Spatial geochemistry of Upper Jurassic marine carbonates (Iberian subplate)

Earth-Science Reviews 139 (2014) 1–32

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Spatial geochemistry of Upper Jurassic marine carbonates (Iberian subplate) Rute Coimbra a,⁎, Adrian Immenhauser b, Federico Olóriz a a b

Departamento de Estratigrafía y Paleontología, Universidad de Granada, Spain Institute for Geology, Mineralogy and Geophysics, Ruhr Universität Bochum, Germany

a r t i c l e

i n f o

Article history: Received 28 October 2013 Accepted 20 August 2014 Available online 4 September 2014 Keywords: Jurassic Carbonate geochemistry Diagenesis Palaeoceanography

a b s t r a c t Chemostratigraphy applied to ancient marine carbonates is commonly based on one-dimensional (stratigraphic) sections or core data. As demonstrated from modern oceans, this approach underestimates the spatial complexity of physico-chemical seawater properties. Here, a several-hundred-kilometer long transect consisting of seven Upper Jurassic sections from settings ranging from (proximal) neritic middle shelf to the epioceanic (distal) fringe across the southern and eastern palaeomargins of the Iberian sub-plate reveals variability in sedimentologic, stratigraphic, and geochemical records. The comparison of isotopic and elemental data from different carbonate materials (matrix micrite, carbonate cements in veinlets, belemnite rostra, and ammonite shells) reveals differential diagenetic pathways. Microfacies, cathodoluminescence and geochemical data retrieved from biostratigraphically well constrained sections reveal that epioceanic matrix micrite geochemical data provide valuable proxies for palaeo-seawater properties. Our data are reviewed in the context of published Late Jurassic records. The outcome shows a higher level of complexity including the potential admixture of marine, continental, and diagenetic geochemical signals in the epicontinental record. The stratigraphic trend in carbon isotopes of epioceanic sections agrees upon that of Upper Jurassic reference sections from the northern Tethyan margins, while oxygen isotope ratios are relatively 18O-enriched. Palaeo-seawater temperatures across the transect investigated were estimated using δ18O as tentative proxy for interpreting distance from shore, differences in water masses, relative depth variations, and potential local forcing factors. Palaeoenvironmental conditions evaluated through the combined record of isotope (δ18O), elemental (Mn), and skeletal content contrast with relative fluctuations in sea level. Site-specific changes in palaeoceanographic parameters such as water depth, seawater temperature, salinity and upwelling are considered in comparison to examples from ancient and modern oceans. © 2014 Elsevier B.V. All rights reserved.

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Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.1. Characterizing ancient sea waters — a message from the present . . . . . . . . . . . . . . . . . . . . . . 1.2. Palaeoceanography, sea-level fluctuations and variations in seawater geochemistry — a Late Jurassic background 1.3. The Ammonitico Rosso Facies Complex: advances in the last decades . . . . . . . . . . . . . . . . . . . . 1.4. Testing spatial geochemistry across a Late Jurassic palaeomargin — the case study . . . . . . . . . . . . . . Regional setting and study areas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1. Regional setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2. Environmental setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Methods and materials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.1. Sampling strategy and analytic procedure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2. Carbonate and silica materials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Petrographic and geochemical data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1. Microfacies analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.1. Carbonate grains and textures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.1.2. CL patterns . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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⁎ Corresponding author at: Dpto. Estratigrafía y Paleontología, Facultad de Ciencias, Universidad de Granada, Av. Fuentenueva s/n, 18071 Granada, Spain. Tel.: +34 665085651. E-mail address: [email protected] (R. Coimbra).

http://dx.doi.org/10.1016/j.earscirev.2014.08.011 0012-8252/© 2014 Elsevier B.V. All rights reserved.

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4.2.

Carbonate geochemistry . . . . . . . . . . 4.2.1. Carbon and oxygen isotope ratios . . 4.2.2. Elemental geochemistry . . . . . . 5. Interpretation and discussion . . . . . . . . . . . 5.1. Preservation and reliability of C and O records 5.1.1. Carbon isotope composition . . . . 5.1.2. Oxygen isotope composition . . . . 5.1.3. Silica diagenesis . . . . . . . . . . 5.2. The epioceanic record . . . . . . . . . . . 5.3. The epicontinental record . . . . . . . . . . 6. The wider context . . . . . . . . . . . . . . . . 7. Conclusions . . . . . . . . . . . . . . . . . . . Acknowledgments . . . . . . . . . . . . . . . . . . . Appendix 1. . . . . . . . . . . . . . . . . . . . . Appendix 2. . . . . . . . . . . . . . . . . . . . . Appendix 3. . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . .

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1. Introduction 1.1. Characterizing ancient sea waters — a message from the present The record of Earth's climate and changes in marine seawater properties through geological time is largely based on evidence provided by a wide range of proxies, usually lithofacies, fossil assemblages or geochemical information. In Earth sciences, observation and interpretation of geochemical data are commonly derived from one-dimensional section or core, or an averaged “composite curve” from several localities (e.g., Shackleton et al., 1993; Immenhauser et al., 2002; Weissert and Erba, 2004; Föllmi et al., 2006; Dera et al., 2011). Whereas the fundamental validity of time-series data sets from particular sites is not questioned here, modern oceans clearly show complex, three-dimensional organization of water masses with space and time changes in physical–chemical seawater properties (Astraldi et al., 2002; Cardin et al., 2011; Turpin et al., 2012, 2014). Moreover, in coastal settings or in shallow epeiric seas, local influences modify regional or global signals, usually through relative changes in sea-level and climate fluctuations that drive changes in continental runoff (Holmden et al., 1998; Fanton et al., 2002; Panchuk et al., 2005; Dopieralska et al., 2006; Panchuk et al., 2006; Immenhauser et al., 2008). The complexity in trace elemental and isotopic properties of present day oceanic water masses stresses the relevance of spatial approaches, both when modern and ancient marine records are concerned. Oceanographers deal with spatial records of modern marine dynamics at all scales, from global and latitudinal distribution of chemical elements in seawater and marine sediments (e.g., Elderfield, 2006; Dessai et al., 2011; Radic et al., 2011; Slemons et al., 2012), to their ocean or basin-wide characterization including the analysis of selected skeletal hardparts embedded in sediments (Chester, 2000; Lacan and Jeandel, 2001; Amini et al., 2004; Kato et al., 2011; Middag et al., 2011; Wei et al., 2012). Recently, the understanding of the three-dimensional structuring of present-day water masses and components of the related ocean/atmosphere coupling has been improved significantly (Cacho et al., 2000; Gutjahr et al., 2010; Voelker et al., 2010; Radic et al., 2011; van de Poll et al., 2013). On smaller scales, the environmental complexity of present-day proximal-to-distal transects has been demonstrated across particular shelves and basins (e.g., Great Bahama Bank; Swart and Eberli, 2005; Gischler et al., 2009; Swart et al., 2009; Turpin et al., 2012). Despite the vast available information for modern day settings, attempts to interpret the spatial variation of ancient marine water masses from the proxy record of carbonate deposits point to a challenging, at present under-explored, research area. Late Jurassic examples mainly refer to selected epicontinental, intra-basin geochemical

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analyses of bulk and skeletal samples rarely covering the full duration of the Late Jurassic (Riboulleau et al., 1998; Bartolini et al., 2003; Lécuyer et al., 2003; Zakharov et al., 2005; Brigaud et al., 2008). Other workers focused on selected stratigraphic boundaries recorded in structured palaeo-margins (Rais et al., 2007) or through correlation of distant data provided by different authors (e.g., Žák et al., 2011). The inherent incompleteness of the stratigraphic record (e.g., Wetzel and Allia, 2000; El Kadiri, 2002; Dogan et al., 2006; McLaughlin et al., 2008), diagenetic alteration (Dickson and Coleman, 1980; Cicero and Lohmann, 2001; Morse et al., 2007; Macouin et al., 2012) and lateral discontinuity of facies belts (Bhattacharya, 2011) commonly limit attempts to reach a precise characterization of ancient water masses and to separate palaeoenvironmental signals from noise. The present research approaches the chemostratigraphic characterization of ancient carbonates and related sea water properties along a proximal-to-distal transect connecting two major marine palaeoenvironments (i.e., neritic/epicontinental and epioceanic waters). To place the outcome in a broader context, a review of the palaeoenvironmental background during the Jurassic “greenhouse” world is relevant for establishing a link between major palaeoceanographic events and how they may relate to potentially recorded geochemical trends.

1.2. Palaeoceanography, sea-level fluctuations and variations in seawater geochemistry — a Late Jurassic background Numerous palaeoclimatic interpretations of the Jurassic world have been published (e.g., Hallam et al., 1993; Abbink et al., 2001; Hallam, 2001; Gröcke et al., 2003; Cecca et al., 2005; Dera et al., 2009). Still, understanding the links between driving mechanisms and related climatic changes is complex (Hallam, 2001; Sellwood and Valdes, 2006) as exemplified in present oceans. Large-scale (global) modeling has provided useful latitudinal information. Examples include Boreal/Tethyan water mass characterization and general climate inferences (Hallam, 2001); latitudinal characterization of ocean temperature in relation to upwelling influence (Gröcke et al., 2003; Muttoni et al., 2005); or the establishment of latitudinal climate belts based on sedimentary and fossil records (Hallam et al., 1993). Concerning inter-tropical areas, it is not obvious if the temperature–depth structure of Late Jurassic oceans should be compared with that of modern tropical seas (Ziegler et al., 2003). Jurassic “greenhouse conditions” would argue for increased evaporation in low-latitude shallow seas (Ziegler et al., 2003) and the formation of Tethyan warm, saline bottom waters has been proposed (Arthur et al., 1987). As a consequence, the Jurassic Tethys Ocean most likely acted as source of warm and saline waters for much of the world's oceans (Kutzbach and Gallimore, 1989).

