchy et al., 2001, 2003, 2007; Iaccarino et al., 2008; Caruso et al., 2020). These species are frequently mixed with poorly preserved and older in age bathyal species (e.g. Caruso et al., 2020).

Planktic foraminifera are represented both by species whose late occurrence pre-dates the MSC (e.g. *Praeorbulina* spp., *Paragloborotalia partimlabiata*, P. siakensis, *Neogloquadrina atlantica* praeatlantica, *Globigerinoides subquadratus*, *Globorotalia saheliana*, *Globorotalia conomiozea*, *Acarinina* sp., *Hedbergella* sp.) and by taxa with extended biostratigraphic ranges (e.g. *Sphaeroidinellopsis* seminulina, *Turborotalita quinqueloba*, *Globorotalia miotumida*, *Globoturborotalita decoraperta*, *Neogloboquadrina acostaensis*, *Neogloboquadrina* spp., *Orbulina universa*, *Globigerinoides trilobus*, *Globigerinoides obliquus*, *Globorotalia scitula*, *Globigerina bulloides*, *G. Mediterranea* and *G. humerosa*; see Fig. 2b for references).

The mixing of foraminifera species with different ecological and salinity requirements and the widespread agreement that the brackish Paratethyan fauna are autochthonous (see subsection 5.1) has always complicated the interpretation of the origin of the foraminiferal assemblages. Among the benthic species, *Ammonia* taxa and the other benthic euryhaline taxa are generally considered autochthonous because they are typically well-preserved and their ecological and salinity requirements could be compatible with those of the Paratethyan ostracods.

The habitat of these benthic foraminifera today in environments both influenced by and disconnected from the open ocean indicates that the Stage 3 sediments in which they occur were deposited in a shallow-water environment subject to salinity fluctuations (Caruso et al., 2020 and references therein), but they do not provide insights into the water provenance. By contrast, their poor preservation, older age and low diversity of the bathyal taxa strongly suggest that these species are reworked (Bassetti et al., 2006; Iaccarino et al., 2008; Caruso et al., 2020). Their mode of life is also incompatible with the shallower water elements of the faunal assemblage. The planktic species which went extinct before the MSC are also undoubtedly reworked (Iaccarino et al., 2008; Caruso et al., 2020). It is more challenging to discriminate between *in situ* and reworked specimens of the long range Neogene taxa. Most of them are considered to be reworked because of their scarcity, their occurrence with *in-situ* brackish organisms and their poor preservation (e.g. Iaccarino et al., 2008; Caruso et al., 2020). A more complex controversy surrounds the long-range dwarf specimens (Fig. 9c) occurring in onshore substage 3.1 (di Tetto Fm. in the Trave section; Iaccarino et al., 2008) and Lago-Mare sediments (Upper Mb. of the Nijar Feos Fm., Fortuin and Krijgsman, 2003; Aguirre and Sánchez-Almazo, 2004;

Bassetti et al., 2006; Sorbas Basin, Roveri et al., 2019a; Bajo Segura Basin, Corbí and Soria, 2016; Colombacci Fm. in Northern Apennines localities, Casati et al., 1976; Colalongo et al., 1976; Rio and Negri, 1988; Popescu et al., 2007; Cyprus, Rouchy et al., 2001) and in some offshore localities (e.g. Sites 124, 125, 129A, 132, 134, 372, 376, 653, 974B, 975, 978; Cita, 1973; Cita et al., 1978; Kastens et al., 1987; Cita et al., 1990; Iaccarino and Bossio, 1999). This fauna is variably interpreted as:

1. reworked and size-sorted during transport, therefore lacking any paleoenvironmental significance (e.g. Kastens et al., 1987; Iaccarino and Bossio, 1999; Fortuin and Krijgsman, 2003; Bassetti et al., 2006);

2. *in situ* and indicating normal marine conditions (Aguirre and Sánchez-Almazo, 2004; Braga et al., 2006) or temporary Atlantic incursions (Rouchy et al., 2001);

3. *in situ* and indicative of high-stress environments (Keller and Abramovich, 2009), such as restricted and/or diluted marine environments (Corbí and Soria, 2016; Corbí et al., 2016, 2020). However, the paleoecological significance of dwarfism in foraminifer tests is not well understood and, given its potential implications for the Lago-Mare environment, it needs to be explored in greater detail.

