Brunhes boundary: les liaisons dangereuses?

Quaternary Science Reviews 226 (2019) 106039

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A Mediterranean perspective on 10Be, sedimentation and climate around the Matuyama/Brunhes boundary: les liaisons dangereuses? Luca Capraro a, *, Fabio Tateo b, Patrizia Ferretti c, Eliana Fornaciari a, Patrizia Macrì d, Daniele Scarponi e, Nereo Preto a, Feng Xian f, g, Xianghui Kong f, g, Xingjun Xie f, g
degli Studi di Padova, Via G. Gradenigo 6, 35131 Padova, Italy Dipartimento di Geoscienze, Universita CNReIGG, Via G. Gradenigo 6, 35131 Padova, Italy c
Ca
Foscari di Venezia, Via Torino 155, 30172 Venezia, Italy Dipartimento di Scienze Ambientali, Informatica e Statistica, Universita d Istituto Nazionale di Geofisica e Vulcanologia, Via di Vigna Murata 605, Ie00143 Roma, Italy e
di Bologna, via Zamboni 63e67, 40126 Bologna, Italy Dipartimento di Scienze Biologiche, Geologiche e Ambientali, Universita f State Key Laboratory of Loess and Quaternary Geology, Institute of Earth Environment, CAS, 710061 Xi’an, China g Shaanxi Key Laboratory of Accelerator Mass Spectrometry and Application, Xi’an AMS Center, Xi’an 710061, China a

b

a r t i c l e i n f o

a b s t r a c t

Article history: Received 27 June 2019 Received in revised form 21 October 2019 Accepted 23 October 2019 Available online xxx

The 10Be/9Be ratio is commonly employed as a tool for establishing the stratigraphic position of paleomagnetic excursions and reversals whenever the traditional paleomagnetic approach fails to provide conclusive results. In particular, it is held that 10Be production rates in the atmosphere depend on the strength of the Earth’s magnetic field, and the fallout and deposition of cosmogenic beryllium at the surface happen on a very short time scale. However, investigations performed on terrestrial and marine successions demonstrate that the 10Be record and the paleomagnetic signal are often asynchronous. Mechanisms that control the conveyance and deposition of cosmogenic 10Be to the seafloor are still ambiguous and poorly documented. Here, we discuss the dynamics of 10Be in a central Mediterranean marginal marine depositional scenario characterized by a pervasive terrigenous influx. Our data show that a very close correlation exists between 10Be concentrations and the local proxy of rainfall rates and regimes (pollen), indicating that a considerable 10Be transport from the mainland may occur in response to the remobilization of terrestrial reservoirs during periods of increased runoff. Superimposed is a dynamic oceanographic setting that further controls the preservation potential of 10Be at the bottom, in terms of changing water chemistry and/or composition of the sedimentary flux to the seafloor. Results of our investigation suggest that, in particular environmental and depositional settings, the interplay between climate, terrigenous yield and oceanography may jeopardize the sedimentary depiction of the meteoric 10Be contribution, thus challenging the use of 10Be for tracking the stratigraphic position of geomagnetic reversals. © 2019 Elsevier Ltd. All rights reserved.

1. Introduction Production of cosmogenic 10Be in the atmosphere (also known as ‘meteoric’ 10Be; McHargue and Damon, 1991) is caused by the cosmic ray spallation of nitrogen and oxygen (e.g. Kovaltsov and Usoskin, 2010). Intensity of this phenomenon varies at different time scales according to several factors, including the longeterm variability in the strength of the Earth’s magnetic field, with production rates of meteoric 10Be that are inversely proportional to the

* Corresponding author. E-mail address: [email protected] (L. Capraro). https://doi.org/10.1016/j.quascirev.2019.106039 0277-3791/© 2019 Elsevier Ltd. All rights reserved.

intensity of the geomagnetic dipole moment (e.g., Elsasser et al., 1956; Frank et al., 1997; Beer et al., 2012). The available datasets indicate that geomagnetic field reversals and excursions were characterized by periods of weak geomagnetic dipole moment and increased 10Be production in the atmosphere, as especially well documented at the MatuyamaeBrunhes boundary (e.g., Raisbeck et al., 1985, 2006; Carcaillet et al., 2003; Zhou et al., 2014; nabre az et al., 2014; Valet et al., 2014; Simon et al., 2016c, Me 2018a,Simon et al., 2018b; Du et al., 2018) and at the Laschamp nabre az et al., 2012; Raisbeck event (e.g. Thouveny et al., 2008; Me et al., 2017). Accordingly, authigenic 10Be measurements from sedimentary successions are employed as an alternative and/or

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support to the traditional paleomagnetic investigations, when tracking the stratigraphic position of prominent geomagnetic events in critical conditions, such as the study of ice cores (e.g., Raisbeck et al., 1985, 2006), or whenever conventional paleomagnetic analyses deliver ambiguous results (e.g., Zhou et al., 2014; Simon et al., 2016a). However, 10Be scavenging from the atmosphere follows complex pathways that may affect the mode and tempo of the cosmogenic fallout (e.g., Field et al., 2006; Monaghan et al., 1986; Willenbring and von Blanckenburg, 2010a; Raisbeck et al., 2006). In the case of marine sedimentary environments, several additional variables may also contribute to the 10Be conveyance and deposition, ranging from ocean circulation patterns to the local depositional style and setting (e.g., Mangini et al.,
s et al., 1989; Kusakabe et al., 1990, 1991; Brown et al., 1984; Bourle 1992; Measures et al., 1996; von Blanckenburg et al., 1996; Willenbring and Von Blanckenburg, 2010b; von Blanckenburg and Bouchez, 2014; Simon et al., 2016b; Capraro et al., 2018). Dynamics within the catchment area, such as the episodic mobilization of continental reservoirs, can promote massive10Be fluxes to the basins, where 10Beerich particles may be involved in intricate depo
s et al., sitional processes before and after their settling (e.g. Bourle 1989; von Blanckenburg et al., 1996, 2012; Frank et al., 2009). Altogether, the magnitude of nonemeteoric 10Be yields to marginal marine settings may be considerably larger than the pure atmospheric input, and might thus hamper a realistic depiction of the “true” cosmogenic signal (e.g., Brown et al., 1988; You et al., 1988; von Blanckenburg and Bouchez, 2014). Documentation on the correspondence between the paleomagnetic and 10Be signals in marginal (hemipelagic) settings is anything but comprehensive, because there is a dearth of marine

stratigraphic successions with high sedimentation rates preserving a welleconstrained record of both 10Be and the Matuyama-Brunhes reversal. Situation is especially critical for the Mediterranean region, where the marine paleomagnetic record of the last reversal is generally poorly preserved. To our knowledge, the Valle di Manche (VdM) section (Crotone Basin, Calabria, Southern Italy; Roda, 1964; Rio et al., 1996; Capraro et al., 2017) hosts the only available and reliable record of the Matuyama-Brunhes boundary for the whole Mediterranean marine stratigraphy, which was recognized by means of paleomagnetic investigations further supported by accurate magnetic mineralogy analyses (Macrì et al., 2018). Therefore, the VdM section offers an unprecedented benchmark for testing whether the 10Be record reconstructed in a landlocked basin with hemipelagic sedimentation should be employed for pinpointing the Matuyama-Brunhes geomagnetic boundary securely and independently from the environmental conditions at the time of deposition. 2. Geologic setting The Valle di Manche (VdM) section is located in the inner part of the Crotone sedimentary basin (Calabria, Southern Italy; Fig. 1), generally interpreted as a forearc basin located above the external submarine part of the Calabrian accretionary wedge (Rossi and Sartori, 1981). The Crotone Basin is bounded north and south by two NWeSE leftelateral strikeeslip faults, namely the RossanoeSan Nicola and PetiliaeSosti, respectively (Fig. 1; Van Dijk and Okkes, 1991). Landward, it is confined by the Sila plateau, mainly constituted by Paleozoic granites (e.g., Ogniben, 1973; Messina et al., 1994). The sedimentary infill of the CB is represented

Fig. 1. Location of the studied section. a): location of the Crotone Basin in Southern Italy (orange square); the green triangle indicates the position of the Montalbano Jonico (MJ) section. b): simplified geologic map of the Crotone Basin, indicating the location of the San Mauro sub-basin (orange square). Borders of the San Mauro syncline are marked by the barbed lines. The blue dot marks the position of the Vrica section (see text). c): panorama view of the Valle di Manche badlands, exposed below the cemetery of San Mauro Marchesato (modified from Capraro et al., 2017). The Cutro 2, SM1, SM2 and SM3 units are indicated. The dashed line in SM2 indicates the position of the “Pitagora ash”. The dashed line above the Cutro 2 unit indicates the inferred position of the SM1 bottom, which is not exposed here. The double vertical line marks the sampling profile. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