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Concerning Jurassic sea-level changes, and particularly those of the Middle to Late Jurassic world, glacio-eustasy may have contributed to variable degrees. This is because the presence of high-latitude ephemeral and/or permanent inland ice caps and/or sea ice cannot be ruled out (Moore et al., 1992; Rowley and Markwick, 1992; Valdes et al., 1995; Price, 1999; Dromart et al., 2003; Immenhauser, 2005), but remains debated (Dera et al., 2011). Moreover, intensified seafloor spreading and volcanic activity forced a combined effect: (1) increased atmospheric CO2 and triggered changes in global climate during the Middle to Late Jurassic (Weissert and Mohr, 1996; Price, 1999; Weissert and Erba, 2004); and (2) tectonic forcing of relative sea-level fluctuations. Hence, tectono-eustasy has been favored as a major trigger for relative sea-level fluctuations through Jurassic times (e.g., Hardenbol et al., 1998) and accentuated during the Middle and Late Jurassic. In fact, plate tectonics had a major influence on relative changes in first and second order sea level fluctuations from the beginning of the break-up of Pangea (Hallam, 2001). The interpretation of third-order fluctuations is open to debate, as it is that of their precise duration. More locally, the evolution of the central North Atlantic Basin was of importance for the westernmost Tethys, especially throughout the Middle and Late Jurassic (e.g., Olóriz et al., 2002b and references therein; Rais et al., 2007). In low latitudes, the geodynamic evolution of western Tethyan areas during the Jurassic is witnessed in the context of irregular bottom physiography related to tectonic instability including submarine volcanism. This has been largely recognized for the epioceanic fringe in western Tethyan margins (e.g., D'Argenio, 1974; Olóriz, 2000). Therefore, in the context of intricate physiography and related hydrodynamics, approaching a spatial structuring of palaeo-water masses is a difficult task. Significant information on Late Jurassic sea water properties and related palaeoenvironmental dynamics can be gained by linking of the chemostratigraphic record preserved in marine carbonate archives with palaeoceanographic patterns. Among the latter, the mid-Oxfordian transgression was a major feature during early Late Jurassic times. This sealevel rise has been documented throughout the Tethyan realm and is widely recorded in carbon isotope records (Bartolini et al., 1999; Cecca et al., 2001; Rey and Delgado, 2002; Savary et al., 2003; Rais et al., 2007; Coimbra et al., 2009). Moreover, evidence for the mid-Oxfordian transgression is also found in the central North Atlantic Basin and the Mexico–Caribbean areas (Olóriz et al., 2002b; Cobiella-Reguera and Olóriz, 2009 and references therein). Increased continental weathering related to this palaeoceanographic– palaeogeographic event, coupled with increased nutrient runoff to the ocean, is reflected in seawater δ13CDIC values (Hoffman et al., 1991; Bartolini et al., 1996; Jenkyns, 1996; Bartolini et al., 1999; Colacicchi et al., 2000; Wierzbowski, 2002). Enhanced burial of sedimentary organic carbonate (12 C rich) results in a 13 C enriched oceanic DIC pool (Hoffman et al., 1991; Bartolini et al., 1996; Jenkyns, 1996; Weissert and Mohr, 1996; Bartolini et al., 1999; Colacicchi et al., 2000). The effects of the mid-Oxfordian transgression-related major oceanographic event on δ13C are widely debated for Tethyan and global interpretations (Bartolini et al., 1999; Cecca et al., 2001; Rey and Delgado, 2002; Savary et al., 2003; Rais et al., 2007; Coimbra et al., 2009; Dera et al., 2011). Later, during the Kimmeridgian and earliest Tithonian, global sealevel rise influenced major palaeoenvironmental parameters in lowto-lower intermediate latitudes. Superimposed higher frequency fluctuations in relative sea level were most probably driven by complex geodynamic and coupled ocean atmosphere interactions (Haq et al., 1988; Hardenbol et al., 1998; Hallam, 1999, 2001) or other regional factors. Expansion of carbonate platforms during second order sea-level highs resulted in higher carbon sequestration in shallow shelves (Weissert and Mohr, 1996). Corg/Ccarb ratio modulated the carbonate burial rate, resulting in a comparatively stabilized Kimmeridgian and Tithonian δ13C pattern, in comparison to Oxfordian δ13C fluctuations (e.g., Weissert and Mohr, 1996). This trend is observed along Tethyan palaeomargins (Weissert and Mohr, 1996; Bartolini et al., 1999; Cecca

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et al., 2001; Rey and Delgado, 2002; Savary et al., 2003; Rais et al., 2007; Coimbra et al., 2009). 1.3. The Ammonitico Rosso Facies Complex: advances in the last decades The Ammonitico Rosso Facies Complex (ARFC) comprises red to gray and clayey to calcareous, more or less nodular deposits in the centralwestern Tethys (see Fig. 1 for geographical distribution of outcrops reporting geochemical information). Numerous workers have studied the Ammonitico Rosso Facies and any approach to a brief historical review must be necessarily selective. Table A.1 (in Appendix 1) presents main items, contribution of geochemistry and related topics to the interpretation of ARFC deposits. More than 150 years of biostratigraphic work have provided essential support for any other research on the ARFC, and present-day intra-subzone ammonite-horizon resolution is commonly available (omitted bibliography refers to local research). A succinct image of the ARFC points to four main phases of research: (1) the 60–70s of the last century as the time for recognizing geodynamic contexts and nodulation processes, including pioneer comparisons with present-day examples. (2) The 80s showing progresses in detailed petrography favoring classification of AR types within the ARFC, which was recognized as resulting from multi-causal palaeoenvironmental forcing. Special steps forward related to increased recognition of the role of biogenic activity and a better understanding of diagenesis, as well as comparisons with present day nodulation processes. Innovative isotope analyses were applied very limitedly to ARFC deposits. (3) The 90s as crucial time recording a new focus on biogenic activity (pigmentation included), the wide application of isotope and elements geochemistry, a global scenario for deposition of the ARFC, the impact of sequence stratigraphy and ecostratigraphy and precise statements concerning diagenesis. (4) Finally, the first one and half decades of the present century showing the special relevance reached by magnetostratigraphy to improve stratigraphic correlation and chemostratigraphy to clarify diagenetic courses and palaeoenvironmental interpretations among others. Selective comments follow with citations, except for the ARFC research in the region covered by this study that is given throughout the main text — i.e., data about the ARFC in the Betic Cordillera will be limited to those focused on geochemical research. During the 1960s of the last century, relevant interpretations concerning the origin of ARFC focused on geodynamical contexts (e.g., Aubouin, 1964; Wendt, 1969, 1970) and below calcite compensation depth (CCD) dissolution for nodulation processes (Hollmann, 1962, 1964), but also shallow water eco-sedimentary contexts were revealed from the occurrence of microbiotic buildups (e.g., Sturani, 1964). During the seventies, accumulated knowledge on cementation and porosity processes favored diverse interpretations about nodulation processes. At the same time, Noble and Howells (1974) assumed rapid carbonate cementation at very shallow burial depths (some tens of centimeters) within calcite–aragonite-rich sediments, active burrowing and differential compaction for nodular limestones of Paleozoic age. Meanwhile, within a generally assumed geodynamic context (e.g., Jenkyns, 1971; Bernoulli and Jenkyns, 1974; Müller and Fabricius, 1974), the bioturbation–dissolution model proposed by Jenkyns (1974) for the origin of red nodular limestones was of particular relevance for avoiding interpretations that were strictly related to great-depth processes (e.g., Hollmann, 1962, 1964 “subsolution” model). Early comparisons with nodulation related processes in present day analogs from different depths were available (e.g., Bernoulli, 1972; Müller and Fabricius, 1974; Wendt, 1974; Larock and Ehrlich, 1975; Calvert and Price, 1977; Mullins et al., 1980), as well as cases of application of geochemical analysis to Mnnodules and crusts (e.g., Drittenbas, 1979) and valuable contributions to identification of micro-biogenic activity in the ARFC (e.g., Massari, 1979). Rapid, differential cementation early during diagenesis was proposed for nodulation (e.g., Bernoulli, 1972; Tucker, 1973, 1974; Schlager, 1974),

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Fig. 1. Geographic distribution of outcrops encasing Ammonitico Rosso deposits on which geochemical information is reported (see also Table 1 in Appendix 1). S Spain, Majorca Island (45 outcrops) and Italy (28 outcrops) after references in Table A1. Tunisia (11 outcrops) after Soussi et al. (1999); Turkey (9 outcrops) after Nicosia et al. (1991), Varol and Gökten (1994) and Varol and Tunay (1996) and Hungary (3 outcrops) after Cronan et al. (1991).

and there were alternative interpretations for cementation being derived from sea water (Kennedy and Garrison, 1975) or skeletal dissolution (Jenkyns, 1974). The state of the art in the early 80s was compiled by Farinacci and Elmi (1981) highlighting classifications and diverse interpretations of AR types of the ARFC resulting from the analysis of particular cases, emphasizing difference between slope versus variable swell, i.e., “seamount” conditions. Within the context of the geodynamic evolution in western Tethys, the local eco-sedimentary context was recognized to discriminate particular AR types within the ARFC, together with increasing recognition of not only macro- but also micro-biotic activity as one of the main factors determining this red to gray, marly to calcareous and more or less nodular facies complex. Hence, multi-causal forcing was recognized for the origin of ARFC deposits. Later in the eighties precise petrological descriptions and interpretations of nodulation and early diagenesis processes in particular examples of the ARFC progressed (e.g., Clari et al., 1984), as well as characterization of Fe–Mn precipitation, the palaeonvironmental meaning of Fe- and Mnnodules and their comparison with present ones (e.g., Grasselly and Polgári-Szentandrássy, 1985; Mindszenty et al., 1986). Isotopic analysis in ARFC deposits were in their initial phase (e.g., Jenkyns and Clayton, 1986; Weissert and Channell, 1989). During the late eighties and the early nineties, pioneer contributions highlighted the role of biogenic activity and the organic matter content, early during carbonate diagenesis, as well as in nodular limestones (e.g., Raiswell, 1987; Massari et al., 1989; Aller, 1990; Scudeler Bacelle and Nardi, 1991; Ballarini et al., 1994; Fenchel and Finley, 1995), thus promoting a new focus for analyzing relevant aspects in the genesis of ARFC deposits. In addition, bulk elemental geochemistry from Lower Jurassic Mn-nodules revealed similarity with modern Fe–Mn nodules and crusts, as well as typical difference between seamounts (ARCF deposits) and basinal settings (e.g., Cronan et al., 1991), and combined with the analysis of microbial activity was crucial for revealing conditions of Fe–Mn crust formation according to depth late in Jurassic ARFC contexts (Martín-Algarra and Sánchez-Navas, 1995; JiménezEspinosa et al., 1997). The combination of biotic, environmental and geodynamical factors were used to characterize AR-types within the ARCF, with no particular dependence of great depth (e.g., Nicosia et al., 1991). In addition, geodynamical and palaeoceanographical considerations served to Cecca et al. (1992) to propose a global scenario for interpreting causal palaeoenvironmental conditions during the

long time period in which the “death” ARFC occurred — Late Permian to earliest Cretaceous (ca. 112 My according to ISC (International Chronostratigraphic Chart, 2013). During the nineties, palaeogeograhical and tectono-eustatic interpretations, sequence stratigraphy analyses, ecostratigraphic interpretations and precise sedimentological observations enlarged the understanding from eco-sedimentary conditions to diagenesis in ARCF deposits (e.g., Martire, 1992; Olóriz et al., 1993, 1995, 1996; Clari and Martire, 1996; Martire, 1996; Caracuel et al., 1997; overviews and references in Olóriz et al., 2002a and Vera et al., 2004). Carbon and oxygen isotope analysis of skeletals supported geochemistry-based early palaeotemperature estimates in ARFC deposits (e.g., Price and Sellwood, 1994), as well as bulk C-isotopes related sea level fluctuations (e.g., Jenkyns and Clayton, 1986; Caracuel et al., 1997) within a context of increasing application of isotope analysis in Jurassic rocks (e.g., Weissert and Mohr, 1996; Veizer et al., 1997, 1999). Of particular relevance for interpreting relationships between palaeoenvironmental and depositional conditions was the similarity in Middle Jurassic carbon isotope curves from swells with intercalated ARFC deposits and troughs with occurrence of siliceous sedimentation in the Betic Cordillera, and those obtained from comparatively expanded and more or less siliceous hemipelagites in the Umbria–Marche– Sabina in the Apennines (e.g., Bartolini et al., 1998). Moreover, the role of microbial activity during the early diagenesis in red limestones throughout the Phanerozoic was well established (e.g., Mamet et al., 1997; Préat et al., 1999). From the earliest 21st century, chemo- and magnetostratigraphy and related analytic procedures, boosted new steps for better understanding ARFC deposits. Colacicchi et al. (2000) demonstrated complex interconnections among palaeoenvironmental parameters during Jurassic times forcing carbon isotope curves in siliceous deposits from Tethyan areas, the type area for ARFC deposits, as well as persistent decline in Upper Jurassic carbon isotopes previously identified by Weissert and Channell (1989) and later confirmed by Cecca et al. (2001) in ARFC deposits. Jenkyns et al. (2002) provided a general picture of Jurassic isotope and element chemostratigraphy. Rey and Delgado (2002, 2005) used comparison of C- and O-isotopes from epioceanic pelagic swells and troughs for interpreting difference in palaeoenvironmental and burial diagenesis conditions, as well as to recognize regional tectonics and global palaeoenvironmental events.