## 5.5. Calcareous nannofossils and the *C. acutus* conundrum

Calcareous nannofossils are the fossil remains of coccolithophores, single-celled marine algae which dwell in the eutrophic and photic zone of the ocean (e.g. Ziveri et al., 2004). The potential recognition of marine calcareous nannofossils in marginal Stage 3 deposits would therefore have implications for the Mediterranean base-level and the hydrological riddle of MSC Stage 3. However, like foraminifera and dinocysts, the *in situ* versus reworking issue also impacts the nannoflora.

MSC Stage 3 is crossed by three important nannofossil bio-events astronomically calibrated in the ocean record: the top of *Discoaster quinqueramus* at 5.537 Ma, the base of *Ceratolithus acutus* at 5.36 Ma and the top of *Triquetrorhabdulus rugosus* at 5.231 Ma (Backman et al., 2012; Agnini et al., 2017). Most of the (few) studies that addressed the nannoflora component of Stage 3 deposits did not report taxa belonging to the biozones defined by these bio-events, but only taxa of Cenozoic and Cretaceous age, clearly physically reworked (e.g. Sites 132, 134, 653, 654A, 967A, 969B, Ryan et al., 1973; Hsü et al., 1978b; Müller et al., 1990; Castradori, 1998; Piedmont Basin, Trenkwalder et al., 2008; Violanti et al., 2009; Trave, Fonte dei Pulcini and Stingeti sections and Mondragone well in the Apennines, Cosentino et al., 2006, 2012, 2018; Iaccarino et al., 2008). An exception is the nannoflora observed in the uppermost Messinian sediments at Sites 978A, 975B and 967A (Levant Basin; Fig. 2b). Here, among the plethora of reworked and long-ranging Neogene taxa, Castradori (1998) reported the anomalous abundance of *Sphenolithus* spp. (mostly *Sphenolithus gr abies/moriformis*). Although the assemblage points to the absence of a primary marine signature, the unlikely possibility that reworking and/or sorting lies behind the observed peak of *Sphenolithus* spp. led Castradori (1998) to conclude that at least one incursion of marine water occurred during the (uppermost) Lago-Mare.

By contrast, some authors (i.e. Popescu et al., 2007, 2015; Do Couto et al., 2014; Clauzon et al., 2015; Pellen et al., 2017) described the nannofossil assemblage the Lago-Mare LM Unit in Malaga, the Zorreras Member in Sorbas, the uppermost di Tetto/Colombacci Fm. in some Apenninic localities and offshore in the Alborán Basin as having good preservation and showing no erratic fluctuations, all characteristics that led to their interpretation as autochthonous and to the conclusion that these sediments were deposited in a Mediterranean already replenished of Atlantic water (Fig. 4b). In addition, these authors reported the low abundance, but continuous presence of the biostratigraphic markers for the Zanclean *Triquetrorhabdulus rugosus* and *Ceratolithus acutus* (Fig. 9d) below the formally defined Miocene/Pliocene boundary (Van Couvering et al., 2000) in several onshore and offshore Mediterranean (as wells as Paratethyan) localities (see Popescu et al., 2017 for details and a
complete list of finding locations).