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by a very thick succession of Miocene to Pleistocene sediments (Ogniben, 1955; Roda, 1964; Di Grande, 1967; Massari et al., 2002, 2010), which were accumulated into smaller subebasins controlled by differential subsidence rates (Capraro et al., 2006, 2011; Massari et al., 1999, 2002, 2010). The younger part of the Crotone Basin stratigraphy, which contains the interval studied for this work, lies on a thick and widely exposed Lower Pleistocene succession, consisting of open marine muds (the “Cutro 1” of Rio et al., 1996) that grade into the cyclothemicallyeorganized “Cutro 2” unit (Rio et al., 1996). In a large part of the CB, the “Cutro” muds are truncated and directly overlain by marine and/or continental terrace deposits in response to the early uplift of the basin (Capraro et al., 2011, 2017). Only in confined sectors of the Crotone Basin, such as the San Mauro subebasin (Rio et al., 1996; Massari et al., 1999, 2002), the local persistence of strong subsidence rates during middle Pleistocene times allowed for the deposition of an expanded openemarine stratigraphy above the “Cutro” muds. The San Mauro sub-basin consists of a synsedimentary growth syncline bounded by two dextral obliqueeslip faults (Fig. 1; Massari et al., 2002, 2010; Macrì et al., 2014). It hosts a cyclothemic, shallowingeupward marine to continental succession (the “San Mauro Sandstone” of Rio et al., 1996), which we will refer to hereby as the “San Mauro Formation”, organized into a stack of cyclothems that reflect the local sedimentary response to changes in both glacioeustasy, tectonics and climate (Rio et al., 1996; Massari et al., 2002, 2007; Capraro et al., 2005, 2017). Rio et al. (1996) subdivided the San Mauro Formation into three lithological units (SM1 to SM3; Fig. 2). Bottom up, these are the coarseegrained SM1, represented by shallowewater sands deposited during the glaciation associated to the Marine Isotope Stage (MIS) 24/22; the mudedominated SM2, spanning from MIS 21 to MIS 18; the coarseegrained SM3, which incorporate shallowemarine deposits such as marine and continental sands and gravels. The soft muds correlative to the MIS 19 interglacial in SM2 contain a prominent pyroclastic layer, known as the “Pitagora ash” (Rio et al., 1996). The material studied for this work was recovered from the Valle di Manche (VdM) section (Fig. 1), which represents the best exposed and more expanded profile available of the San Mauro Formation (Rio et al., 1996; Capraro et al., 2005, 2015, 2017). The section was laid in a depocentral setting of the San Mauro subbasin, where the stratigraphic record is basinal, continuous and complete (Massari et al., 2002, 2007). Indeed, the stratigraphic log shows a closely spaced alternation of facies (Fig. 2), however all changes are subtle and gradual, and no evidence occurs of erosional surfaces, hiatuses, and/or abrupt variations in sediment composition, being the sedimentation style persistently dominated by hemipelagic mud settling within the interval of interest (Massari et al., 2002, 2007; Scarponi et al., 2014). From a physical stratigraphic perspective (Fig. 2), the studied interval includes the entire full MIS 19 interglacial, spanning from the top of sandy facies E (
3.25 m level) to the basal part of the fossiliferous silts of facies C1 (þ3.57 m). With respect to the subdivision presented in Capraro et al. (2017), to which we refer for further information, we have singled out a new unit (Facies A0) that marks the gradual, yet rapid, transition from the fossiliferous sands of Facies E to the dark muds of Facies A1 (Fig. 2). The Matuyama-Brunhes boundary is documented immediately above the “Pitagora ash” (Macrì et al., 2018). For paleomagnetic analyses, ca. 100 in situ-oriented samples were collected in the stratigraphic interval between
24.80 m and þ6.16 m, and characteristic remanent magnetization (ChRM) directions were calculated from the stepwise AF demagnetization diagrams (Macrì et al., 2018). Computed ChRM directions and VGP latitudes indicate that the M-B magnetic reversal develops over a 3-cm thick interval centered at ca. þ12.5 cm. Two short intervals of normal polarity

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were detected below the “Pitagora ash” between
3 and
20 cm and between
105 and
110 cm, respectively (Fig. 3). The DGRM/ DNRM ratio, a sensitive proxy of the occurrence of ferrimagnetic greigite, shows that a spurious gyromagnetic remanent magnetization (GRM) effect occurs variably only in the lowermost part of the studied interval, below
26 cm, with a significant increase between
105 cm and
115 cm (Fig. 3). A number of rock magnetic analyses on representative discrete samples (Macrì et al., 2018) confirm that magnetic properties of the sediments in the interval straddling the Matuyama-Brunhes geomagnetic transition are very good. 3. Materials and methods We have measured the total organic carbon content (TOC) and mineralogical composition of a batch of samples previously employed for reconstructing the published d18O, pollen, calcareous nannofossils (Fig. 4) and ostracods records (Capraro et al., 2015, 2017; Rossi et al., 2018; Azzarone et al., 2018). In addition, we analyzed the mineralogy of an extra sample picked from the “Pitagora ash”. Based on the age model of Capraro et al. (2017), this dataset (25 cm resolution) allows for an average time resolution in the order of ca. 0.8 kyr. For measuring beryllium concentrations and the stable isotope composition of Melonis barleeanum, we have processed a different sample set collected with a ca. 5 cm resolution along the VdM profile in the interval between
2.54 m and 3.75 m, which encompasses the entire MIS 19c interglacial (Capraro et al., 2017). Since we analyzed one sample out of two, this data series offers an average time resolution in the order of 0.35 kyr. 3.1. Stable isotopes We measured the stable isotope composition (O and C) of the benthic foraminifer M. barleeanum from 54 samples, which were prepared by washing ca. 100 g of sediment on a 63emm sieve using distilled water and ovenedrying overnight the residue on the sieves at 50  C. Between 10 and 35 pristine foraminifer specimens were handpicked from the fraction >150 mm under a binocular microscope and processed for each isotopic analysis. Tests were crushed gently using two clean glass plates, transferred to the massespectrometer vials, and soaked in 3% hydrogen peroxide in order to remove any possible organic contaminant. Analytical grade acetone was added and samples were cleaned ultrasonically, after which the excess liquid was siphoned off. Finally, samples were ovenedried overnight at 50  C. Carbon dioxide for isotopic analysis was released using orthophosphoric acid at 70  C in an automated continuous flow carbonate preparation GasBench II device, and analyzed in a Thermo Scientific Delta V Advantage Isotope Ratio Mass Spectrometer at the Department of Geosciences of the University of Padova. The laboratory standard (Carrara marble MAQ1) was calibrated to VPDB through NBS 19, using the values of
2.20‰ for the d18O and þ1.95‰ for the d13C, as recommended by Coplen (1988, 1995). The longeterm analytical reproducibility was ±0.07% for the d13C and ±0.09% for the d18O (1s), and was evaluated by systematically analyzing a quality control standard (GR1 marble) along with the samples. 3.2. Mineralogy The mineralogical composition of sediments was measured for 30 samples, which were cleaned, ovenedried and then crushed and pulverized by hand. Analyses were performed via Xeray diffraction by a Philips X’Pert Pro diffractometer (Co tube and secondary monochromator) at the Department of Geosciences, University of

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Fig. 2. Columnar log of the VdM section. From left to right: stratigraphic thickness with respect to the base of the “Pitagora ash”; facies labels; lithology; facies description and interpretation of the associated depositional setting.

Padova. Powders were backloaded in order to minimize the chances of preferred orientations. The mineral constituents were computed on a semiequantitative basis according to the soecalled   et al., 2001; Chipera and Bish, 2002) fullepattern approach (Srodo n and using the mineral database by Eberl (2003, revised on 2011). 3.3. Total organic carbon The Total Organic Carbon (TOC) of 15 bulk sediment samples was measured at the Department of Geosciences, University of Padova. Ca. 5 mg of sample were weighted in silver capsules and etched with 10% HCl overnight on a hot plate at 65  C, repeatedly, until the reaction was complete. The dry capsules were then wrapped and analyzed with a Thermo Scientific Flash 2000 Elemental Analyzer. Calibration was made using a pure carbonate standard and blank capsules, these latter treated identically to the samples. All samples were run in duplicate. Results are given in wt %. 3.4. Beryllium For reconstructing the records of “authigenic” 10Be and 9Be, we analyzed 68 sediment samples that were prepared at the Xi’an AMS center of the Institute of Earth Environment CAS (IEECAS) and the Xi’an Jiaotong University, according to the chemical procedure described in Zhou et al. (2007) with minor modifications accounting for the difference between loess and marine sediments. “Authigenic” 10Be measurements were performed using the 3eMV

accelerator mass spectrometer (AMS) at the Xi’an AMS Center, IEECAS. This facility is a multieelement system (14C, 10Be, 26Al, 129I) with a single beam line dedicated to AMS applications. For the current analyses, chemical procedure blanks (for quality control) were ca. 1  10
14, in keeping with optimal routine measurements. 9 Be measurements were performed using the Inductively Coupled Plasma Atomic Emission Spectrometry at the Xi’an AMS Center, IEECAS. The measured uncertainties of 10Be data are less than 2.5% (average ca. 1.4%), while the uncertainties of 9Be data are less than 5.4% (average ca. 2.0%). 3.5. Curve fitting For data handling, we employed the Solver© addeon in Microsoft Excel©, a tool primarily conceived for data mining and problem solving that also allows for an effortless estimate of the coherence and similarity between different data series. In order to avoid any possible bias and distortion of the original data sets, we decided to employ a very simple analytical approach, which was basically aimed at finding the best fit between a given set of measured values (“proxy” series) and a target record. The Solver© tool generates a “calculated” curve (calc) by multiplying all the data in the proxy series (P) by a given coefficient (r) and shifting the result horizontally by an arbitrary value (k), as follows:

calc ¼ ½P * r þ k The calculated series (calc) is tentatively compared to the target, and the routine is reiterated automatically with changing (r) and (k)

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Fig. 3. Computed paleomagnetic VGP latitudes correlated with the d18O of benthic foraminifer curve Uvigerina peregrina and the rock magnetic DGRM/DNRM parameter, as sensitive proxy of the occurrence of ferrimagnetic greigite along the stratigraphic section (see Macrì et al., 2018, for details). The high-resolution VGP paths indicate that the Matuyama-Brunhes transition record is extremely sharp, as it occurs across a thin stratigraphic interval (ca. 3 cm) above the base of the “Pitagora ash” where the magnetic mineralogy is remarkably homogeneous.