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Morano et al. (2003) identified microbially mediated paragenesis of calcite in AR. The incidence of microbial activity in differential pigmentation of ARFC deposits was firmly demonstrated by Mamet and Préat (2003, 2005, 2006) and Préat et al. (2004, 2006, 2008), as well as the microbial mediation in Fe–Mn crusts related to the ARFC sediments from different depths (e.g., Jiménez-Millán and Nieto, 2008; Reolid and Nieto, 2010). The impact of diagenesis on preservation of microbial– fungal structures in ARFC has been mentioned by Mamet and Préat (2006b). Carbon-isotope curves reflecting changes of the palaeocarbon cycle have been related to ammonite biostratigraphy, ARFC deposits included, showing main faunal turnovers and radiation episodes linked to changes in the carbon isotope record potentially forced by global fluctuations in the carbon-cycle (e.g., Sandoval et al., 2004; O'Dogherty et al., 2006). Préat et al. (2011) investigated the behavior of Fe and Mn in ARFC deposits and related major differences among AR types to geochemical composition and depositional setting, thus giving geochemical basis for lithofacies identification within ARFC deposits. Hence, these authors gave complementary support to the analogous AR types' differentiation proposed in the earliest 80s of the twenty century by Elmi (in Farinacci and Elmi, 1981), and compared with younger analogous such as the Cretaceous Oceanic Red Beds (CORB) of Late Cretaceous age. In a different

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way, geochemical analyses across a transect of several hundred kilometers in the S–SE palaeomargin of the Iberian Plate provided new insights for the understanding of a large scope of topics from diagenetic courses to palaeoenvironmental interpretations, with mention of local and global controlling factors epicontinental and epioceanic water masses, including the identification of episodes of hydrothermal influences and exportation of Sr-rich fine clastics affecting areas characterized by ARFC deposits (e.g., Coimbra et al., 2009; Coimbra and Olóriz, 2012a,c; Coimbra et al., 2014a,b). In addition, taphonomic and magnetostratigraphic analyses conducted on ARFC deposits refined stratigraphic interpretations and correlations from local to “global” scale, with recognition of a potential global warming during the Late Jurassic and the influence of vital effects on C-isotope records retrieved from belemnite rostra (e.g., Caracuel et al., 2000; Pavia et al., 2004; Pruner et al., 2010; Žák et al., 2011). 1.4. Testing spatial geochemistry across a Late Jurassic palaeomargin — the case study In an attempt to test Upper Jurassic micritic carbonates as archives of the complex three-dimensional organization of palaeo-water masses across a Tethyan palaeomargin, sections from the Iberian sub-plate

Fig. 2. Regional distribution of geological units along the studied areas: (A) Southern Iberia. Note present day distance among sections, not considering palinspastic reconstructions. (B) Majorca Island (adapted from García-Hernández et al., 1980; Caracuel and Olóriz, 1998). Stars indicate locations of studied sections (note key to abbreviations for sections studied). (C) Main oceanographic, bathymetric and sedimentological parameters of studied sections with indication of sources used. Upper Jurassic ammonite biochronozones registered throughout the studied sections include: Antecedens, Transversarium, Bifurcatus, Bimammatum, Planula (Middle to Upper Oxfordian); Platynota, Strombecki, Divisum Compsum, Beckeri (epioceanic Kimmeridgian); Platynota, Hypselocyclum, Divisum (epicontinental Lower Kimmeridgian); Hybonotum, Albertinum, Semiforme/Verruciferum, Richteri, Admirandum/ Biruncinatum, Burckhardticeras (Tithonian).

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(Spain and Portugal) were investigated for their sedimentological, stratigraphic and geochemical records. As a pre-requisite for spatial characterization of a mixed carbonate-siliciclastic platform and adjacent seaward deposits, this study relies on a well-established biochronostratigraphic control based on ammonites (Olóriz, 1978; Marques, 1983; Marques and Olóriz, 1989a,b; Cariou et al., 1997; Geyssant, 1997; Hantzpergue et al., 1997; Caracuel and Olóriz, 1998; Olóriz et al., 2002a). This allowed for the correlation of separate sections across depositional environments and lithofacies (see caption of Fig. 2 for identified ammonite biozones). The aims of this paper are to: (1) present a detailed Late Jurassic geochemical record from seven spatially distant stratigraphic sections in Southern Iberia along a several-hundred-kilometer long, proximal-todistal transect connecting two major marine palaeoenvironments — the neritic/epicontinental and the epioceanic domain; (2) interpret the geochemical trends retrieved from carbonate materials in terms of transect-wide spatial and temporal variations of palaeo-seawater properties taking into account local and diagenetic (noise) features; and (3) compare and contrast these data against a review of published coeval records. With this, we aim at encouraging research focused on the recognition of spatial chemostratigraphic patterns and their relationships to palaeoceanographic conditions. 2. Regional setting and study areas 2.1. Regional setting Data come from carbonate samples collected in Southern Iberia (Spain and Portugal) and the Majorca Island (Fig. 2A). For details in the regional geology and stratigraphy see Olóriz et al. (2002a) and Vera et al. (2004 and references therein). During the Late Jurassic, the

southeastern Iberian margin was a tectonically unstable area. It was affected by movements between the Iberian and African plates (Fig. 3A) along the Maghrebian–Gibraltar Transform Zone, an extensional–transtensional margin linked to the geodynamic history of the Central-North Atlantic Basin and the westernmost Tethys. In the northwest Tethyan margin (Figs. 3A and 4), the physiography of southern Iberia resulted in differentiation of two major palaeogeographic areas. The more proximal region, within a wide epicontinental shelf system, is represented by the Algarvian platform and the Prebetic Zone and northeast equivalents, while the more distal epioceanic fringe is represented by the Subbetic Zone and lateral equivalents (Figs. 2, 3A and 4). The Balearic Archipelago (Figs. 2A, B and 4) constitutes the northeastern extension of the Betic Cordillera (the Betic–Balear Domain in Fontboté et al., 1990). Pre-Cenozoic palaeogeography and Jurassic stratigraphy suggest that the Majorca Island formed part of a northeastern segment in the East-Iberian palaeomargin during Jurassic times (see Olóriz et al., 2002a for references). Sections on Majorca belong to the Sierra Norte domain, with detailed regional geology reported by Álvaro et al. (1984) and detailed Upper Jurassic stratigraphy described by Caracuel and Olóriz (1998). The data presented here, including bio-chronostratigraphic ranges, thicknesses, and main lithofacies (see Fig. 2) result from the analysis of strata along a proximal-to-distal transect. Ammonite biochronostratigraphy allows correlations from the middle shelf to the epioceanic fringe, ranging from ca. 60 to 250 m water depth according to Olóriz et al. (2003) and Olóriz and Villaseñor (2010) (Fig. 2). 2.2. Environmental setting The most proximal section, Rocha Poço (Figs. 2A, C, 3 and 4), includes Lower Kimmeridgian epicontinental silty limestones and sponge

Fig. 3. (A) Late Jurassic paleogeographic reconstruction of western and central Tethys. Schematic location of sections is indicated by stars and labeled 1 through 7. (B) Lithological columns of the studied sections. Biochronostratigraphy as in Fig. 1 caption: OXF (Mid-Upper Oxfordian); KIMM (Kimmeridgian); TITH (Tithonian); BERR (Berriasian). Plate tectonic setting is from Stampfli and Borel (2002). Depositional environments are from Thierry et al. (2000a,b).

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Fig. 4. Composite schematic representation of the south and eastern Iberian paleomargins showing assumed location of the studied sections (1 to 7, as in Fig. 2) and major paleoenvironmental parameters during deposition (adapted from Olóriz, 2000). Note that the most proximal end-member of the studied transect is represented by the Rocha Poço section (1) (southern Portugal; Figs. 1A, 2 and 3), covering mainly Kimmeridgian deposits (Marques, 1983; Marques and Olóriz, 1989a,b; Olóriz et al., 1994) from the epicontinental shelf system represented eastwards by the wider Prebetic Shelf. The Cala Fornells section (5) holds a similar, albeit somewhat more distal position compared to the Cuber section (4) in tectonic-sheet IV of Álvaro et al. (1984) according to differential compression and shortening (Gelabert et al., 1992) in the piggy-back thrust-system of the Sierra Norte or Serra de Tramuntana (Majorca island). The most distal sections are those of Salcedo (6) and Cardador (7) belonging to the Internal Subbetic Zone (Figs. 1A, 2 and 3).

bioherms (Marques, 1985; Ramalho, 1988), a biofacies that suggests palaeo-water depths ranging from 25 to about 200 m (Keupp et al., 1990; Olóriz et al., 2003 and references therein). Ramalho (1988) estimated the depth for the euphotic zone in the range of 50–150 m for the sponge reefs in the Rocha area. Recent work on Jurassic coral and sponge–microbialite buildups estimate depth ranges of 60–100 m (Olivier et al., 2007). Eastwards along the south Iberian palaeomargin, mid-shelf depths (50–70 m) are assumed for Oxfordian spongemicrobial buildups in the Prebetic Zone (Olóriz et al., 2003). Thus, the shallowest depths estimated for the most proximal section would correspond to a palaeodepth of 50 to 60 m, but no precise information about photic zone depth is available. The epioceanic record is dominated by nodular limestones recorded on structural highs. Averaged low sedimentary rates favored the deposition of carbonate mud, resulting in condensed, nodular Ammonitico Rosso and related facies on marine highs (Fig. 4), at maximum water depths of around 250 m (Jenkyns, 1974; Olóriz et al., 1996, 1997; Olóriz, 2000). Synsedimentary nodulation resulted from interplay among sediment input, current sweep and winnowing, and burrowing that enhanced pore-water circulation (Flügel, 2004; Coimbra et al., 2009). Nodular structure formed early during diagenesis (Jenkyns, 1974; Müller and Fabricius, 1974; Mullins et al., 1980; Clari et al., 1984; Clari and Martire, 1996) resulting in an early stabilization of carbonate minerals (Jenkyns, 1971). Often, active burrowing triggered subsurface dissolution of aragonite skeletal grains, mainly ammonites, gastropods, coccolithophorids, and “filaments” corresponding to valves of young pelagic bivalves, and favored the irrigation of taphonomically active firmground zone. This contributed to the availability of dissolved carbonate and elements for nodule formation early during diagenesis

(Jenkyns, 1974; Farinacci and Elmi, 1981 and references therein; Coimbra et al., 2009). Both sections from the Island of Majorca (Figs. 2B, 3 and 4) are distinct in lithofacies. At Cuber and Cala Fornells, Ammonitico Rosso facies show thin intercalations of wavy to platy limestone, partly bearing silica nodules, suggesting a locally higher silica input, perhaps from dissolution of marine siliceous skeletons, radiolaria rather than sponge spicules. The only deviation from epioceanic Ammonitico Rosso facies is the Aumedrá Fm., which has been interpreted as resulting from fine sediment sorting during transport from a distant shallow carbonate shelf (Coimbra and Olóriz, 2012c). The sedimentary archive of the epioceanic fringe points to a general picture of Late Jurassic carbonate deposition on west-Tethyan epioceanic swells resulting in the condensed Ammonitico Rosso and related facies (for a comprehensive review see Farinacci and Elmi, 1981). Precursor sediments of Ammonitico Rosso and related facies were an admixture of suspended mud and “planktic snow”, including the possibility for local variable frequency of episodes of fine sediment winnowing and reworking. Among preserved macroscopic skeletals, those of vagrant organisms (cephalopods) overwhelmingly dominate over scarce benthos in our record, but local deviation to benthos-rich settings are known elsewhere. 3. Methods and materials 3.1. Sampling strategy and analytic procedure A total of 316 hand specimens averaged 19 samples per meter of section for the more condensed profiles, i.e., the Cañada del Hornillo,