Such findings (especially that of *C. acutus*) are in sharp disagreement with most of the existing literature and have resulted in an important debate amongst the MSC community (e.g. Popescu et al., 2007, 2008 vs Roveri et al., 2008c and Stoica et al., 2016 vs Popescu et al., 2017), not only for their paleoenvironmental implications (i.e. presence of Atlantic water in the Mediterranean), but also for the chronostratigraphic repercussions (Fig. 4b). The chronostratigraphic value of *C. acutus* lies in its short temporal distribution straddling the M/P boundary (astrochronologically calibrated at 5.332 Ma; Van Couvering et al., 2000; Lourens et al., 2004). However, the corresponding biozone is established in oceanic areas (Zone CNPL1: 5.36-5.05 Ma; Backman et al., 2012; Agnini et al., 2017) and is considered not applicable to the Mediterranean region during the MSC due to the harsh physicochemical conditions that are unsuitable for marine biota (Di Stefano and Sturiale, 2010). The interpretation of these nannofossil assemblages in the westernmost areas of the Mediterranean has been countered with several observations: (1) the observation of these age-diagnostic taxa is often not replicated by other studies (e.g. Roveri et al., 2008a; Van Baak et al., 2015; Krijgsman et al., 2020b); (2) *Ceratolithus acutus* is very rare also in fully marine open-ocean sediments (e.g. Di Stefano and Sturiale, 2010); (3) despite being rare in the late Messinian Mediterranean, this species has never been documented together with other long-range taxa, generally predominant in the assemblage, in Stage 3 deposits (see discussion in Krijgsman et al., 2020b). Recently, Golovina et al. (2019) showed that the morphology and size of *C. acutus* overlaps with the shape and dimensions of destroyed ascidian spicules (i.e. calcareous elements produced by benthic tunicates; Fig. 9d), providing an explanation for erroneous identification of *C. acutus* in the Black Sea Basin (Golovina et al., 2019) and perhaps in the western Mediterranean Lago-Mare sediments as well.

## 5.6. Fish

Fossil fish remains provide information about salinity and depth and have been used to contradict the brackish nature of the Lago-Mare deposits by Carnevale et al. (2006a, 2006b, 2008, 2018) and Grunert et al. (2016). Euryhaline fish species inhabit marine to brackish environments and dominate settings with strong salinity variations while stenohaline fish have specific salinity requirements (marine, brackish, or freshwater) and cannot survive under different conditions. Demersal fish (i.e. those living in or immediately above the sea floor) have specific depth requirements, whereas pelagic fish occupy the water column within a wide range of depth variable from species to species. Fossil fish remains are found either as articulated or disarticulated skeletal parts, including teeth and otoliths, which are identified to the species level. Articulated fish skeletons typically indicate autochthonous deposition because of the difficulty in reworking and transporting intact skeletons. Otoliths and fish teeth are much more likely to be transported.

Otoliths and rare articulated skeletons (Fig. 9e) of marine and Paratethyan species have been reported from Stage 3 deposits, but commonly huge volumes of sediment are required to find even quite small numbers of these fossils (e.g. 20 tons from Monccucco, 6 tons from Cava Serredi, 700 kg from Capanne di Bronzo; Schwarzhans et al., 2020), much more than what is expected for normal marine deposits (i. e. < 30 kg; Agiadi et al., 2017; Karakitsios et al., 2017b).

Substage 3.1 sediments contain articulated skeletons (Fig. 9e) of the marine fish species *Lampanyctus licatae* and *Maurolicus muelleri*, and the shallow water, euryhaline species *Aphanius crassicaudus* in the Lower Feos Member in the Nijar Basin (de la Chapelle and Gaudant, 1987) and the marls of the first UG cycle in the Polemi Basin (Manzi et al., 2016a; Fig. 3a). Cava Serredi (Tuscany), Verduno and Monccucco (Piedmont) are the only other localities in which fish remains (only otoliths) in (claimed) substage 3.1 sediments are known (Carnevale et al., 2006a, 2008, 2018; Grunert et al., 2016).