Fig. 4. Left panel: stratigraphic log and records of d18O and d13C reconstructed in the VdM section for Uvigerina peregrina and Melonis barleeanum. The d18O record of U. peregrina is redrawn from Capraro et al. (2017). Middle and right panels: a) abundances of Gephyrocapsa spp., given as number of specimens counted within 1 mm2 of the analyzed slides. This genus was selected as it represents the most common taxon within the nannofossils assemblage, and roughly approximates the total abundance of native nannofossils in the investigated samples; b) percentile abundances of grains belonging to the “mountain” conifer group within the total pollen assemblage; data are redrawn from Capraro et al. (2017). The thin dashed lines mark the facies boundaries; the red dashed line indicates the stratigraphic position of the “Pitagora ash”. All curves are plotted against depth. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

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until the best fit between the target and calculated values is accomplished. This result is marked by the smallest possible standard deviation (SD) between the records, calculated as follows:

vffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi u n u 1 X SD ¼ t ðxo
xc Þ2 n
1 1 where n is the number of data points considered for calculation, x0 is the measured value and xc is the calculated value. Computations made separately on individual proxy series could be merged arithmetically into a single calculated curve. If the measured and calculated series contain a different number of values e as in the present case, given the different sampling resolutions e missing data points are generated automatically by linear interpolation. 4. Results 4.1. Stable isotopes Melonis barleeanum is a deepeinfaunal foraminifer species that essentially provides information on the chemistry of interstitial waters (e.g., Jorissen, 1987). Still, d18O values measured on this species, oscillating between ca. 0.70‰ and 1.69‰ vs. VPDB, confirm both the shape and stratigraphic position of the MIS 19 interglacial as previously documented by the shalloweinfaunal benthic foraminifer Uvigerina peregrina (Fig. 4; Capraro et al., 2017). The d13C record of M. barleeanum oscillates between ca. 0 and
1.80‰ vs. VPDB, in phase with that of U. peregrina (Fig. 4). This wide range confirms that profound oceanographic perturbations took place within the San Mauro sub-basin across MIS 19, especially in terms of basin circulation and primary and export productivity rates and preservation (e.g., Rohling et al., 2004, 2015; Grimm et al., 2015; Grant et al., 2016). Extreme negative values, indicative of increasing flows of isotopically “light” organic matter to the seafloor (e.g., Duplessy et al., 1984; Curry and Oppo, 2005), are observed between ca.
2 m and ca.
0.70 m (Facies A0eA1; Figs. 2 and 4). From this interval upwards, d13C values increase gradually and attain a maximum at ca. 2.90 m, in the upper part of Facies B, where conditions of maximum local sea level depth and minimal terrigenous input are documented (Capraro et al., 2017; Azzarone et al., 2018). Paleontological and physical stratigraphic data collected from this interval (Capraro et al., 2015, 2017; Scarponi et al., 2014; Rossi et al., 2018) point to an optimal oxygenation at the seafloor, with thorough exploitation of the available organic matter by benthic communities. 4.2. Mineralogy The mineralogical composition of sediments does not change significantly across the studied interval (Fig. 5). We only detected limited variations in the relative abundances of the principal mineral components that, however, are consistently present in all the investigated samples, with the exception of that picked from the “Pitagora ash” (Fig. 5). The mineral assemblage is persistently dominated by quartz and phyllosilicates, among which disordered mixed layers (MeL) and micaceous phases (i.e., illite and true micas) represent the most abundant populations. Less abundant are calcite and feldspars, while chlorite and kaolinite only occur with small percentages. Pyrite and gypsum are present with noticeable, yet very small, abundances only in the interval corresponding to facies A1 (between
2 and
0.5 m). Notably, their highest concentrations occur in the lower part of facies A1 (from
2 to
1.25 m), where calcite attains to its minimal abundances

Fig. 5. Relative abundances of the minerals recognized in the VdM section, plotted against depth. All values are given as wt%. Far right: estimate of the average sediment grain size (see text for details). Horizontal dashed lines as in the previous figure.

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(Fig. 5). The coexistence of sulphates and sulfides at the same stratigraphic level suggests that gypsum occurs as a weathering product of pyrite, which is common in the interval of Facies A0eA1 (Capraro et al., 2017; Rossi et al., 2018; Scarponi et al., 2014). These evidences point to an acidic, sulfideerich environment at the seafloor, probably sustained by the release of hydrogen sulfide (H2S) from chemoebacterial communities (e.g., Calvert, 1983; Mercone et al., 2001). In presence of sulfidic acid, either in equilibrium with pyrite and/or as the product of late oxidation of pyrite, fineegrained carbonates were partly subjected to dissolution, as demonstrated by the bad preservation of calcareous shells and tests in Facies A0 and A1 (Capraro et al., 2015, 2017; Scarponi et al., 2014). We employed the [(quartz þ plagioclase)/(MeL þ kaolinite)] ratio as a rough proxy of the average sediment grain size (Fig. 5). Specifically, quartz and plagioclase are representative of the coarser fraction, whereas disordered clays minerals (MeL) and kaolinite distinguish the fineegrained fraction (Chamley, 1989; Dinelli et al., 2007; Feng et al., 2011). As expected, higher values are found in correspondence to facies E, which is constituted by fineegrained siliciclastic sands. Beginning from the very base of facies A0 (
2.75 m), grain size decreases gradually down to a minimum in the middle part of Facies A1 (
1 m). An upwardecoarsening trend leads to the base of the grey muds of Facies A2, where grain size values attain a lesser yet distinctive peak (Fig. 5). The base of the “Pitagora ash” is marked by an abrupt shift to extremely low values that persist up to þ30 cm, suggesting the presence of abundant fineegrained pyroclastic material dispersed within the hemipelagic fraction. From this point upwards, a long interval is observed of slowly increasing grain size, which terminates in correspondence to the very fine silts of Facies C1.

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4.3. TOC At VdM, Facies A0 and A1 (Fig. 2) are characterized by an overall dark color of sediments, the common occurrence of large sulfide nodules and a depleted and poorly preserved benthic fossil assemblage, suggesting low oxygenation rates and high fluxes of organic matter to the seafloor (Scarponi et al., 2014; Rossi et al., 2018). Simon et al., 2016a found similar conditions in the MJ section at a matching stratigraphic position, i.e., immediately below the isotopic “warm” peak of MIS 19. These physical stratigraphic evidences are consistent with those generally associated to the periodic deposition of Mediterranean PrecessioneRelated Sapropel layers (MPRS; e.g., Rossignol-Strick, 1985; Lourens, 2004; Rohling et al., 2015; Grant et al., 2016; Andersen et al., 2018). In order to further substantiate this correlation, we measured the total organic carbon (TOC) content for 15 sediment samples in the interval between
3.35 m and 2.48 m. Values range from ca. 0.3% to 0.6% of the total sample weight, with maximum concentrations that, as expected, are found in the interval belonging to Facies A1 (Figs. 2 and 6). A very good match exists between the local d13C and TOC records (Fig. 6), suggesting that the amount of organic carbon measured at VdM reflects the primary biological productivity in the basin. We conclude that intervals of higher TOC concentrations at VdM point to a period of increased surface productivity and/or augmented preservation potential of the organic snow at the basin bottom, in agreement with what expected at the deposition of a MPRS (e.g., Cita and Grignani, 1982; Rohling et al., 2015). The presence of an impoverished benthic fossil fauna indicates however that the oxygen depletion at the seafloor was not as severe as that documented in the case of “true” MPRS (e.g., Van Straaten, 1972; Capotondi et al., 2006; Abu-Zied et al., 2008). Still, TOC concentrations measured in Facies A1 (ca. 0.6%) remain

Fig. 6. Total organic carbon (TOC) values (reverse values) compared to the d13C record of Uvigerina peregrina and Melonis barleeanum; values are plotted against depth. Horizontal dashed lines as in the previous figures.