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Cardador, Puerto Escaño and Salcedo sections, and one sample per meter of section for the Cuber, Cala Fornells and Rocha Poço sections (Fig. 3B). Some of these samples (40%) were previously presented in a different context (Coimbra et al., 2009 and Coimbra and Olóriz, 2012c). Here, we combine published evidence with new data and provide a review of resulting patterns. From each hand sample, two slabs were cut: one for macro- and microscopic inspection of the microfacies and cathodoluminescence (CL) patterns and a twin slab for geochemical analysis. Photomicrographs of thin sections from each hand specimen provided a quantitative estimate of skeletal abundance through pixel counting (Coimbra and Olóriz, 2012b). CL analysis of selected samples of each carbonate material used a hot stage CL microscope (HC4-LM) at the facilities of the Institute for Geology, Mineralogy and Geophysics, Ruhr University Bochum, Germany. Matrix micrite powder sub-samples were hand-drilled using diamond drill tips with diameters of 1 to 1.4 mm. An average of four powder samples per slab resulted in a total of 1150 matrix micrite sub-samples, but also included veinlet carbonate cements (n = 102), ammonite shells (n = 2), and belemnite rostra (n = 30). At the Rocha

Poço section, sedimentary infills of cavities in sponge bioherms (Fig. 5E and beige areas in Fig. 6D) were sampled systematically. Matrix micrite encasing silica nodules in the Cuber section (Figs. 2B, 3 and 4) was tested for the potential influence of silica diagenesis on carbonate geochemistry. Three transects crossing nodule areas were sampled, resulting on a total of 51 matrix micrite samples. Geochemical analyses were performed at the facilities of the Institute for Geology, Mineralogy and Geophysics, Ruhr University Bochum, Germany using the same equipment and methods described in Coimbra et al. (2009) and Coimbra and Olóriz (2012c). Analytical precision (±1σ) for carbon and oxygen-isotope data, controlled by NBS19 and internal standards, was better than ±0.03 and ±0.07‰ for δ13C and δ18O, respectively. Duplicate samples presented a deviation of ± 0.02‰ for δ13C and ± 0.08‰ for δ18O. The palaeotemperature equation of Anderson and Arthur (1983) was applied, considering a δ18Oseawater between −1.0 and + 0.5‰ SMOW according to Shackleton and Kennett (1975), hence, providing palaeotemperature ranges rather than single palaeotemperature values. All reported palaeotemperatures follow this approach, supported by the variability in values assumed for

Fig. 5. Field images documenting sampling strategy and lithologic features. (A) Typical Ammonitico Rosso succession of the Salcedo section with numbers in white labels as an example of the sampling strategy (resolution) applied during field work (white labels in A). Hammer for scale, white arrow. (B) Calcareous Ammonitico Rosso horizon (Cardador section). (C) Detail of nodular structure (Cardador section). (D) Siliceous nodules intercalated in Ammonitico Rosso succession (Cuber section, siliceous nodules indicated by white dashed line). (E) Spongiolithic bioherms (Rocha Poço section).

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Fig. 6. Polished slabs showing macroscopic features. (A) Indistinct nodularity in Ammonitico Rosso facies with macrofossils (Salcedo section). (B) Ammonitico Rosso nodule and internodule (arrow) facies with characteristic red color pattern (Cardador section). (C) Siliceous nodule in wackestones matrix crosscut by a vein filled with late burial cement (Cuber section). (D) Slab from spongiolithic bioherms (Rocha Poço section). Arrow indicating “infill” refers to fine-grained sediment infilling primary cavity in intra-spongiolitic space. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Jurassic δ18Oseawater as shown in Price and Sellwood (1994), Price et al. (1998), Lécuyer and Bucher (2006), Brigaud et al. (2008), Wierzbowski and Joachimski (2009), and Price and Harwood (2011). Maximum elemental scatter for duplicate samples was in the order of 3%. The elemental dataset is the core of research in progress and will be briefly mentioned whenever elemental information provides a better understanding of the isotope dataset.

and Olóriz, 2012a). Here only two cases of ammonite shell preservation (n = 2) were available, in which cementation patches were identified.

3.2. Carbonate and silica materials

4.1.1. Carbonate grains and textures Selected photomicrograph images illustrating relevant features described in this subchapter are given in Appendix 2 (Figs. A.1, A.2 and A.3). This detailed information provides visual support for composition, fabrics and interpreted sedimentary processes discussed in the following text, which results from an extensive analysis of hundreds of thin sections produced for each sampled horizon. Typically, Ammonitico Rosso facies in Cardador, Salcedo, Cañada del Hornillo, Puerto Escaño, Cuber and Cala Fornells sections (Figs. 2 through 4) are predominantly wackestones with variable contents of radiolaria; calcisphaerulids; calcareous dinoflagellate cysts; planktic crinoids represented by mainly Kimmeridgian and Lower Tithonian Saccoccoma; Oxfordian Conoglobigerinidae, Globuligerina, and other unidentified planktic and benthic foraminifera; Upper Tithonian microgranular-organic and mainly hyaline tintinnoids; ostracods; green algae zoospores, Globochaetes; cephalopods, belemnites and ammonite remains including ammonitella, shell fragments and aptychi; benthic molluscs mainly represented by fragments of gastropods and bivalves; brachiopods; plates, spines and fragments of echinoderms; and sponge spicules. Locally, packstone horizons are mainly composed of pelagic bivalves, so-called “filaments” and/or Sacoccoma sp., planktic foraminifera such as Conoglobigerinidae, and radiolaria, with a variable amount of peloids. Aside from pebbly mudstone horizons and hardgrounds, background microfacies containing variable amounts of the microfossils

Matrix micrite samples discussed here are dominated by nodular Ammonitico Rosso facies (Figs. 5A, C and 6A, B) alternating with stratigraphic horizons characterized by less distinct to indistinct nodulation (Figs. 2C and 3B for facies distribution, and Fig. 5B). The occurrence of facies with palaeoenvironmental relevance other than Ammonitico Rosso adds complexity to this study. These include silica nodule bearing horizons at the Cuber section (Figs. 5D and 6C) and silty limestones and bioconstructed carbonate deposits (Figs. 5E and 6D) at the Rocha Poço section. Overall, cemented veinlets are uncommon in the carbonate slabs, but blocky calcite cements and microsparite patches were identified and sampled where possible. The sampling procedure consisted on retrieving bulk carbonate cement values, since the aim was to assess the degree of diagenetic imprint, rather than the characterization of each diagenetic stage. Cementation in skeletal remains is considered separately since they constitute a particular case. All belemnite rostra (n = 30) were screened for diagenetic alteration using elemental composition obtained from bulk samples of the calcitic rostra and catodoluminescence (Fig. 7H, J). Inner cast preservation of ammonites in Ammonitico Rosso indicates dissolution of aragonite, which agrees with metastability of aragonite under marine conditions (Veizer, 1983). Notwithstanding, neomorphic ammonite shells are known in red-gray cephalopod-rich pelagic limestones that are distinct from typical Ammonitico Rosso facies (Coimbra

4. Petrographic and geochemical data 4.1. Microfacies analysis

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Fig. 8. Carbon and oxygen isotopic composition of matrix micrites, carbonate cements (veinlets), and belemnite rostra from all sampled locations. Mean values are presented for each analyzed studied material and section (vertical lines), as well as maximum and minimum values (width of the bars). The number of samples analyzed is indicated in each case (white numerals). Note the overall narrow δ13C range (except belemnites) and fairly homogeneous δ18O ratios (except carbonate cements in veinlets).

mentioned show fabrics indicating common reworking and hence no preservation of microstratigraphic order between micro-erosion/ omission surfaces — i.e., uncommon macroscopic recognition of sedimentary laminae. These microfacies correspond to usually burrowed soft- to firmgrounds. The spatial arrangement of elongated skeletals such as “filaments” and Sacoccoma is not in agreement with deposition surfaces, and these skeletals usually show variable size and orientation, while size gradation in other microfossils such as foraminifera, radiolaria, and skeletal fragments is not evident. Sorting is variable but rather low showing cases of millimeter-size gastropods embedded in Globuligerina-rich wackestone, as well as particle sorting in horizons with closely packed skeletals. Sedimentary packing varies and allows differentiation of horizons with loose and close packing. Matrixsupported, wackestones correspond to heterometric and more heterogeneous bioclastic accumulations with particle sizes mainly b500 μm and more frequently in the range of medium to very fine sand, and lower fragmentation in wackestones with variable content of skeletals but uncrushed spherical particles up to 500 μm. Packstones usually are dense accumulations of a given skeletal type as it is the case of grain supported “filaments” and/or sacoccoma-rich horizons, or of more heterogeneous skeletals showing local nesting and less common stacking in the case of more densely packed horizons. Fragmentation is highly variable above 500 μm but commonly intact shell's geometries are evidenced by shell curvature in bivalves and other mollusks, with or without punctual breakage. Additional cases are ammonite camerae with partially recrystallized or matrix in-filled interiors between 1 and

3 mm and up to 4 mm embedded in wackestone and mudstone matrixes. Cases of accumulation of silty sized skeletal fragments are known. In contrast, skeletal telescopage is a rarity if recorded. All these features, together with macroscopic occurrence of trace fossils (Chondrites, Planolites, Thalassinoides) agree with usual burrowing as also revealed by cloudy aspect of the matrix in polished surfaces and less commonly in filled tubes. Carbonate dissolution/precipitation processes are widespread and highly variable, but preservation of subtle shell structure in small, originally aragonitic fossils is known (e.g., ammonitella). Less common are fabrics showing well preserved micro-horizons, which typically are rich in “filaments”, radiolarian, pellets and/or mudstone. Micro-erosional surfaces are identified, as well as differential infill of shell cavities with respect to the enclosing matrix and the occurrence soft, partly lithified, intraclasts. Microfacies from the more proximal Rocha Poço section clearly differ from those described above for typical epioceanic sections, and correspond to a substantially thicker section of Lower to lowermost Upper Kimmeridgian deposits. This epicontinental setting shows distinct biogenically-induced sedimentary structures with peculiar fabrics (Fig. 6D). A silty (quartz) marly–clayey limestone (mudstone to wackestone) rythmite rich in ammonites, bivalves and plant remains characterizes the lower 20 m of the Rocha Poço section. Silty to marly intervals include disarticulated ostracods and scarce variably preserved small benthic foraminifera. The overlying 25 m thick spongiolitic facies belongs to the “biohermes stromatolitiques à spongiaires silicieux” of Ramalho (1988) within the Peral Formation (Marques, 1985). The