The more diverse and abundant ichtyofaunal record occurs in

substage 3.2 in a few marginal sections on the Italian peninsula (Ciabot Cagna in the Piedmont Basin; Cava Serredi and Podere Torricella in Tuscany; Capanne di Bronzo, La Vicenne and Ca' Ciuccio in thrust-top basins of the Northern and Central Apennines). The Lago-Mare fish remains mainly comprise otoliths of both euryhaline and stenohaline taxa indicative of marine, brackish, and freshwater habitats (Carnevale et al., 2018). Three articulated skeletons of the euryhaline marine taxa *Mugil* cf. *cephalus* (Fig. 9e), the marine Indo-Pacific species *Spratelloides gracilis* and of *Gobius* sp. have been identified at Cava Serredi in a horizon < 1 m below the Miocene/Pliocene boundary (Carnevale et al., 2006b). The dominant stenohaline families in these assemblages are Gobiidae, a family of demersal fish occupying shallow-water marine, brackish and freshwater environments, and Myctophidae, which are marine mesopelagic fish that live below 200 m depth during the day, but feed at night in surface waters. A recent review of the Tortonian-Zanclean Gobiidae of the Mediterranean (Schwarzhans et al., 2020) showed that the otoliths of this family, described by Carnevale et al. (2006a, 2008, 2018) and Grunert et al. (2016) as belonging to marine Atlantic species, instead belong to brackish and freshwater species of Paratethyan affinity inhabiting sheltered prodelta environments. In fact, no normal marine demersal taxa were recognized in these assemblages by Schwarzhans et al. (2020). As for the Myctophidae, the vast majority of the taxa belonging to this family were recovered in Monccucco and Verduno from alluvial plain silty mudstones along with terrestrial mammals (Dela Pierre et al., 2011; Colombero et al., 2017 and references therein), pointing to a physically reworked origin. When $^{87}$Sr/$^{86}$Sr isotope ratios are measured (Carnevale et al., 2008; Grunert et al., 2016), the resulting Sr-based age of the otoliths is > 7 Ma, therefore further arguing against their in-situ origin. Since the good preservation of the otoliths suggests they did not suffer physical reworking (Carnevale et al., 2006a, 2006b, 2008, 2018; Grunert et al., 2016), predators foraging in open marine settings and migrating to marginal environments are proposed as a way out of the enigma (Carnevale et al., 2008, 2018; Grunert et al., 2016; Colombero et al., 2017). However, Carnevale et al. (2006a) also rule out that so well preserved otoliths may have suffered post-mortem transport and action of the digestive acids in the stomach of predators. Rare findings of Myctophidae from Ciabot Cagna (3 species), Cava Serredi (1 species), Capanne di Bronzo (1 species) and Podere Torricella (6 species) (Carnevale et al., 2018) are all from sections where the host sediments have not been studied in sufficient detail to be clear about the *in situ* or reworked nature of the fossil assemblage. This lack of sedimentological uncertainty also extends to the stratigraphic position of many samples, because a stratigraphic log is provided for only a few sections (i.e. Ca' Ciuccio, Cava Serredi and Monccucco; Carnevale et al., 2006a, 2006b). What this stratigraphic information suggests is that euryhaline fish taxa are widespread throughout substage 3.2, whereas strictly Myctophidae, which are an oceanic, marine stenohaline species, only occur very close to the base of the Pliocene, plausibly corresponding to the uppermost lithological cycle in substage 3.2 (~5.35-5.33 Ma; Carnevale et al., 2018).

## 5.7. Summary of the Stage 3 paleontological record

The aquatic fossil record of MSC Stage 3 indicates that substage 3.1 in onshore sections is mostly barren, while diverse assemblages characterize substage 3.2 deposits. By contrast, the deep record as a whole contains relatively few, low diversity assemblages. This might be as a consequence either of the limited sample locations recovered from the offshore areas (see Fig. 2b) or because the environmental conditions in the intermediate-deep basins were less favorable for sustaining the life forms typical of the onshore domain. Nevertheless, the assemblages that are found in both marginal and deep locations comprise mixed brackish and marine species.

Brackish species are mostly represented by ostracods and dinocysts (and mollusks here not addressed because poorly studied; see Esu, 2007 and Guerra-Merchán et al., 2010). Prominent is the affinity of these late
<figure><img src="image_23.png" /></figure>