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well below the threshold conventionally associated to sapropel layers (TOC  2%; Olausson, 1961; Kidd, 1978). In the Vrica section, likewise located within the Crotone Basin, (e.g., Pasini and Colalongo, 1982; Hilgen, 1991), TOC concentrations range from ca. 0.7e0.9%, in the exposed laminitic, MPRSeequivalent layers to 0.2e0.3% in the intervening clays (Raffi and Thunell, 1996). This range compares well to that measured in the VdM section (Fig. 6). It follows that Facies A1 is fully consistent with a MPRSeequivalent layer subjected to subaerial exposure and weathering, in keeping with our physical stratigraphic and paleontological interpretations (Scarponi et al., 2014; Capraro et al., 2017; Rossi et al., 2018). Core records recovered from the eastern Mediterranean, however, show no lithological evidence of MPRS layers in the interval between MIS 21 and MIS 18 (e.g., Castradori, 1993; Lourens, 2004; Konijnendijk et al., 2014, 2015). One may speculate that the development of hypoxic/dysoxic conditions and deposition of organicerich, MPRSelike sediments could only take place in marginal embayment areas such as VdM and MJ, where the weak orbital forcing exerted on local evaporation/precipitation rates and marine circulation by precession cycle (iecycle) 74, at the onset of MIS 19 (Lourens, 2004; Konijnendijk et al., 2015), was magnified (e.g., Tyson and Pearson, 1991; Melles et al., 2013).

4.4. Beryllium In our beryllium record (Table 1; Fig. 7), authigenic 10Be and 9Be concentrations oscillate generally out of phase, demonstrating that they respond to different sources and/or transport mechanisms to the sedimentation area. In marginal marine settings such as VdM, 9 Be concentrations are essentially likely to reflect the terrigenous
s et al., 1989; von siliciclastic yield to the basin margin (e.g., Bourle Blanckenburg et al., 2012; von Blanckenburg and Bouchez, 2014). In contrast 10Be, a radioactive cosmogenic isotope with a halfelife of ca. 1.39 Myr (Chmeleff et al., 2010; Korschinek et al., 2010), can be conveyed by vastly different carriers other than the pure atmospheric fallout, such as mineral fragments, airborne dust, soil particles, vegetal debris, etc. (e.g., Stensland et al., 1983; Pavich et al., 1986; You et al., 1988; Barg et al., 1997; Graly et al., 2011; Bacon et al., 2012; Dixon et al., 2018; Mishra et al., 2019, and references within). Our 9Be concentrations (Table 1; Fig. 7a) range from ca. 3.6  1016 to ca. 34  1016 atoms per gram of sediment. Maximum values are restricted to a prominent spike centered at ca. 4 cm level, which we arbitrarily limited to concentrations >7  1016 at/g. Accordingly, the 9Be peak extends over a ca. 45 cm thick interval straddling the 6 cmethick “Pitagora ash” layer (Capraro et al., 2017). This scenario is consistent with a massive injection into the water column of pyroclastic material rich in highly soluble 9Be, this being very abundant in mantle sources (Baroni et al., 2011). Once the geochemical anomaly associated to the “Pitagora ash” is graphically removed (Fig. 7b), the 9Be record reveals a longeterm trend of 9Be concentrations diminishing upwards, although a dense succession of oscillations is shown that follow very closely the d13C record of M. barleeanum (Fig. 7b). As confirmed by both the abundance patterns of calcareous nannofossils (Fig. 4), which reflect the local budget of biogenic/pelagic carbonate production and export (Capraro et al., 2017), mineralogic data (Fig. 5) and trends in the 9Be record, Facies A2 and Facies B preserve a signal of decreasing terrigenous siliciclastic yield to the basin, probably in response to the gradual deepening and shift towards offshore conditions (Fig. 2). When compared to the 9Be record, authigenic 10Be (decayecorrected) concentrations are considerably smaller, ranging from ca. 0.43 to 2.24  108 atoms per gram of sediment (average of

Table 1 Beryllium data from the Valle di Manche section. Authigenic trations are given as atoms/gram of sediment.

10

Be and 9Be concen-

Level (m)

Authigenic 10 Be (Eþ8 at/g)

9 Be (Eþ16 at/g)

Authigenic (E-8)

3.48 3.38 3.29 3.20 3.11 3.01 2.92 2.83 2.74 2.65 2.55 2.46 2.37 2.28 2.19 2.09 2.00 1.91 1.82 1.73 1.63 1.54 1.45 1.36 1.27 1.17 1.08 0.99 0.90 0.81 0.71 0.62 0.53 0.44 0.35 0.25 0.16 0.07
0.10
0.20
0.30
0.40
0.50
0.60
0.70
0.80
0.90
1.00
1.10
1.20
1.30
1.40
1.50
1.60
1.70
1.80
1.90
2.00
2.10
2.20
2.30
2.40
2.50
2.60
2.70
2.80
2.90
3.00

2.009 2.068 2.051 2.089 2.166 2.054 2.240 2.032 1.627 1.488 1.569 1.527 1.471 1.366 1.280 1.228 1.128 1.006 1.046 0.995 0.945 0.939 0.936 0.911 0.928 1.008 0.766 0.788 0.798 0.816 0.794 0.891 0.816 0.764 0.759 0.795 0.577 0.430 0.814 0.864 0.967 0.931 1.031 1.076 1.189 1.264 1.223 1.107 1.141 1.228 1.265 1.269 1.343 1.198 1.389 1.399 1.344 1.429 1.385 1.319 1.242 1.322 1.167 1.148 1.146 1.159 1.065 0.960

3.913 4.125 4.174 4.318 4.293 3.568 3.631 3.872 3.616 3.732 4.115 4.299 3.831 3.978 4.153 4.138 4.029 3.793 3.959 3.877 4.215 4.070 4.300 4.036 4.005 4.196 4.386 4.380 4.784 5.333 5.567 6.657 6.524 5.874 5.602 7.438 17.991 33.928 7.887 7.511 4.503 4.159 4.396 4.632 4.864 4.738 4.676 4.605 4.483 4.467 4.486 4.589 4.687 4.529 4.879 4.632 4.452 5.390 5.101 4.640 4.821 4.925 4.263 4.829 4.068 4.469 3.950 5.221

0.513 0.501 0.492 0.484 0.505 0.576 0.617 0.525 0.450 0.399 0.381 0.355 0.384 0.343 0.308 0.297 0.280 0.265 0.264 0.257 0.224 0.231 0.218 0.226 0.232 0.240 0.175 0.180 0.167 0.153 0.143 0.134 0.125 0.130 0.135 0.107 0.032 0.013 0.103 0.115 0.215 0.224 0.234 0.232 0.244 0.267 0.262 0.240 0.255 0.275 0.282 0.276 0.286 0.265 0.285 0.302 0.302 0.265 0.272 0.284 0.258 0.268 0.274 0.238 0.282 0.259 0.270 0.184

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Be/9Be

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9

Fig. 7. The beryllium record of the VdM section, plotted against depth. at/g: atoms per gram of sediment. a): 9Be concentrations, full record. b): 9Be concentrations limited to values < 6, which emphasize the background dynamics of 9Be in the sedimentation area with respect to the anomalous 9Be peak associated to the emplacement of the “Pitagora ash”; the d13C record of Melonis barleeanum (thin blue line) is also reported for comparison. c): 10Be concentrations (decay-corrected). d): 10Be/9Be ratio, i.e., 10Be concentrations (decay-corrected) normalized to 9Be. Be1 and Be2 indicate the most relevant peaks of 10Be concentration recognized in the local record; PM (Pitagora Minimum) marks the interval of minimal 10Be concentrations at the “Pitagora ash”. Horizontal dashed lines as in the previous figures. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

1.21  108 at/g; Table 1; Fig. 7c). Beginning from very low concentrations in Facies E, 10Be abundances attain a modest peak at
2.00 m, exactly at boundary between Facies A0 and A1 (Be1 in Fig. 7c). Values decrease gradually up to ca.
0.30 m, at the transition from Facies A1 to A2. This horizon marks a dramatic decrease in 10Be concentrations, which attain the absolute minimum at the top of the “Pitagora ash” (Pitagora Minimum: PM in Fig. 7c), where the accelerated injection of 9Beerich pyroclastics further subdued the weak 10Be signal. Between ca. 0.30 and 1.08 m, 10Be values recover but remain persistently low. From this point upwards, 10Be concentrations increase and maximum amounts of 10Be, in the order of twice the background values, are achieved at the top of the condensed muds of Facies B (ca. 2.90 m; Be2 peak in Fig. 7c), which are associated to conditions of maximum local sea level depth and sediment starvation (Fig. 2). The uppermost part of the record, which develops in the lower part of silty Facies C1, is characterized by a minor decrease in 10Be concentrations (Facies C1; Fig. 7c). In order to normalize the cosmogenic flux to the terrigenous
s et al., 1989; input, a10Be/9Be ratio is usually employed (e.g., Bourle Barg et al., 1997). At VdM, the 10Be/9Be record follows very closely that of 10Be, with small differences that are however restricted to the lower part of the section (Table 1; Fig. 7d). In particular, the muted Be1 peak documented in the 10Be curve virtually disappears in the 10Be/9Be record, being replaced by an interval of virtually steady values up to the base of Facies A2 (Fig. 7d). Still, the principal longeterm trends and traits of the 10Be record are preserved, especially the concentration minimum in correspondence with the “Pitagora ash” (PM) and the Be2 peak in Facies B (Fig. 7d). 5. Discussion Marine successions offering both a10Be stratigraphy and a sound paleomagnetic record of the Matuyama-Brunhes transition have been essentially reconstructed in deep oceanesediment cores (e.g., Henken-Mellies et al., 1990; Valet et al., 2014; Simon et al., 2016c,