Fig. 7. Comparison of transmitted light (images to the left) and CL imaging (to the right) of characteristic features. (A/B) Typical epioceanic cement paragenesis. Cement generations are labeled 1 to 3. Matrix micrite presents dull luminescence. (C/D) Carbonate cement from the upper part (spongiolithic horizon) of the epicontinental section. Note the presence of one single cement phase. (E/F) Ammonite shell showing a complex paragenetic infill. (G/H) Belemnite rostrum. Note overall intrinsic blue luminescence, cross-cut by bright orange veinlet. Rare interlaminae orange luminescence is present. (I/J) Belemnite rostrum. Belemnite rostrum shows similar luminescence pattern as (D). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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Fig. 9. Spatial chemostratigraphy: (A) carbon isotope composition and (B) oxygen isotope stratigraphy of all studied sections (Cardador, Salcedo and Cañada del Hornillo sections after Coimbra et al., 2009 and Cuber and Cala Fornells section as in Coimbra and Olóriz, 2012c, not previously discussed). Intra-sample isotopic variability is indicated for each investigated horizon. Full line indicates the mean value of all subsamples for each hand specimen (carbon and oxygen in all sections) and discontinuous line (only in the Cuber section) connects veinlet carbonate cement samples. Key to symbols for different analyzed materials is given. The portion of the sections that represents the Kimmeridgian time interval is represented by horizontal dashed lines. Lithology, biochronostratigraphy and abbreviations as in Figs. 1 and 2.

spongiolitic limestones are mainly composed by peloidal wackestone, and minor packstone, with well-preserved microbial/algal structures corresponding to planar-to-wavy laminated or stromatolitic microbialites, and thrombolites. Subrounded rather than angular micritic intraclasts and micritized aggregate grains are common. Small, monocrystalline quartz grains occur, mainly at the bottom and top of the section. Microbial encrustation of skeletals is widespread, and clotted peloidal fabrics occur locally. Common microbial/algal oncoids and lumps include skeletals, and are usually embedded in brown micrite matrix or pale-gray microspar patches most probably related to infilling/recrystallization of voids. Very fine mudstone, automicrite and/or allomicrite, to microspar infill of tubes related to boring and/or burrowing locally interrupting microbialite layering. In addition to patches of microspar, larger spaces most probably due to biogenic activity show coarser particulate sediment, including rare intraclasts derived from the enclosing wackestone. Uncommon geopetals show pale-brown fillings containing rounded intraclasts, skeletals and pellets. Silicification is common above levels with large sponges. Skeletal abundance was estimated semi-quantitatively for all epioceanic sections and revealed fluctuations in bioclast abundance between 1 and 45%, with a mean value of 15%. The epicontinental

record was not included in this quantification because the lower part of the Rocha Poço is rich in quartz but poor in calcareous bioclasts and the overlying spongiolithic facies limits abundance estimations. In spongiolithic limestones, the matrix is composed largely of automicrite sensu Neuweiler and Reitner (1993) and Reitner and Neuweiler (1995) and the precise contribution of allomicrite cannot be established.

4.1.2. CL patterns CL microscopy of matrix micrite revealed dull to slightly brighter luminescence (Fig. 7B, D, F and H). In general, carbonate cements from veinlets gathered from the epioceanic samples revealed three cement generations (1 to 3 in Fig. 7B). The first generation (1) is a nonluminescent, subhedral calcite phase, followed by a bright-luminescent orange generation (2). The latest cement generation (3) is a dull orange luminescent blocky calcite, which occluded the center of voids. This pattern was present in all epioceanic samples investigated. In contrast, epicontinental samples yielded only dull blocky calcite cements (see Fig. 7D) with no evidence for a more complex succession of diagenetic stages.

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Two ammonite shells retrieved from two epioceanic sections locally showed a three-phased luminescence pattern (Fig. 7F), but differences regarding epioceanic veinlet carbonate cements deserve attention. In these ammonite shells, CL phase 1 accounts for 23 to 40% of thin section area, while phases 2 and 3 represent 60 to 77%, respectively. Belemnite rostra show an overall intrinsic dark blue luminescence with the exception of rare, orange luminescent growth increments that are crosscut by bright orange veinlets (Fig. 7H, J).

4.2. Carbonate geochemistry 4.2.1. Carbon and oxygen isotope ratios Carbon and oxygen isotope data are shown in overview in Table 1 and in Figs. 8 and 9. All epioceanic sections (Cañada del Hornillo, Puerto Escaño, Cuber, Cala Fornells, Salcedo and Cardador) share similar C-isotope stratigraphic trends (Fig. 9). Although absolute values of δ13C ratios fluctuate between sections, they range from 2.0 to 2.5‰ at the base of the recorded Oxfordian, with a pronounced positive shift of about 1‰ during the Middle Oxfordian. With the exception of the Aumedrá Formation of Majorca Island (Coimbra and Olóriz, 2012c), carbon isotope ratios progressively decrease in Kimmeridgian and Tithonian strata. Late Jurassic minima of ca. 1‰ are reached for the topmost Tithonian. The most proximal section, Rocha Poço (Figs. 2 through 4) is characterized by depleted δ13C values near the base of the Kimmeridgian. Further up-section, isotope values progressively increase from −6.0 to + 2.0‰, with the most pronounced shift across the boundary to spongiolithic limestone (Fig. 9A). At the base of the spongiolithic facies, the δ13C record of infilling material differs from the biologically mediated sediments, showing depleted values, a distinction no longer recorder in overlying horizons. Thus, positive δ13C ratios are maintained throughout the Kimmeridgian spongiolithic facies except for a single, punctual negative deviation of about 6.0‰ (Fig. 9A). An overview of the δ18O characteristics of different carbonate materials (Table 1) reveals matrix micrite samples with mean values around −1.3‰, whereas veinlet carbonate cements are the most depleted set of samples, with average of −3.7‰ (Fig. 8 and Table 1). Belemnite δ18O values are higher than the average of matrix micrite δ18O values.

Fig. 10. Crossplot including major diagenetic domains identified from mean isotope values for the studied carbonates. Epioceanic matrix micrite samples fall within the marine domain while veinlet carbonate cements agree with progressively more depleted δ18O ratios typical for the shallow marine burial domain. Data from epicontinental samples reveal a more complex pattern due to the interference of Tethyan with local paleoceanographic patterns and a differential diagenetic pathway including a possible meteoric diagenetic overprint.

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Oxygen isotope values resulting from ammonite shells differ from each other (−1.2 against −3.4‰). The geochemical signature from different carbonate materials (Figs. 8 and 9) shows sections from the epioceanic fringe grouped based on their matrix oxygen isotope values. More landward sections (Cañada del Hornillo and Puerto Escaño; Figs. 2A and 4) display similar values between −1.9 and 0.5‰. The Cala Fornells and Cuber sections from the Majorca Island (Figs. 2B, C, 3 and 4) include the same average matrix micrite δ18O values (− 2.0‰) and similar maximum values (− 0.2‰), although the Cuber section reaches lower values (− 6.2‰) than the Cala Fornells section (−3.7‰). Matrix micrite from the more distal, epioceanic sections of Cardador and Salcedo share the same range of matrix micrite δ18O values (− 0.3 to 0.9‰, mean of 0.2‰; Fig. 8). The most proximal epicontinental section at Rocha Poço is characterized by low mean matrix micrite δ18O values (− 4.9 to −0.4‰).

Fig. 11. CL images from ammonite shells in relation to their C and O isotope composition. (A) and (B) CL images obtained from two ammonite shells from Puerto Escaño and Cañada del Hornillo sections, respectively. Paragenetic phases labeled 1 to 3 (see also Fig. 6F and text for main differences regarding cement generations 1 to 3 recognized in later carbonate cement phases). (C) and (D) Same images applying pixel counting. Phase 1 was eliminated, allowing a better view of carbonate cement phases 2 and 3. Dashed line outlines the shape of the shell. (E) Carbon and oxygen isotope composition for bulk ammonite shells (phases 1 to 3), host matrix micrites and mean carbonate cement values from later diagenetic veinlets. Note different significance of phase 1 in each example, and differences in oxygen isotope composition of ammonite shells. Data obtained from Puerto Escaño materials are shown in green and those from Cañada del Hornillo are shown in red. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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Carbonate cements from veinlets can be separated in two groups based on their oxygen isotope signature (Fig. 8). One group includes δ18O values lower than those of matrix micrite (Cardador, Salcedo, Cañada del Hornillo, Puerto Escaño and Cala Fornells sections; Figs. 2 through 4) and a second group with carbonate cement oxygen isotopes that partly overlap with matrix micrite data (Cuber and Rocha Poço sections; Figs. 2 through 4). Oxygen isotope values for belemnites show minor fluctuations, ranging from − 1.7 to − 0.2‰ (Fig. 8). Most belemnite samples plot within the more positive end of the range of matrix micrite data. Matrix micrite δ18O isotope values of all sections (Fig. 9B) reveal an increase in the order of 1‰ during the mid-Oxfordian in all epioceanic sections (Puerto Escaño, Cañada del Hornillo, Cala Fornells, Cuber, Salcedo and Cardador). The Kimmeridgian and Tithonian record is more variable, but a first-order isotope pattern is recognizable in all epioceanic sections (Fig. 9B). After the mid-Oxfordian positive shift, δ18O ratios decrease towards the Oxfordian/Kimmeridgian boundary. Kimmeridgian matrix micrite δ18O values increase, and the Tithonian record reflects the return to δ18O-depleted values that shift to more positive values towards the end of the Late Jurassic. Overall, a marked negative shift in δ18O values coincides with silica nodule-bearing horizons (Fig. 9B, Cuber section). Micrite and veinlet carbonate cements follow the same stratigraphic trend, although the latter are shifted towards more negative values (about 4‰ lower than micrite values). Samples chosen to test the potential influence of silicification on the matrix micrite surrounding silica nodules show no specific variation, with only minor random fluctuations. In the epicontinental Rocha Poço section (Fig. 9B), the Early Kimmeridgian record is rather stable (maintaining values of ca. − 4‰; Fig. 9B), with a gradual increase towards δ18O-enriched average values of − 2.7‰ upwards in Lower Kimmeridgian strata. 4.2.2. Elemental geochemistry Although an in-depth interpretation of elemental data for different carbonate materials is beyond the aims of this study, elemental evidence provides insight into the degree of diagenetic alteration, such as the preservation of belemnite rostra. Elemental threshold values separating well preserved and altered belemnite rostra have been proposed (Voigt et al., 2003; Rosales et al., 2004; Nunn et al., 2009; Malkoč and Mutterlose, 2010; Price, 2010), but no consensus values are reported. In general, relatively low Mn and Fe contents, along with high Sr

Fig. 13. Stratigraphic distribution of δ18O and Mn abundances from epioceanic Cuber section. Calcium stratigraphic abundance and Mn/Ca are also shown in order to discard this proxy as a cause for fluctuations in Mn abundance. Note the inverse relationship between δ18O and Mn, with marked shifts across horizons bearing siliceous nodules. The brownish wavy limestones refer to the Aumedrá Formation. Biochronostratigraphy as in Fig. 1, caption. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

abundance agree with a fair preservation state of belemnite rostra (Veizer, 1983; Saelen, 1989; van de Schootbrugge et al., 2000; Nieburh and Joachimski, 2002; Malkoč et al., 2010; Benito and Reolid, 2012).

Fig. 12. Small-scale geochemical transect across matrix micrite encasing silica nodules. (A) to (C) Isotope values, with black numerals for δ13C and white for δ18O, with location of sampling area. (D) Summary of the obtained values, with indication of number of samples (n), mean values and relative variation of carbon and oxygen isotope composition (Δδ) along each transect.