(caption on next page)
**Fig. 10.** Isotopic record of MSC Stage 3 for the Mediterranean Basin. (a) Compilation of MSC Stage 3 $^{87}$Sr/$^{86}$Sr isotope data sourced from ostracod valves and gypsum crystals (see Supplementary material 1 and subsection 6.1 for references). Data are plotted with the global $^{87}$Sr/$^{86}$Sr seawater curve (McArthur et al., 2012). Error bars indicate analytical error, which is so small in some cases that no error bars are visible at this scale. To not complicate the figure, horizontal error bars have not been added for the sections/cores unprovided of a chronostratigraphic framework and for which age uncertainties are present (i.e. all but Nijar and Vera basins, Eraclea Minoa and onshore Cyprus; see Fig. 3). Note that none of the $^{87}$Sr/$^{86}$Sr isotope ratios but one from Nijar plot on the ocean curve. In the inset is shown the Mediterranean Sr record for the entire MSC as well as the time-equivalent Eastern Paratethys record (modified after Andreetto et al., 2021). (b) Plot of $\delta^{34}\text{SO}_4$ and $\delta^{18}\text{O}\text{SO}_4$ in Stage 3 gypsum and anhydrite beds from onshore and offshore localities (see Supplementary material 1 and subsection 6.2 for references). No measures are available from the marginal basins, where gypsum did not deposit during Stage 3. The dark blue and black rectangles represent the sulfate isotopic composition of the Global Messinian ocean and Stage 1 (PLG) evaporites, respectively. The light blue area represents the sulfate isotopic composition of mixtures of Messinian marine waters with non-marine sources. The red area represents the isotopic composition of the residual sulfate ion in a basin where marine Messinian sulfate is consumed by microbial sulfate reduction to produce $H_2S$. The arrow represents the isotope trajectory of dissolved sulfate resulting from the mixing of residual $^{34}$S-enriched sulfate produced by MSR and $^{34}$S-depleted sulfate produced by $H_2S$ oxidation. All the published $\delta^{34}\text{SO}_4$ and $\delta^{18}\text{O}\text{SO}_4$ values are provided corrected with the fractionation factors $\delta^{34}\text{S}=+1.65\%$ and $\delta^{18}\text{O}=+3.5\%$ to smooth the isotopic fractionation effects experienced by dissolved sulfate and to reason on values reproducing the isotopic composition at the time of gypsum precipitation. (c) $\delta\text{D}$ isotopes of C$_{29}$ and C$_{31}$n-alkanes and C$_{37}$ and C$_{38}$ long chain alkenones recorded in the Stage 3 gypsums and marls of the Eraclea Minoa section (modified from Vasiliev et al., 2017). Blue lines indicate the values recorded in the present day lacustrine settings for the n-alkanes (Sachse et al., 2006) and in the alkenones from the Mediterranean in the recent times (Van der Meer et al., 2007). Error bars indicate standard errors of the mean.

Messinian Mediterranean brackish species with the same species that were simultaneously dwelling in the Eastern Paratethyan basins (Dacian, Euxinic and Caspian) and in the North Aegean. Since these organisms were not present in the Mediterranean at any time before the MSC, they are considered, with a broad consensus, as *in situ*. This conclusion is further corroborated by the mixing of adult and juvenile forms in the ostracod assemblages and by the good preservation of the specimens, which do not show typical evidence of physical reworking like abrasion, dissolution, or fragmentation. Still problematic is the time of their arrival in the Mediterranean and their likelihood as biostratigraphic tool. From our review it seems more likely that truly Paratethyan species of ostracods entered the Mediterranean already during substage 3.1, when they colonized intermediate-deep settings, while they entered the marginal basins at different times during substage 3.2. As for dinocysts, characteristic is their occurrence only in the uppermost Messinian. However, it must be noted that samples from the substage 3.1 interval are rarely processed for dinocysts, especially in age model-equipped sections (Fig. 3a). The route followed by the Paratethyan immigrants is equally contested and important for paleoenvironmental and paleohydrological interpretations. In view of a desiccated Mediterranean, their migration can only have taken place passively by means of aquatic migratory birds. Conversely, the homogeneity of the ostracod assemblages throughout the Mediterranean marginal basins is more indicative of the presence of a water body fed by Eastern Paratethys and connecting all Mediterranean subbasins, therefore implying relatively high water-level conditions (at least at times when ostracod-bearing sediments deposited; see Andreetto et al., 2021).