Simon et al., 2018a,Simon et al., 2018b), which inherently lack an adequate documentation on terrestrial climates and processes at the basin margin. This information is best preserved in open marine successions accumulated in proximal (hemipelagic) settings, where sedimentary influences from the continent are strong and pervasive. In this respect, the continuous and expanded oneland hemipelagic record of the MIS 19 interglacial preserved at VdM, which originated by pure mud settling in a mide to outer shelf environment (Capraro et al., 2017), is a fundamental point of reference. Compared to the open ocean, the central Mediterranean is an extremely sensitive area where the sedimentary record allows for an accurate registration of both the regional and global climatic variability (e.g., Krijgsman, 2002). In this respect, the San Mauro sub-basin provides the opportunity to compare directly between a higheresolution beryllium stratigraphy and a sharp paleomagnetic record of the Matuyama-Brunhes boundary in the context of a comprehensive set of marine and terrestrial paleoclimatic proxies (Capraro et al., 2005, 2017; Massari et al., 1999, 2007; Scarponi et al., 2014; Rossi et al., 2018; Azzarone et al., 2018). Besides VdM, oneland open marine stratigraphic records straddling the MIS 19 interglacial to compare to are unfortunately very scarce. The Chiba section, recently investigated for its beryllium record (Simon et al., 2019), was deposited in a deep marine setting (lower slope; Kazaoka et al., 2015) at the western margin of the Pacific Ocean. Numerous physical stratigraphic, sedimentary and paleontological evidences (e.g., Ito and Kathura, 1993; Ito, 1997, 1998; Mulder et al., 2008; Fukuda et al., 2015; Kazaoka et al., 2015; Suganuma et al., 2015, and references therein) demonstrate that the local sedimentation style was strictly turbiditic. Unfortunately, depositional settings subjected to the accumulation of massive, coarseegrained turbiditic bodies are inherently characterized by episodic (discontinuous) sedimentation and prone to the development of stratigraphic unconformities, which ultimately result in stratigraphic gaps and an uneven documentation of the geological record (Normark et al., 1993; Mutti et al., 2003, 2009). A pervasive

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occurrence of these adverse geological processes have been demonstrated to occur throughout the Kazusa Group, which includes the Chiba composite section (e.g., Ito, 1996, 1998; Kitazaki and Majima, 2003; Ito et al., 2014). Therefore, the time, mode and tempo of the most relevant conveyors of 9Be and 10Be to the depocenter at Chiba were largely different from those recognized at VdM and, in absence of further independent constraints and evidences, we cannot consider the Chiba section as a suitable analog for the scope of this work. The MJ section (Ciaranfi et al., 2001, 2010) crops out in the northern Ionian Sea, only ca. 140 km far from VdM (Fig. 1). The interval straddling MIS 19 at MJ was laid in depositional (Ciaranfi et al., 2010), climatic and environmental conditions that are very similar to those documented at VdM (e.g., Capraro et al., 2005; Bertini et al., 2015, and references within). Regrettably, the MJ section lacks a paleomagnetic record, due to postedepositional remagnetization (Sagnotti et al., 2010). Still, it preserves an extraordinarily rich and manifold data set (e.g., Ciaranfi et al., 2001, 2010; D’Alessandro et al., 2003; Stefanelli, 2003; Girone et al., 2013; Bertini et al., 2015; Nomade et al., 2019) that also includes a higheresolution beryllium record (Simon et al., 2016a). It follows that, for the time being, the MJ section provides the most appropriate and reliable benchmark for testing the consistency and implications of the results achieved at VdM.

5.1. The beryllium record in the Mediterranean A simple visual correlation (Fig. 8) demonstrates that the 10Be, Be and 10Be/9Be records reconstructed at VdM compare extremely well with those obtained at MJ (Simon et al., 2016a). The overall good match is confirmed by the similar d18O stratigraphies for M. barleeanum, demonstrating that both records are essentially in phase (Fig. 8). Background 9Be concentrations are consistently larger at VdM (Fig. 8). This difference is likely to reflect the terrigenous sources for each basin, these being 9Beerich Paleozoic crystalline rocks at VdM (Ogniben, 1955; Ryan, 2002) and recent marine muds, supplied from the emerging Apennines, at MJ (Boccaletti et al., 1990; Popov et al., 2004, 2006; Vitale and Ciarcia, 2013). Furthermore, the 9Beebearing fraction is probably more diluted by 9Beefree deposits at MJ, where the sediment accumulation rates (between ca. 80 and 170 cm/kyr; Simon et al., 2016a) are significantly higher than those calculated at VdM (between ca. 20 and 94 cm/kyr; Capraro et al., 2017). We plotted the 10Be and 10Be/9Be records of VdM and MJ versus time (Fig. 9) according to the age models developed individually for each section by Capraro et al. (2017) and Simon et al. (2016a), respectively. Three prominent tie points can be identified in both records (these being the V3 minimum, Pitagora minimum and Be2 peak in Fig. 9), which provide obvious correlation between the 9

Fig. 8. The beryllium and d18O records reconstructed for the VdM and MJ sections. Data series are plotted against depth, according to the thicknesses measured for each of the stratigraphic section, but to different scales (thicknesses are indicated separately for each section by axis labels) in order to ensure the best visual overlap. Horizontal scales are either ascending or descending, as indicated by the empty triangles, and may refer either to both or individual sections, as indicated by axis labels. Horizontal dashed lines indicate the stratigraphic position of the V4 tephra at MJ and the “Pitagora ash” at VdM, according to their own vertical reference scale. a): 10Be concentrations, decay-corrected (MJ in dark orange, VdM in light blue). b): 9Be concentrations (MJ in light orange, VdM in grey). Chrons are referred to the paleomagnetic record obtained in the VdM section. c): 10Be/9Be ratio (MJ in red, VdM in blue). d): d18O records of M. barleeanum for MJ (orange line) and VdM (green line) compared to the d18O of U. peregrina reconstructed at VdM (purple line). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

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Fig. 9. Comparison between the 10Be/9Be (left) and 10Be, decay corrected (right) measured for the VdM (black line) and MJ (blue line) records. The Be1 and Be2 peaks in the measured 10Be record are indicated. Data are plotted against age, according to the chronological models proposed by Capraro et al. (2017) for VdM and Simon et al. (2016a) for MJ. Different horizontal scales have been employed for plotting 10Be values according to the different concentrations of 10Be at VdM and MJ. Horizontal orange bars indicate the vertical spread of the ages associated to the prominent tie points between the sections (numerical values are also indicated), which indicated by couplets of identical symbols (stars, triangles, squares). Note that the V3 minimum at MJ falls outside the age range considered in the drawing (as indicated by the ? symbol), however it appears in the 10Be/9Be record. The V4 minimum was not considered for correlation, because a) it falls very close to the prominent Be2 peak and b) its signature at VdM is very weak. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

sections. Accordingly, the chronological offset between the curves is  2 kyr, which is virtually negligible as the uncertainty associated to the age models in question is in the order of ±5 kyr (Capraro et al., 2017). This tight correspondence between VdM and MJ not only confirms that the individual chronological frameworks developed for both sections are sound and consistent, but also proves that their stable isotope and 10Be records are fully validated at the regional scale. In both records, prominent spikes in 9Be concentrations are found in correspondence to pyroclastic layers, these being the V4 tephra at MJ and the “Pitagora ash” at VdM (Fig. 8). The “Pitagora ash” is accountable in both sections for the PM minimum, which provides a robust correlation point (Figs. 8 and 9). In spite of this pronounced 10Be signature, which is further emphasized by a concomitant, yet small, peak in 9Be concentrations, no physical evidence exists of the “Pitagora ash” in the MJ section (Petrosino et al., 2015). Still, in their physical description of the MJ section, Ciaranfi et al. (2001) reported on the presence of dispersed pumice fragments at the very same stratigraphic level as the PM. It is thus tempting to infer that the “Pitagora ash” was produced also at MJ, but dismantled afterwards. Similarly, both the V3 and V4 tephra layers of MJ (ca. 774 and 801 ka, respectively; Petrosino et al., 2015) are not found at VdM, where their correlative 10Be signatures can be however detected. In particular, we identified a tentative “V3 minimum” at the very base of the VdM record (Fig. 9) and a more persuasive “V4 minimum” in correspondence to the muted 10Be decrease at the top of the VdM section, immediately above the Be2

peak (Figs. 8 and 9). Notably, both these 10Be minima are not associated to increases in 9Be concentrations at VdM. Lower sediment accumulation rates can be invoked for explaining the higher background 10Be concentrations at VdM, which however mimic very closely the oscillations recorded at MJ (Fig. 8). As expected, 10Be concentrations in both the Mediterranean records are considerably lower than those documented in the much vaster open ocean basins (e.g., Simon et al., 2018a,Simon et al., 2018b, and references within), where 10Be concentrations at the seafloor are considerably larger due to the combined contributions from the atmosphere, scavenging and advection (e.g., von Blanckenbourg and Bouchez, 2014). The prominent Be2 peak recorded in the upper part of the VdM section, albeit documented only partially, is fully consistent, both in terms of timing and shape, with that found at MJ, which Simon at al. (2016a) proposed as the geochemical marker of the Matuyama-Brunhes boundary. 5.2.