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diagenetic history (Dickson and Coleman, 1980; Allan and Matthews, 1982; Melim and Scholle, 1992; Bickert, 2000; van der Kooij et al., 2009; Madden and Wilson, 2013). Carbon-isotope ratios reflect balance changes among components of the carbon cycle (Berger and Vincent, 1986; de Boer, 1986; Walter et al., 2007; Derry, 2010). Dissolution of metastable carbonate mineralogies, neomorphism and latter recrystallization processes may alter the original marine δ13 C composition (Given and Lohmann, 1985; Carpenter and Lohmann, 1989, 1997). The source of diagenetic carbonate precipitated is often the host carbonate itself (Hudson, 1975) and therefore, this proxy is rather conservative. Pervasive diagenetic processes can be reflected by depleted carbon isotope values, since late diagenetic fluids are commonly 13C depleted relative to matrix carbonates (e.g., Carpenter and Lohmann, 1997; Schneider et al., 2008; van der Kooij et al., 2009; Brand et al., 2010; Geske et al., 2012). Oxygen isotope composition of marine carbonate materials largely reflects changes in seawater temperature and/or salinity (Craig and Gordon, 1965; Marshall, 1992; Spero et al., 1997; Kohn and Welker, 2005; Chacko and Deines, 2008; Suarez et al., 2010), as well as sedimentary porosity and permeability. Other factors may cause oxygen isotopic fractionation, such as mineralogy or the biotic versus abiotic origin of different carbonate materials (see Marshall, 1992 for details). Oxygenated pore waters buffers δ18O composition of later carbonate precipitates (Allan and Mathews, 1977; Lohmann, 1987). Post depositional changes are usually accompanied by increased fluid temperature in the burial realm or the influence of 18O depleted meteoric fluids (Given and Lohmann, 1985; Hoefs, 1997), causing original δ18O values to be significantly lowered (Plunkett, 1997; Bartolini et al., 2003; van der Kooij et al., 2009; Madden and Wilson, 2013).

Fig. 14. Skeletal abundance plotted against oxygen isotope chemostratigraphy. (A) Statistically relevant, linear correlation computed for data from all epioceanic sections. Note that variable fabrics characterize epioceanic Ammonitico Rosso facies. The resulting heterogeneous lithofacies comes from a wide array of local depositional conditions where skeletal abundance is not conducted by a unique sedimentation/maturation process. (B) Stratigraphic distribution of skeletal abundance (dashed line) and oxygen isotope composition (full line) for the Cardador and Cala Fornells sections (biochronostratigraphy as in Fig. 1, caption). Note co-variance of both proxies.

From 30 analyzed samples, 19 fall within ranges of the published elemental composition of “well-preserved” belemnite low-Mg calcite (according to references above; Tables 2 and 3). It was also taken into account that their respective oxygen isotope values was not conspicuously low and that no clear evidence for diagenetic alteration was observed under CL. Manganese content at the Cuber section was compared with oxygen isotope composition, revealing opposite trends between these proxies, further discussed. 5. Interpretation and discussion 5.1. Preservation and reliability of C and O records Carbon and oxygen isotope ratios in Mesozoic marine carbonates may present a wide range of values, depending on primary parameters (including carbonate mineralogy, seawater physico-chemical properties, metabolic effects and precipitation kinetics) and their subsequent

5.1.1. Carbon isotope composition Values for most of the analyzed carbonate materials (except belemnite rostra, matrix micrite from the lower part of the Rocha Poço section, and respective veinlet carbonate cements; Fig. 8) oscillate around δ13C ratios of 2.0‰. These values are in agreement with previous data from Upper Jurassic Tethyan margins (Joachimski, 1994; Jenkyns, 1996; Weissert and Mohr, 1996; Bartolini et al., 1999; Rais et al., 2007; Coimbra et al., 2009; Boulila et al., 2010). Depleted δ13C ratios (b2.0‰) from the lower part of the Rocha Poço section (Fig. 9A) relate to more silty lithofacies. Carbon isotope data from matrix micrite and veinlet carbonate cements from these horizons are depleted as much as 8.0‰ relative to the overlying biogenic carbonate facies (Figs. 8, 9 and 10). Considering primary values forced by occurrence of non-marine water influence on the middle shelf, the diagenetic component of these values probably is best explained in the context of interstitial fluid flow during burial of porous, permeable deposits (van der Kooij et al., 2009). Within this diagenetic context, later veinlet carbonate cements from the Rocha Poço section with very negative δ13C values (Fig. 10) yield a distinct luminescence pattern (Fig. 7C and D) relative to the patterns obtained for epioceanic veinlet carbonate cements (Fig. 7A and B). Microbially induced automicrite from spongiolithic facies in the upper part of the Rocha Poço section shows a range of values subtly lower than those obtained from epioceanic Ammonitico Rosso deposits. This is most probably due to forcing by active organic matter decay related to significant microbialite occurrence, but maintains the same decreasing δ13C trend as Lower Kimmeridgian epioceanic records (Fig. 9A). Similar stratigraphic trends and absolute carbon isotope values (Fig. 9A) reflect the reliability of the marine δ13C signal, buffering a local deviation to mesotrophy in a mid-shelf setting by comparison to dominant oligotrophy in the epioceanic fringe where scarcity in macrobenthos was typical (Olóriz et al., 2002a and references therein). The average 2‰ depletion of belemnite guards in respect to the host matrix micrite (Fig. 8) is interesting because low-Mg calcite belemnite guards commonly are considered a diagenetically stable carbonate

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archive (Veizer, 1983; Saelen, 1989; Longinelli et al., 2003; Brand, 2004). The latter promoted the wide use of belemnite rostra for palaeoenvironmental reconstructions (e.g., Price and Sellwood, 1994; Saelen et al., 1996; Podlaha et al., 1998; Wierzbowski, 2002; Voigt et al., 2003; Rosales et al., 2004; McArthur et al., 2007a; Price, 2010) as they are expectedly low sensitive to post-mortem diagenetic alteration. However, a direct application of this interpretation is not straightforward (Saelen et al., 1996; McArthur et al., 2007b; Richter et al., 2011; Benito and Reolid, 2012). In the investigated case, the obtained offset is broader than expected for the δ13C range in a water column of hundreds of meters depth. This is because a 2‰ offset is reported for present water depths in the range of

kilometers (b2.5 km, Key et al., 2004; Hilting et al., 2008). Given the non-luminescent nature of the rostra (Fig. 7H, J), significant diagenetic alteration seems unlikely. Possible interpretations for these lower values include the incorporation of metabolic 13C-depleted organic carbon (vital effect) or the decay of organic laminae (Price and Page, 2008; Price et al., 2009). Alternative interpretations to in vivo forcing for the low δ13C-values in belemnites (vital effect) include the decay of organic laminae obscurae (Müller-Stoll, 1936; Saelen et al., 1996) that alternate with inorganic ones. Assuming a rather closed diagenetic system, isotopically light organic carbon potentially is contained in early secondary cements occluding pore space in belemnite guards. All of these interpretations are geologically reasonable, but difficult to prove and evidence is

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circumstantial at best in the case investigated here. The application of belemnite data in palaeoceanographic studies must be considered with care given controversial interpretations of their ecology and/or metabolism (Podlaha et al., 1998; McArthur et al., 2004; McArthur et al., 2007b; Price et al., 2009; Richter et al., 2011). 5.1.2. Oxygen isotope composition Matrix micrite oxygen isotope values from the epioceanic sections (Figs. 8 and 9) are 18O-enriched relative to data from coeval depositional settings reported elsewhere (Price and Sellwood, 1994; Veizer et al., 1999; Bartolini et al., 2003; Préat et al., 2006). In absence of remarkable CL reaction to diagenesis, such a difference may result from particular temperature and/or salinity conditions as further discussed. Dense and saline bottom waters have the potential to affect marine porewater diagenesis (Immenhauser et al., 2002; van der Kooij et al., 2009). With reference to the epioceanic sections (Figs. 2 through 4), major seawater salinity changes would affect the sparse benthic communities. However, no evidence for faunal changes related to fluctuations in salinity levels is evident in these sections (Olóriz, 2000). Hence, this scenario can be discarded. Moreover, the epioceanic setting is distant from shelf areas where surface seawater evaporation and bottom topography may result in the formation of locally dense, saline

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water masses or plumes. All this is inconsistent with salinity as the main driver of the relatively consistent behavior of matrix micrite δ18O values obtained from the epioceanic samples studied. For three out of the six epioceanic sections presented, changes in Ammonitico Rosso matrix micrite δ18O values were interpreted more reasonably in terms of changes in water masses, including factors such as upor downwelling, bottom currents, and related spatial palaeotemperature and/or palaeosalinity distribution (Coimbra et al., 2009). Particularly, matrix micrite δ18O ratios from Cardador, Salcedo, Puerto Escaño and Cañada del Hornillo sections (Fig. 9B and Coimbra et al., 2009) are thus considered to represent an uncommon case of micrite isotope values reflecting nearbottom seawater isotopic composition. This fair state of preservation is attributed to very early marine stabilization of thermodynamically unstable carbonate phases (e.g., Coimbra et al., 2009 and references therein). Such a process promoted rapid near-seafloor lithification and incipient nodulation during an early shallow marine burial stage, under the influence of marine pore waters. Therefore the most typical feature of the studied Ammonitico Rosso — i.e., “freezing” of bottom seawater conditions — is interpreted to be closely related to conservative porewater geochemical signals. Sections in the Majorca Island (Figs. 2 through 4) are characterized by lithological differences regarding sections from southeastern Iberia