Marine assemblages are composed by foraminifera, nannofossils, dinocysts and calcareous nannofossils. Their reworked or *in situ* nature is in many cases contested but critical for paleoenvironmental interpretation. The picture that emerges from our review is that an open marine signature is questionable in the foraminifera, nannofossils, dinocyst and fish records, as well as in other biotic groups (e.g. corals, echinoids and mammals) here not tackled (and for which we refer the reader to Dominici et al., 2018 and Carnevale et al., 2019). All marine representatives of the above mentioned categories were reintroduced into the Mediterranean only at the beginning of the Pliocene and at the expense of the Paratethys species that, instead, disappeared. Collectively, these observations lead us to conclude that the marine model as conceived by Carnevale et al. (2006a, 2006b, 2008, 2018) and Grunert et al. (2016) has no foundation and therefore will not be further discussed.

## 6. The geochemical perspective

Variations in the water sources draining into the Mediterranean are expected to be reflected also in (geo)chemical properties of the paleo-depositional environments. Important information about the nature of the connectivity framework of the Mediterranean can be gained by

interpreting geochemical signals that respond to the presence or absence of an exchange with a chemically-unique water body. Four main geochemical proxies have been applied so far to MSC Stage 3 sedimentary and paleontological records. These includes both radiogenic (Sr isotope ratios) and stable isotopes (sulfate and oxygen) measured on fossils and minerals and hydrogen isotopes on molecular biomarkers. This section summarizes the dataset available for geochemical proxies (Fig. 10; Supplementary material 1) and its interpretation(s) for MSC Stage 3.

### 6.1. Strontium isotope ratios ($^{87}$Sr/$^{86}$Sr)

The available strontium isotope data for Stage 3 (Fig. 10a; Supplementary material 1) derive from measurements on both Ca-bearing fossils (ostracod valves, mollusk shells, fish otoliths; Fig. 9a) and minerals (calcite and gypsum), where $Sr^{2+}$ dissolved in an aqueous solution substitutes Ca atoms due to their similar ionic radius (e.g. Hajj et al., 2017). Here we screen the available dataset and discuss only results that (1) reflect the original primary isotopic signal, i.e. the isotopic signal of the fluid at time of shell calcification or mineral precipitation, and (2) for which timing of mineral precipitation can be constrained. This screening excludes bulk carbonate samples (e.g. *Colombacci* limestones; Bassetti et al., 2004), which contain carbonate compounds of various and/or unknown provenance, measurements from mollusk shells and otoliths (e.g. Carnevale et al., 2008; Grunert et al., 2016; Roveri et al., 2019a), because they are made of mineral phases easily altered during diagenesis (e.g. aragonite; Marcano et al., 2015), and data coming from reworked material (e.g. all reworked gypsum or transported foraminifera). $^{87}$Sr/$^{86}$Sr isotope ratios have also been measured by Müller and Mueller (1991) and Roveri et al. (2014b) on the halite beds recovered at Sites 134, 374 and 376 (Ryan et al., 1973; Hsü et al., 1978b). Although they provide interesting interpretative aspects, we do not consider these Sr measurements because the position of Sr in the crystal lattice of halite is unknown and the removal of all contaminants, that is not a straightforward procedure (see Meillijson et al., 2019), is not clear it was achieved by Müller and Mueller (1991) and Roveri et al. (2014b). As a matter of fact, there is no consistency between data generated from roughly the same interval in Core 134 by Müller and Mueller (1991) (0.708968) and Roveri et al. (2014b) (0.708800-0.708896). Added to this is the uncertainty over the provenance of halite in Sites 134 and 374 (see subsections 4.2 and 4.6.1), which violates both criteria mentioned above.

The general trend of the Mediterranean $^{87}$Sr/$^{86}$Sr isotope ratio during the MSC deviates from the ocean curve towards the less radiogenic values of the major peri-Mediterranean rivers and Paratethys and returns abruptly to oceanic values at the Miocene/Pliocene boundary (Fig. 10a inset). This trend is regarded to reflect the progressive restriction of Mediterranean-Atlantic exchange and the relative increase in