10

Be and the Matuyama-Brunhes boundary

As anticipated, the VdM section preserves the only known dependable record of the Matuyama-Brunhes reversal for the whole Mediterranean oneland marine record (Head and Gibbard, 2015). At VdM, the Matuyama-Brunhes boundary is located in the midst of full MIS 19, at ca. 12.5 cm, immediately above the “Pitagora ash” (Fig. 3; Macrì et al., 2018). Based on the chronology proposed by Capraro et al. (2017), an age of ca. 786 ka was obtained for the reversal, which falls in the older part of the age range globally

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associated to the Matuyama-Brunhes transition (Head and Gibbard, 2015) but very close to that calculated by Sagnotti et al. (2014) for the Sulmona lacustrine record (Central Italy) and Mark et al. (2017) (see Macrì et al., 2018 for discussion). In the MJ section, where magnetostratigraphic investigations proved to be inconclusive (Sagnotti et al., 2010), Simon et al. (2016a) proposed to locate the geomagnetic reversal in correspondence to the midepoint of a prominent 10Be spike that, based on an astronomicallyederived age and geochronologic evidences (age of the V4 tephra; Petrosino et al., 2015), was calculated to fall at ca. 774 ka (beginning of the MIS 19 e MIS 18 transition). This chronology is in keeping with the Matuyama-Brunhes reversal age obtained for a number of North Atlantic sediment cores (Channell et al., 2009), and consistent with that derived by the recognition of a10Be peak in the midst of a geomagnetic relative paleointensity (RPI) minimum in openeocean sedimentary successions (e.g., Simon et al., 2018a,Simon et al., 2018b). Nevertheless, recognition of the Matuyama-Brunhes transition at MJ, and especially its precise stratigraphic positioning, are purely speculative, since the 10Be evidence is not further crossevalidated by paleomagnetic measurements and/or RPI determinations. Additional evidences call for a cautious interpretation of the MJ record. Oceanic 10Be records (Simon et al., 2018a,Simon et al., 2018b) are generally characterized by the occurrence of a massive peak of 10Be overproduction extending over most of the RPI minimum, which is estimated to last between ca. 15 and 20 kyr (Valet et al., 2005; Channell et al., 2009), while the calculated duration of the main 10Be peak at MJ is dramatically shorter, in the order of ca. 4 kyr (Simon et al., 2016a). Furthermore, the simple anatomy of the 10Be peak preserved globally in the oceanic sediment cores contrasts sharply with the record found at MJ, which exhibits a swarm of 10Be spikes, of similar magnitude to the main peak, above the putative Matuyama-Brunhes boundary, at the late MIS 19eMIS 18 transition (Simon et al., 2016a). Interestingly, such 10 Be peaks, which have no obvious geomagnetic substantiation, are in phase with the oscillations documented in the in local d18O stratigraphy (Simon et al., 2016a), suggesting that environmental/ climatic conditions e such as seaelevel, regional precipitations and temperatures e might have influenced the buildup of the local 10Be record. At VdM, it was not possible to locate and constrain stratigraphically the RPI minimum associated to the Matuyama-Brunhes reversal, because the relatively coarseegrained lithologies immediately above and below the investigated segment (Capraro et al., 2015, 2017) are not amenable to paleointensity determinations (Macrì et al., 2018). However, our paleomagnetic data (Macrì et al., 2018) suggest that the stratigraphy investigated for this work falls completely within the interval of geomagnetic intensity low. The 10Be spike at VdM occurs ca. 3.5 m above the MatuyamaBrunhes boundary that, based on our chronology (Capraro et al., 2017; Macrì et al., 2018), corresponds to a ca. 12 kyr interval. Evidences of a decoupling between the 10Be peak and the MatuyamaBrunhes transition, this being older by ca. 10 kyr, are also found in openeocean sediment cores and terrestrial sequences with low sediment accumulation rates, where the decoupling is usually explained by the combined effects of lockein processes and bioturbation (e.g., Suganuma et al., 2010; Zhou et al., 2014). This explanation is not justified for the expanded VdM section, where the diagenetic overprint possibly related the occurrence lock-in effect, if any, could not account for the important stratigraphic distance between the Matuyama-Brunhes boundary and the 10Be peak (Be2; Fig. 7; e.g., Sagnotti et al., 2005). 5.3. Climate, oceanography and the

10

BE record

The 10Be record produced within landlocked basins such as the

San Mauro sub-basin, where sedimentation is hemipelagic and takes place in a shelf setting, may originate from a manifold array of contributions from both atmospheric and terrigenous sources, such as aerosols, mineral fragments, plant debris, soil particles, humic acids, etc. (Pavich et al., 1986; Bacon et al., 2012; Maher and von Blanckenburg, 2016). Efficiency of these mechanisms is likely to change in time, according to both the inherent capacity of 10Be reservoirs, type of substrate which 10Be atoms are adsorbed onto, and the strength of potential 10Be carriers to the depocenter, these being precipitations, coastal runoff, winds, ocean currents, etc. (e.g., Anderson et al., 1990; Kusakabe et al., 1991; Brown et al., 1992; Lao et al., 1992; von Blanckenburg et al., 1996; Frank et al., 1997; Igel and von Blanckenburg, 1999; Wagner et al., 2001; Mishra et al., 2019). In order to single out the environmental parameters that mostly influenced the 10Be flux to the San Mauro sub-basin and gauge their efficiency as 10Be carriers, we attempted a simple compositional approach by means of the Solver© tool, as described above. The VdM 10Be record was selected as the target curve, to which each potential 10Be carrier and/or related transport process (“proxies” hereafter, for the sake of simplicity) was compared. With reference to the case in question, the analysis was expected to provide either of the following scenarios: a) a substantial covariance/anticovariance exists between the measured 10Be series and the “calculated” curve, implying that the proxies considered for computation oscillate coherently and in phase with 10Be concentrations. This scenario would suggest that 10Be deposition was not independent from the local environmental and climatic conditions, these being possibly involved in the transport and/or stabilization of 10Be to the seafloor; b) no correlation is observed between the “calculated” series and the 10Be record, suggesting that no relationships exist between 10 Be concentrations and the regional climatic and environmental conditions considered for calculation. This is the expected (and most desirable) scenario when interpreting the 10Be record as a pure cosmogenic signal and an objective marker of the Matuyama-Brunhes boundary, as in the case of the MJ section (Simon et al., 2016a). According to Simon et al., 2016a, the 10Be record at MJ oscillates independently from the local climatic and environmental changes, so that it reflects solely the variable productions rates of cosmogenic 10Be in the atmosphere. By contrast, a number of climatic and environmental proxies seem to follow the oscillations observed in the beryllium record at VdM (Figs. 4e7). In particular (Fig. 10), authigenic 10Be concentrations are approximated very closely by the calculated record based on the relative abundances of “mountain” conifers pollen (% on the total pollen assemblage, Fig. 4; Capraro et al., 2017) and the calcite content in bulk sediments (wt%; Fig. 5). Both the proxies in question hold a climatic and environmental significance that is fully consistent with the potential 10Be dynamics in a marginal marine setting, as discussed henceforth.

5.3.1. “Mountain” pollen This informal definition points to the palynological contribution provided by mesoe and microthermal wateredemanding conifers such as Abies, Picea, Cedrus and Tsuga. In the Pliocene and Pleistocene pollen record of the central Mediterranean area, increasing abundances of this vegetational assemblage are associated to the development of a yeareround rainfall regime and high precipitation rates, possibly in concomitance with decreasing temperatures zel and Me dail, 2003). Based on presenteday (e.g., Bonin, 1981; Que