Fig. A.1. Microfacies details that illustrate Section 4.1.1. Scale bar corresponds to 1 mm. (A) Bioclast-rich and poorly sorted wackestone mainly containing mollusk remains (bivalves), rare cephalopods (aptychus, fragmented septa?) and indeterminate fragments embedded in matrix micrite with isolated or coupled green algae zoospores (Globochaetes) and dinoflagellate cysts. Distinct skeletals range in size from fine to medium sand without evidence of microstratigraphic order (biogenic rather than physical mixing). Note common occurrence of fragmentation with variable orientation of elongated bioclasts, well preservation of shell geometry with no crushing, rare articulate shells and inverted geopetals, and shelly remains of variable thickness. [Puerto Escaño section, sample 18, Kimmeridgian]. (B) Above: Bioclast-rich packstone showing extreme packing of elongated, fine and recrystallized skeletals (stacking of pelagic bivalves — “filaments”). Shelly lag composed of dominant parallel arrangement of coarse sand sized bioclasts, most probably due to winnowing of matrix micrite, rapid wanishing of energy and redeposition of sorted skeletals (“filaments”). No significant lateral transport is favored rather than reworking and redeposition by relatively distant energy sources (distal storms?; shelly turbidite lag due to slope sliding of muddy sediment?) because of the occurrence of “filaments” as part of a more variable bioclastic content below the sharp contact indicating micro-erosion surface. Below: Bioclast-rich wackestone containing pelagic bivalves (“filaments”), pellets, sponge spicules and radiolaria. Parautochtonous fosssil assemblage. [Cuber section, sample 2, Oxfordian]. (C) Poorly sorted bioclast-rich wackestone containing recrystallized skeletals (bivalves, fragments and spines, planktic foraminifera, gastropods) of mainly fine sand size, with scarce larger particles (brachiopods, isolated and coupled green algae zoospores — Globochaetes- and common echinoderm plates). Underlying is a very fine silty, bioturbated mudstone to bioclast-poor wackestone with scattered planktic foraminifera (Globuligerina) and isolate green algae zoospores (Globochaetes), disarticulated bivalves, radiolaria, and plates of echinoderms (out of view). Note sharp contact indicating a rapid accumulation event revealed by deposition of a shell lag resulting from increased energy, winnowing, sorting and deposition on microerosion surface (laterally depressed by shell lag excavation out of view where circular and irregularly elongated micrite-to microspar patches reveal bioturbation). Flattened matrix patches (burrows) agree with sedimentary load crushing of tiny shells as shown below and left in the underlying mudstone horizon. [Cañada del Hornillo sample 7, Kimmeridgian]. (D) Bioclast-rich wackestone corresponding to micro-horizon defined by relative enrichment of elongate skeletals (coarse sand sized pelagic bivalves — “filaments”) determining subtle subparallel lamination due to episode of lowered net accumulation rate. Other biogenic remains are sponge spicules, pellets and less common radiolaria. Note overlying dilution of smaller “filaments” (b500 μm) due to subtle change in environmental energy [Cuber section, sample 35, Kimmeridgian]. (E) Unsorted wackestone showing millimeter sized ammonite shells with differential infilling of camerae and well preservation of unflattened but broken shells embedded in matrix micrite with mainly indeterminate, silty sized carbonate particles, rare disarticulated bivalves, echinoderm plates, spines and indeterminate fragments, tiny gastropods, common isolated and coupled green algae zoospores (Globochaetes), and dinoflagellate cysts. Low energy deposit containing parautochtonous fossil assemblage and experiencing very early lithification. [Cardador section, sample 10, Kimmeridgian]. (F) Mudstone with scarce sponge spicules, rare small radiolaria, scarce isolated green algae zoospores (Globochaetes) and dinoflagellate cysts, and thin, elongated skeletals (“filaments”) enclosing uncrushed, ammonite shell with recrystallized interior. Note perfect preservation of shell geometry in transverse section and a single, very restricted break allowing fine sediment to enter the shell interior. Event of muddy deposition (advection from winnowing) rapidly buried the ammonite shell most probably containing fleshy tissue impeding the sedimentary infilling of the carcasse; lithification occurred very early during diagenesis preventing crushing by sedimentary load and preceding later cementation of the empty shell. [Cala Fornell section, sample VII, Tithonian]. (G) Detail of spongiolitic limestone showing broken sponge frame encrusted by microbial growth with peloids, small indeterminate intraclasts and scarce gastropods, and embedded in mudstone to bioclast-poor wackstone with sponge spicules and indeterminate carbonate particles of silt size. [Rocha Poço section, sample 15A, Kimmeridgian]. (H) Wavy, irregularly laminated fabric in stromatolitic microbiallite included in spongiolitic limestone. Note local merging of brown cianobacterial films as well as their interruption by fine silt sediment. Out of view there are thrombolitic growths. [Rocha Poço section, sample 16, Kimmeridgian]. (I) Articulated bivalve trapped in microspar infilling of intersticial space in spongiolitic limestone. Note enclosing bioclast-poor wackestone matrix containing scarce sponge spicules, lumps, small oncoids with large recrystallized nuclei (space for air or fleshy material?) and nubeculariids, as well as the continuity of the mudstone-to-bioclastic poorer wackestone inside the bivalve. Except for the bivalve (2.5 mm high), biogenic particles range from silt to less common medium sand size. Out of view there are scarce isolated green algae zoospores (Globochaetes), indeterminate calcitic and agglutinated benthic foraminifera, echinoderm plates, microbial/algal growths enclosing a distinct, large recrystallized voids corresponding to probable space for air or fleshy material, small intraclasts and peloids, and recrystallized voids separating patches of the wackestone matrix. [Rocha Poço section, sample 19, Kimmeridgian]. (J) Lumpy/intraclastic/microbialitic limestone showing micrite-to-microspar matrix containing grains in the range of fine to very coarse sand. Common intraclasts, aggregate grains, microbial/algal lumps, oncoids with sponge spiculae and indeterminate carbonate particles in the nuclei, as well as cases of microbial/algal growths enclosing a distinct, large recrystallized void corresponding to probable space for air or fleshy material. Note peripheral nubeculariids embedded in microbial/algal structures and rare remains of probable calcitic foraminifera. Out of view there is an incomplete ammonite phragmocone, 0.5 mm in size, showing recrystallized nucleus, whorls third and forth in filled by intraclast/bioclast free micrite-to-micropar, and outer whorl preserved in filled according to the enclosing lumpy matrix, a brachial valve of brachiopod, 2.2 mm in size, with dominant micrite in filling and scarce lumps, scarce indeterminate foraminifera (agglutinated forms), and common nubeculariids and Tubiphytes. [Rocha Poço section, sample 24, Kimmeridgian]. (K) Lumpy/intraclastic/microbialitic limestone with micrite-to-microspar matrix containing grains in the range of fine to very coarse sand. Common irregular microbial/algal growths (lumps) of variable size and oncoids enclosing carbonate particles and fossil remains (biseriate lituolacean forams?, peripheral nubeculariids, sponge spiculae, mollusks) and micritized and/or algal intraclasts. Note microbial/algal encrustment of a large void (algal air bladder, gas bubble or decayed fleshy material) infilled by micritic-to-microsparitic matrix with small fragments of micritized intraclasts or microbial/algal growths inside, as well as the combined record of uncoated and coated grains (cortoids). Out of view are rare coupled green algae zoospores (Globochaetes) and scarce remains of calcitic and agglutinated benthic foraminifera, small disarticulated bivalves and ostracods enclosed or not en microbial/algal growths, and echinoderm fragments. [Rocha Poço section, sample 25, Kimmeridgian]. (L) Lumpy/intraclastic/ microbialitic limestone with micrite-to-microspar matrix with low sorting of grains ranging from coarse silt to very coarse sand. Note thrombolitic growth encasing microbial/algal lumps, oncoids and sponge spiculae, and adjacent micritic to microsparitic matrix containing floated lumps, micritized oncoids, uncoated and locally coated mollusk fragments, common fragments of microbial/algal structures of variable size (micritized and/or algal intraclasts, tuberoids?) including encrustment by nubeculariids, and a large recrystallized and locally bored fragment showing local coating, as well as coated grains (cortoids). Out of view are large fragments of sponge frameworks, common sponge spiculae, disarticulated bivalves, ostracods, echinoid fragments, plates and spines, fragments of camerate ammonite shells, rare miliolids, and coupled green algae zoospores (Globochaetes) [Rocha Poço sample 29, Kimmeridgian].

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(Figs. 2 and 4). Specifically, Majorcan sections show Ammonitico Rosso facies, but not as typical/nodular as in other epioceanic sections, as well as siliceous horizons and common reworked horizons in the upper part of these sections. Considering early, near-seafloor lithification and nodule formation (Coimbra et al., 2009), these features might point to differential depositional conditions and diagenetic paths. Nevertheless, marked diagenetic alteration is excluded, because (1) the chemostratigraphic record of Majorcan sections is very similar to that from southeastern Iberia, and (2) matrix micrite δ18O values differ from those of late diagenetic cements in these sections. These observations point to chemostratigraphic patterns that are largely independent of local macroscopic deviations in lithofacies within the depositional context corresponding to the array of Ammonitico Rosso facies investigated. Epioceanic carbonate cement oxygen-isotope values from different bright luminescent cements phases (Fig. 7B) in veinlets are up to several per mil lower than those from the encasing matrix micrite (Figs. 8 and

10). These lower values provide evidence for the influence of early to later burial fluids (Fig. 10), characterized by increasingly elevated temperatures. In general, matrix micrite δ18O was stabilized diagenetically at a shallow burial marine porewater stage, as confirmed by the characteristic luminescence pattern (Bruckschen and Richter, 1994) recognized for carbonate cements from veinlets in the epioceanic sections (Fig. 7B and overlapping values in Fig. 10). Later diagenetic stages are not excluded for carbonate cements with lower oxygen isotope values (Fig. 10). This observation further reinforces the “conservative” character of matrix micrite in the studied Ammonitico Rosso facies. The epicontinental matrix micrite oxygen isotope record from the Rocha Poço section is the most persistently depleted section, markedly negative in the stratigraphically lower siliciclastics-rich deposits (about − 4.5‰). This value is interpreted as the most probable effect of mixing of marine and freshwater primary signals (as deduced from common remains of plants and quartz) combined with later phreatic

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Figs. A.2 and A.3. Close-up views (white rectangles) providing visual support to some features described in Section 4.1.1. Scale bar corresponds to 1 mm. (A) Bioclast-rich wackestone showing millimeter sized gastropods embedded in well sorted matrix micrite to microspar with very fine to fine sand size planktic skeletals (Globuligerina), and no evidence of microstratigraphic order. Note partial recrystallIzation of shell interior in incomplete gastropods (normal geopetal orientation) and outer whorls in filled by the surrounding matrix with skeletals. Transportation of partially empty gastropods as allochtonous grains, and deposition in distal tempestite or turbidite front that affected pelagic sediments — calcareous ooze with foraminifera. Out of the magnified view, aptychus, ammonitella?, echinoderm plates, bivalves, ostracods, benthic foraminifera (lageniids and indeterminate calcitic forms) and green algae zoospores (Globochaetes) can be found; matrix heterogeneity, micrite patches and variable orientation in geopetals could indicate burrowing. Parautochtonous– allochtonous fossil assemblage. [Cardador section sample 3B, Oxfordian]. (B) Poorly sorted and matrix micrite-rich bioclastic wackestone showing preservation of elongated and variably oriented fragments of mollusks and echinoderms (out of the maginified view are aptychi, common ammonitella and less common juvenile specimens, gastropods, bivalves, ostracods and green algae zoospores — Globochaetes). Note vertically embedded, uncrushed bivalve shell and the occurrence of relatively thick fragments of medium to coarse sand size carbonate skeletals. Distinct patches of micrite-to-microspar devoid of skeletal indicate burrowing, while combined microsparitization and ferruginization selectively affect shells' interiors. Microbioerosion is recorded and severely affected to skeletals showing cracked aspect due to later ferruginization and irregular crystal growth (cementation). Background sedimentary horizon, biogenically mixed, lithified very early during diagenesis and containing parautochtonous, averaged fossil assemblage. [Cañada del Hornillo section, sample 23, Kimmeridgian]. (C) Bioclast-rich wackestone to local packstone of mainly elongated, recrystallized skeletals (mainly echinoderm fragments and Saccocoma; rare indeterminate calcitic, benthic foraminifera out of the magnified view). Note variable thickness and orientation of medium to fine sand size skeletal fragments embedded in matrix micrite, as well as low rounding indicating low transportation. Depositional event forced by environmental energy higher than background, irrelevant lateral transportation, and mixing of long-lasting exposed skeletals punctured by micro-bioerosion with assumed more recent ones without micro-bioerosion marks. Parauthochtonous, averaged fossil assemblage. Highly variable orientation and packing, and no evidence of microstratigraphic order are relevant traits out of the magnified view [Cañada del Hornillo section, sample 18, Kimmeridgian]. (D) Poorly sorted and bioclast-rich wackestone with common preservation of uncrushed, inflated and convex shells (ammonitella, molluscs) embedded in matrix micrite with common green algae zoospores (Globochaetes). Out of the magnified view are echinoderm plates, saccocoma, tiny bivalves and gastropods, aptychi and rare, broken benthic calcitic forams, as well as variable record of micro-bioerosion. Note dominant coarse sand size of distinct bioclasts. Background sedimentation with parautochtonous fossil assemblage, lithified very early during diagenesis. [Cañada del Hornillo section, bed 21, Kimmeridgian]. (E) Bioclast-rich wackestone containing common planktic foraminifera (Globuligerina), less common small bivalves and ostracods, echinoderms fragments and indeterminate carbonate particles. Except for rare, larger remains of ammonite phragmocones (out of the magnified view), note particle sorting at the fine sand range, micritized and ferruginized infilling of foraminifera, local recrystallization of camerae, variable orientation of bivalve shells, and rare articulate ostracods. Out of the magnified view micro-bioerosion is widespread, and burrowing accords with irregular distribution of matrix micrite and skeletals' concentrations (echinoderms plates, spines and fragments, common gastropods and bivalves, uncommon fragmented ammonite phragmocones, and rare aptychi, ammonitella, fragmented benthic calcitic forams and green algae zoospores — Globochaetes). Winnowed and bioturbated, pelagic background sediment with parautochtonous fossil assemblage. Common fragmentation but no shell's crushing indicates very early lithification [Puerto Escaño section, sample 15, Kimmeridgian]. (F) Bioclast-rich wackestone showing disordered, uncrushed skeletals (mainly mollusks, articulated and disarticulated small bivalves and ostracods, saccocoma, echinoderm fragments, planktic forams), variable infilling of shells' interiors and local nesting. Out of the magnified view are Globuligerina, gastropods, echinoderm plates, ammonitella an some fragments of ammonite phragmocones. Storm shelly bed, sorted in the range of medium sand (except rare, slightly larger ammonite fragments), bioturbated (patchy distribution of skeletal abundance and fossil-poor micrite matrix) and lithified early during diagenesis [Cañada del Hornillo section, sample 12, Kimmeridgian]. (G) Bioclastic wackestone containing parautochtonous fossil assemblage, with common ammonitella, aptychi and echinoderm plates and spines, saccocoma, gastropods, bivalves, mollusks fragments, sponge spicules, benthic calcitic forams, and green algae zoospores — Globochaetes (out of the magnified view). Matrix micrite with small indeterminate fragments of silty size carbonate particles. Note preservation of geometry of the tiny carbonate structure in ammonitella showing micritized protoconch, nepioconch in filled by surrounding matix, primary varyx, nepionic constriction and uncrushed shell, as well as well preserved curvature in bivalve shells. Background sediment showing irregular distribution of skeletal concentration indicating burrowing (out of the magnified view is a 4 mm wide burrow piping the wackestone in filled by fine to coarse sand size saccocoma and echinoderm fragments) [Cañada del Hornillo section, sample 25, Tithonian]. (H) Bioclast-rich packstone showing dense packing of disordered and mainly elongated skeletals (stacked and nested pelagic bivalves — “filaments”) most probably due to storms. Note locally trapped peloids, variable recrystallization of “filaments”, and matrix micrite patches. The latter are more evident out of the magnified view where some of them could relate to burrowing (“filaments” adapt to their margins) and there are good examples of cemented shelter porosity voids, and rare trapped calcareous forams (lageniids?), green algae zoospores — Globochaetes — in stems and thick mollusks fragments. Storm, locally bioturbated shelly lag composed of dominant coarse sand size homogeneous skeletals. Parautochtonous fossil assemblage [Cardador section, sample 1, Oxfordian]. (I) Mudstone to bioclast-poor wackestone showing some sponge spicules among indeterminate skeletal remains and cloudy aspect of micrite-to microspar matrix most probably due to burrowing. Background sedimentary horizon containing parautochtonous fossil assemblage with skeletals belonging to common echinoderm plates, ossicles, fragments, saccocoma and green algae zoospores (Globochaetes), less common tiny gastropods, disarticulated bivalves, radilolaria and ammonitella, and rare belemnites and juvenile ammonites and aptychus out of the magnified view. [Cañada del Hornillo section, sample 33, Tithonian]. (J) Medium to coarse silt with common quartz grains, small intraclasts and peloids, very rare benthic foraminifera (Lenticulina) and isolated disarticulated ostracods. Out of the magnified view are scattered plant remains and indeterminate, fragmented benthic, calcareous foraminifera. [Rocha Poço section, sample 10, Kimmeridgian].