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Fig. 10. Calculated values for the proxies considered in our computation (see text for details). Data are plotted against depth. a) Calculated values for calcite (wt%) as proxy: r ¼ 0.05, k ¼
0.25, limited to the considered interval. b) Calculated curve for “mountain” pollen (% on the total pollen sum) as proxy: r ¼ 0.03, k ¼ 0.59 for the Facies E - A1 interval; r ¼ 0.03, k ¼ 0.01 above. c) Combined calculated values for calcite and “mountain” pollen as proxies (red line), computed arithmetically via merging of the component records. d) Comparison between the combined record (red line) and the 10Be record measured at VdM (green dots). SD ¼ 0.05 for the lower part of the record, where calcite was not computed; SD ¼ 0.12 above, where both proxies have been considered. Numbers 1e3 indicate small-scale peaks in the merged calculated curve. The Be1 and Be2 peaks in the measured 10Be record are also indicated. Horizontal dashed lines as in the previous figures. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

analogs, the expansion of “mountain” conifers group in Southern Italy suggests an average precipitation in excess of 2000 mm/yr (Fenaroli, 1970; Noirfalise et al., 1987; Capraro et al., 2005), a threshold significantly higher than that observed today in the Ionian Calabria region (e.g., Ozenda, 1975; Bonin et al., 1976; Bonin, zel and Me dail, 2003). The pollen signal 1981; Gentile, 1982; Que from this vegetational group is especially well represented and dependable in the Pleistocene records of the Crotone Basin if compared to the coeval perieApennine basins, because the Sila massif, which exceeded the altitude of 2000 m a.s.l. since Pliocene times (Fauquette and CombourieueNebout, 2013), allowed for the growth of extensive altitudinal forests. In marginal marine settings surrounded by steep catchment basins, such as the San Mauro subbasin, increased precipitations are expected to promote the massive mobilization and riverine transport of coarseegrained sediments to the depocentral areas (e.g., Posamentier et al., 1988; Pratson et al., 2007), along with vegetal debris and pollen from the higher elevations (e.g., Mudie, 1982; Beaudouin et al., 2007). Instead, the higher abundances of “mountain” pollen across the studied segment are found in the very fineegrained sediments of Facies B (Fig. 4), which reflect conditions of sediment starvation (Scarponi et al., 2014). Massari et al. (2007) explained this apparent paradox by observing that the climatic conditions associated to the development of “mountain” conifer communities would favor the growth of a thick forest coverage all over the catchment basins surrounding the San Mauro sub-basin, with dramatic decreases in the potential land erosion (Mishra et al., 2019) and, ultimately, in the conveyance of coarseegrained sediments to the basin margin. Associated to the development of a densely forested landscape is

the widespread formation of deep, humic soils that e provided that their pH is > 4, these being the vast majority of soils found in mide and lowelatitudes (Van Breemen et al., 1983; Jenny, 1994; Slessarev et al., 2016) e behave as huge reservoirs of meteoric 10Be (Monaghan et al., 1983; Pavich et al., 1984, 1985; Willenbring and von Blanckenburg, 2010a,b). In addition to exerting a mechanical strain on the emerging 10Beerich organic soils (Mishra et al., 2019), meteoric waters can exploit their moderately acidic pH to promote the chemical desorption and solubilization of 10Be from terrestrial mineral substrates (You et al., 1988; Boschi and Willenbring, € m et al., 2018). Once 2016a,Boschi and Willenbring, 2016b; Åstro conveyed via coastal runoff to the sea, where pH is slightly alkaline (~8; e.g., Stumm and Morgan, 1981; Hofmann et al., 2011), terrestrial 10Be particles may be eventually adsorbed and stabilized onto the resident mineral fraction. Altogether, the remarkable similarity that is observed between the abundance of “mountain” pollen and 10 Be concentrations at VdM (Fig. 10) suggests that the local 10Be yield was mostly provided by terrestrial sources, according to the regional variability in rainfall rates and regimes (e.g., Lundberg et al., 1983). 5.3.2. Calcite The total amount of calcite preserved in the fine sediment fraction at VdM was essentially supplied by two independent inputs, these being a biogenicepelagic snow, mainly composed of planktic foraminifer tests and calcareous nannofossils, and a terrigenous (extrabasinal) yield. Intensity of both these carbonate flows is likely to have changed in time across the interval of relevance, according to the local primary productivity (for the pelagic

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snow) and/or the extent and composition of the mineral sediment yield to the basin margin (for the terrigenous fraction). Regardless of the primary source, the carbonate flux to the seafloor was subjected to changing dissolution rates, in response to the chemical conditions (especially pH) of the water column. A rough estimate of the terrigenous carbonate input at VdM was obtained by Capraro et al. (2017) by comparing the relative abundance of reworked nannofossils to the “native” population, demonstrating that the extrabasinal input across the studied interval was generally very small to negligible. In support to this scenario is the dearth of significant carbonate sources exposed in the surroundings of the San Mauro sub-basin, where igneous rocks and siliciclastic sediments are dominant (e.g., Ogniben, 1955; Roda, 1964; Massari et al., 2002, 2010; Capraro et al., 2006, 2011). In fact, longeterm trends in 9Be concentrations, which approximate the overall terrigenous input (Fig. 7), are not in phase with the calcite record (Figs. 5 and 7). Therefore, limited to the interval of relevance, we may interpret the fine carbonate fraction at VdM as essentially provided by biogenicepelagic sedimentation. Changes in the relative concentration of biogenicepelagic calcite in offshore muds reflect the intensity of biological productivity in the surface waters, which has important implications on the distribution of 10Be into the water column (Kusakabe et al., 1987, 1990). Calcite itself is a poor beryllium carrier (Hervig, 2002), as demonstrated by the very low beryllium concentrations found in carbonate sediments (Ryan, 2002). However, 10Be is seized vigorously by organisms living in the upper photic zone because of its nutrientelike behavior (Measures and Edmond, 1982; Kusakabe et al., 1987) to form strong, pHeindependent ligands with the organic matter (Boschi and Willenbring, 2016a,Boschi and Willenbring, 2016b). Once stabilized onto an organic carrier, beryllium can only be released to solution upon digestion of the substrate in an oxidant environment (Kusakabe et al., 1987). Ideally, this implies that the biogenicepelagic sedimentation (calcite þ organic matter) in an offshore setting should be at equilibrium with the concentration of meteoric 10Be in the upper water column. In “normal” conditions, however, most of the organic matter e and its 10Be content as well e is promptly recycled within, or immediately below, the photic zone (e.g., Meyers and Eadie, 1993; Hedges and Keil, 1995), thus hindering the 10Be flow to the seafloor. Therefore, it is unlikely that the sedimentary record, even in fortunate circumstances, may offer an exact estimate of the 10 Be concentrations in the upper water column at a given time. Differently from the organic matter, 10Be retention on mineral sedimentary particles at the seafloor depends strictly on the environmental pH (Boschi and Willenbring, 2016a,Boschi and Willenbring, 2016b). Specifically, the moderately alkaline pH associated to “normal” (oxic) conditions would favor the stabilization of beryllium onto mineral substrates, while acidic conditions e especially in the presence of sulfur and phosphorous (Boschi and Willenbring, 2016a,Boschi and Willenbring, 2016b) e would promote its desorption and release to solution (Dickin, 2018). At VdM, anomalous (sulfidic) conditions at the seafloor can be inferred to have occurred during the deposition of the sapropelelike interval in Facies A0 and A1 (Figs. 2, 5 and 6). As expected, minerals sensitive to acidic environments, such as calcite, were partly damaged and/or dissolved, as demonstrated by the overall bad preservation of the calcareous microe and meiofaunas (Capraro et al., 2017; Rossi et al., 2018). 10Be formerly adsorbed on the dissolving carbonate particles was therefore released to solution, while that carried by the organic matter (possibly the most abundant fraction; Kusakabe et al., 1987) was retained and preserved into the sediment. In the stratigraphy above and below the sapropelelike layer, oxidic (alkaline) conditions allowed for a prompt and thorough digestion of the organic snow and release to

solution of its beryllium content, which could be reclaimed and stabilized onto the mineral fraction (Boschi and Willenbring, 2016a,Boschi and Willenbring, 2016b). The postedepositional oxidation suffered by the organicerich sediments of Facies A0 and especially A1 (see discussion above) suggests that 10Be scavenging from the organic matter might have commenced immediately at the reestablishment of oxidic conditions (i.e., “burnedown” of the organic matter buried into the sapropelic layer; Higgs et al., 1994; Van Santvoort et al., 1996). If that were the case, it is likely that 10Be released to solution from the decaying organic matter was partly advected by bottom currents, partly stabilized onto the local mineral substrate. Nonetheless, a stratigraphic displacement is expected to have occurred with respect to the “source” layer. 5.3.3. Assessing the climatic and environmental influence on the Mediterranean 10BE record In our calculation, we discarded the contribution of calcite in the stratigraphy below the “Pitagora ash” (Facies A0 and A1), where the available evidences point to a pervasive dissolution and/or bad preservation of carbonates that would distort the results. Still, the calculated pollen record alone provides an excellent agreement to the measured 10Be concentration curve (Fig. 10). From the bottom of Facies A2 upwards, where both proxies can be employed, the calculated values follow the measured 10Be curve (Fig. 10). Three prominent oscillations, generated by the mutual contributions of both pollen and calcite (1e3 in Fig. 10), are however observed. Among these, only the upper and most prominent one (3 in Fig. 10) finds a match in the measured record. Considering in our calculation other potential proxies, such as the foraminiferal d18O and d13C, sediment grain size, etc., did not improve the overall correspondence. At VdM, the records of both 9Be (Fig. 7) and calcareous nannofossils (Fig. 4) point to a longeterm decrease of the terrigenous siliciclastic input and a concomitant increase in the flux of pelagicebiogenic calcite up to the top of Facies B (Scarponi et al., 2014; Rossi et al., 2018). The upwardeincreasing disagreement between the measured and calculated 10Be curves in Facies A2 and B (Figs. 10 and 11) may therefore reflect a gradual loss in sediment sensitivity to the beryllium input, in response to decreasing concentrations of clay minerals and organic matter, these being the most efficient substrates for 10Be adsorption and stabilization in our depositional setting (Boschi and Willenbring, 2016a,b). By contrast, the 10Be record at MJ exhibits a number of shorteterm oscillations in the correlative stratigraphic interval (Figs. 8, 9 and 11). Possibly, sediment composition and accumulation rates at MJ allowed for a more detailed registration of smallescale changes in the regional 10Be input compared to VdM, which in turn offers a more comprehensive depiction of the climatic and environmental evolution on mainland across the MIS 19 interglacial. Because the time, mode and tempo of the geochemical records reconstructed at VdM and MJ are very similar (Figs. 8, 9 and 11), we attempted reconstructing a common age model for both the sections based on the tie points indicated in Fig. 10. For the sake of simplicity, ages were calculated by linear interpolation relative to the “Pitagora minimum”, which provides an obvious marker in both records (Fig. 9). An impressive correlation can be observed (Fig. 11) between the combined influx of “mountain” pollen þ calcite calculated at VdM and the 10Be values measured at MJ, especially in the upper part of the succession, where peaks 1 and 2 in the “calculated” curve match the local 10Be record. This correspondence is also confirmed by a weak but congruent pollen signal from “High altitude forest taxa” at MJ (Bertini et al., 2015). Our results are hence validated at the regional scale, implying that a strong environmental imprint exists on the conveyance and preservation of 10Be at the northern margin of the Ionian Sea. In fact, if

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Fig. 11. Comparison between the combined record of calcite and “mountain” pollen, as calculated at VdM, and the 10Be record of VdM (left panel) and MJ (right panel). Data are plotted against time, considering the prominent 10Be minimum associated to the “Pitagora ash” as the reference zero value. For the MJ section, ages were recalculated by linear interpolation between the tie points presented in Fig. 9 (here indicated in red), for which we retained the chronologies proposed by Capraro et al. (2017) for the VdM section. Numbers 1e3 indicate small-scale peaks in the “calculated” curve; a-c mark the corresponding peaks in the 10Be record at MJ. The Be1 and Be2 peaks in the 10Be record measured at VdM are also indicated. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

the Be2 peak (Fig. 11) were solely the result of pure beryllium fallout and settling following a period of anomalous 10Be overproduction in the atmosphere, 10Be concentrations measured at both VdM and MJ would exceed significantly those calculated, these being representative of the amount of 10Be conveyed via riverine transport. Instead, no significant decoupling exists between the measured and calculated curves, which demonstrates that the regional climate variability and oceanographic conditions may be held fully accountable for the 10Be oscillations observed at both VdM and MJ, independently from the changes in cosmogenic 10 Be concentrations in the atmosphere. It is likely that the transport and stabilization of 10Be followed two main pathways, namely: a) mobilization of terrestrial reservoirs (Graly et al., 2010; West et al., 2013; Willenbring and Von Blanckenburg, 2010b) consistent with the regional rainfall rates and regimes, which ultimately control the amount and composition of the solid and dissolved riverine loads conveyed to the basin; 2) changes in the intensity and composition of the biogenicepelagic and terrigenous yields to the seafloor, in response to climate, marine biological productivity and riverine input. Water chemistry, especially pH values, would ultimately rule the 10Be stabilization onto sediments (Boschi and Willenbring, 2016a,Boschi and Willenbring, 2016b). Independently from the dominant process at any given time, the mineralogic composition of sediments is likely to play a key role in controlling the local sensitivity to the 10Be input, being carbonates sensitive to acidic conditions and less dependable than clayey muds as substrates for the adsorption of beryllium (Hervig, 2002; Ryan, 2002). Our results challenge the interpretation of 10Be records

reconstructed for the central Mediterranean hemipelagic record as a genuine representation of the sole meteoric 10Be fallout and settling. In fact, the tight correlation between the measured 10Be and calculated records (Fig. 11) proves that the 10Be peak found at VdM and MJ does not provide an independent and unbiased depiction of the period of cosmogenic 10Be overproduction and fallout that is expected to have occurred globally at about the Matuyama-Brunhes boundary (e.g., Simon et al., 2018a,Simon et al., 2018b, and references within). Our data confirm that, in marginal marine settings, the terrestrial 10Be input can be large enough to completely overshadow the cosmogenic flux (Brown et al., 1988; You et al., 1988; Simon et al., 2016b). It is worth stressing that, limited to the available Mediterranean records, intervals of higher 10 Be concentration correlate with periods of climate cooling and/or increasing rainfall. This is especially obvious in the upper part of the MJ record (Simon et al., 2016a; Nomade et al., 2019), at the late MIS 19eMIS 18 transition, where the d18O of M. barleeanum and 10Be values oscillate perfectly in phase, with higher 10Be concentrations during “cold” d18O events (Simon et al., 2016a). This pattern is also observed at the main 10Be peak, which occurs at both VdM and MJ at the early demise of MIS 19 under conditions of increasing d18O (i.e., climate cooling; Capraro et al., 2017; Simon et al., 2016a). Possible correlations between changes in the riverine 10Be yield and climate are not completely surprising. The evidence of an orbitallyepaced forcing on sedimentation and biota, where regional precipitation rates and regimes play a pivotal role, is pervasive in the MJ and VdM sections (e.g., Bertini et al., 2015; Capraro et al., 2005; Ciaranfi et al., 2010; Girone et al., 2013;

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Nomade et al., 2019; Simon et al., 2016a) and, more in general, throughout the extensive Cenozoic stratigraphic record exposed in the central and eastern Mediterranean (e.g., Krijgsman, 2002; Rio et al., 2003; Giusberti et al., 2019), as especially well documented at the deposition of MPRS layers (e.g., Rohling, 2015, and references within). 6. Conclusions Data discussed in this paper do not, and are not intended to, put in question the correlation between conditions of weak geomagnetic dipole moment and periods of massive overproduction and fallout of meteoric 10Be, which is convincingly demonstrated by deepesea records globally. Instead, they suggest that 10Be records in marginal marine settings subjected to terrigenous sedimentation might not be respectful of the primary 10Be concentrations in the atmosphere at that very moment, as variable amounts of authigenic 10 Be remobilized from terrestrial reservoirs may distort the pure marine signal. In addition, changes in water chemistry and sediment composition may disrupt the stabilization of beryllium at the seafloor. Ultimately, our conclusions challenge the potential of the 10 Be/9Be approach as a substitute for paleomagnetic investigations, at least in specific stratigraphic contexts, and suggest that most care is needed in interpreting the 10Be signal whenever robust and independent validations are not available. Although our assumptions only arise from the study of two Mediterranean shelf records, implications at a larger scale may exist. Authigenic 10Be released to solution, both from marine and terrestrial sources, can be advected by ocean currents on very long distances before it eventually settles, suggesting that the 10Be signature associated to periods of intense river runoff and/or accelerated soil erosion at the regional scale, such as those induced by the orbitallyemodulated climatic variability, might be felt also in offshore hemipelagic settings, or even beyond. Acknowledgments This research was funded by the University of Padova, Italy (Progetto di Ateneo 2011 to L.C. and ex-60% to L.C. and E.F.). We are grateful to Prof. Antje Voelker, Editor of Quaternary Science Reviews, E. Puejo and two anonymous reviewers for their valuable comments and suggestions that greatly improved the manuscript. Appendix A. Supplementary data Supplementary data to this article can be found online at https://doi.org/10.1016/j.quascirev.2019.106039. References Abu-Zied, R.H., Rohling, E.J., Jorissen, F.J., Fontanier, C., Casford, J.S.L., Cooke, S., 2008. Benthic foraminiferal response to changes in bottomewater oxygenation and organic carbon flux in the eastern Mediterranean during LGM to Recent times. Mar. Micropaleontol. 67 (1e2), 46e68. https://doi.org/10.1016/ j.marmicro.2007.08.006. Andersen, M.B., Matthews, A., Vance, D., BareMatthews, M., Archer, C., de Souza, G.F., 2018. A 10efold decline in the deep Eastern Mediterranean thermohaline overturning circulation during the last interglacial period. Earth Planet. Sci. Lett. 503, 58e67. https://doi.org/10.1016/j.epsl.2018.09.013. Anderson, R.F., Lao, Y., Broecker, W.S., Trumbore, S.E., Hofmann, H.J., Wolfli, W., 1990. Boundary scavenging in the Pacific Ocean: a comparison of 10Be and 231Pa. Earth Planet. Sci. Lett. 96, 287e304. https://doi.org/10.1016/0012e821X(90)90008eL. € € m, M.E., Yu, C., Peltola, P., Reynolds, J.K., Osterholm, Åstro P., Nystrand, M.I., Augustsson, A., Virtasalo, J.J., Nordmyr, L., Ojala, A.E.K., 2018. Sources, transport and sinks of beryllium in a coastal landscape affected by acidic soils. Geochim. Cosmochim. Acta 232, 288e302. https://doi.org/10.1016/j.gca.2018.04.025. Azzarone, M., Ferretti, P., Rossi, V., Scarponi, D., Capraro, L., Macrì, P., Huntley, J.W., Faranda, C., 2018. Earlyemiddle Pleistocene benthic turnover and oxygen isotope stratigraphy from the central Mediterranean (Valle di Manche, Crotone

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