(freshwater/brackish) water entering porous sediments prone to diagenetic imprint. In contrast, above a marked lithological change, slightly higher δ18O values (−3.5‰ on average) in the spongiolithic facies might suggest decreasing continental influence in these more carbonate-rich sediments. The averaged δ18O value of −3.5‰ agrees with values of −4.0‰ reported for Late Jurassic analog materials (Keupp et al., 1993; Ruf et al., 2005), in which diagenesis was somewhat relevant. Clearly, epicontinental

spongiolithic limestones underwent a diagenetic pathway that is different from that of typical Ammonitico Rosso facies from the epioceanic fringe. Given limitations for interpreting pristine δ18O values from silty limestone and spongiolithic facies, a suitable record of water column, oxygen isotope composition for the Rocha Poço section was measured from the single and exceptionally well preserved belemnite rostrum retrieved from spongiolithic limestones in the upper part of the section.

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Table 1 Mean carbon and oxygen isotope composition.

Matrix micrite Belemnite rostra Cemented veinlets

δ13C (‰) (V-PDB)

δ18O (‰) (V-PDB)

n

1.9 ± 0.8 0.4 ± 0.7 2.0 ± 0.3

−1.3 ± 1.5 −0.6 ± 0.3 −3.7 ± 1.9

1150 19 102

The δ18O value of −0.9‰ (Table 3) corresponds to a seawater temperature range of 16–23 °C (for details on belemnite ecology see Rexfort and Mutterlose, 2006, 2009, among others). Even considering that microbialite isotopic values represent signals from benthic microbial/ algal consortia, while belemnites inhabited the water column, diagenesis did lower the original microbialite oxygen isotope values down to the present value of −3.5‰. Although no data were available from intermediate phases, a rather complex diagenetic course is envisaged to account for lower (negative) oxygen isotope values based on: (1) an earliest but difficult to quantify imprint most probably forced by biodiagenesis (Neuweiler et al., 2003; Olivier and Boyet, 2006; Desrochers et al., 2007; Neuweiler et al., 2007) in sediments that experienced high microbial activity; (2) limited reworking resulting in microbially induced automicrite to be recycled into detrital allomicrite (Neuweiler et al., 2003) in a context of sedimentary contribution of fine clastics and inner-shelf carbonate grains; and (3) potential phreatically forced diagenesis affecting bulk, mixed carbonate-fine clastic sediments, depleting the oxygen isotope signal towards values of −3.5‰. Obtained mean values of δ13C and δ18O ratios per sampled horizon in matrix micrite from epioceanic sections (Fig. 10) reflect isotope signals from open marine bottom (pore) waters, while data from veinlets point to diverse burial fluids. Marine patterns and continental influx may interact in the shallow epicontinental Rocha Poço section. A distinct feature of the epicontinental geochemical signature is the coherence between carbonate cement and encasing matrix micrite values (Fig. 10). A non-marine influence on the primary matrix micrite signal is likely, further enhanced by later phreatic influence. Overall, epioceanic belemnite δ18O values of screened specimens (Table 3) oscillate around a mean value of − 0.6‰. This value was used for palaeotemperature estimates using the equation of Anderson and Arthur (1983), rendering a range of 15–21 °C. Such results are in agreement with coeval, well-preserved belemnite data from Majorca Island (Price and Sellwood, 1994) and also consistent with direct temperature measurements from modern analogs such as the Hancock seamount, situated at latitude of approximately 30°N on the Hawaiian Ridge (Boehlert, 1988, see Coimbra et al., 2009 for further analogs). This provides evidence on the preferred habitat of these nektonic organisms, preferentially inhabiting selected/particular marine settings in the oceanic fringe providing water masses with temperatures in the range of 15–21 °C. The ammonite shells show similar luminescence patterns (Fig. 11A, B), whereas ammonite δ18O ratios are variable (Fig. 11E). The ammonite shell from the Cañada del Hornillo section presents a δ18O value close to that of the encasing matrix (− 1.1 and − 0.7‰, respectively, Fig. 11E). In contrast, the specimen from the Puerto Escaño section

Table 2 Mean elemental composition for analyzed carbonate materials.

Matrix micrite Belemnite rostraa Cemented veinlets a

Mg (ppm)

Fe (ppm)

Mn (ppm)

Sr (ppm)

n

2707 2575 3165

1657 150 2379

331 45 757

213 1040 278

1150 19 102

Only best preserved specimens (see Table 3).

shows a δ18O value similar to that of late cement δ18O ratios from the same locality (− 4.9 and − 5.7‰, respectively, Fig. 11E). CL images and pixel counting provided data to quantify the relative proportion of different paragenetic phases (Fig. 11C, D). Phase 1 is abundant in some cases (Fig. 11B and D), occluding a total of 40% (locally almost 100%, see Fig. 7F) of the former aragonitic shell. No evidence of shell collapse was observed, which would be prevented by very early, progressive aragonite stabilization into aggrading neomorphic calcite. If so, the alternative of dissolution producing moldic porosity does not apply for this initial phase 1. Supporting this assumption, O-isotope values from ammonite shell material where phase 1 dominates (Fig. 11D) match those of matrix micrite δ18O as would be expected for neomorphic calcite (Fig. 11E). For this reason, ammonite isotope ratios (both C and O) plot close to the assumed marine values since neomorphic phase 1 calcite and matrix micrite both formed early during diagenesis. Hence, this phase is interpreted as aggrading neomorphic calcite replacing original aragonite at a very early diagenetic stage (e.g., Hendry et al., 1995). In the other ammonite shell analyzed (Fig. 11A and C), phase 1 comprises only about 20% of the total bulk shell material (Fig. 11C). It is hence dominated by the presence of luminescent carbonate phases 2 and 3, both bright luminescent. Ammonite δ18O plots within the range of late diagenetic veinlet carbonate cements (Fig. 11E). Based on this pattern, phases 2 and 3 in ammonite shells are interpreted as later cements, in contrast with phase 1. Explanations for the presence of luminescent phases 2 and 3 relate to cement in filling of previous voids adjacent to neomorphic calcite, or to partial dissolution of preserved aragonite patches and later in filling by cement precipitation (e.g., Maliva and Dickson, 1992; Caron and Nelson, 2009). 5.1.3. Silica diagenesis Siliceous horizons of the Cuber section mainly occur in Upper Oxfordian, Kimmeridgian and Middle to Upper Tithonian deposits (see Figs. 3 and 9). Marine siliceous skeletons, e.g., sponge spicules, diatoms, and radiolarians, could provide biogenic opal A via dissolution of metastable opaline SiO2 (Fanning and Schink, 1969; James et al., 2000). Given the open distal marine depositional setting at Cuber (Fig. 4), sponges can be ruled out as the main source of opal, in agreement with the palaeogeographic range of the so-called “sponge megafacies” identified for the Upper Jurassic in epicontinental southern Europe (Pisera, 1991; Pisera et al., 1992). Hence, planktic marine organisms might be responsible for opaline silica enrichment in sediments. Since records of diatoms are not reported prior to the Cretaceous (Gersonde and Harwood, 1990; Falkowski et al., 2004), radiolaria rather than diatoms would respond to potential fertilization, consistent with the observation of abundant radiolaria in thin sections of the Cuber section (Fig. A.1B). The high manganese content for these silica-bearing horizons (Figs. 12 and 13) may suggest a hydrothermal influence (Bender et al., 1970; Kickmaier and Peters, 1990; Corbin et al., 2000), which would operate as water column fertilizer. Alternative Mn sources (e.g., diagenetic enrichment or continental influx) can be ruled out since no significant diagenetic imprint could be recognized for the Cuber section and distance from shore makes continental Mn supply unlikely (see Fig. 4). In fact, further evidence of fair preservation was demonstrated by smallscale geochemical isotope transects showing no influence of silica diagenesis in matrix micrite isotope values surrounding siliceous nodules (Fig. 12). Again conservative matrix micrite geochemical signals emerge even in these less typical Ammonitico Rosso horizons with indistinct nodulation. The overall depleted matrix micrite δ18O values of silicabearing horizons pointing towards warmer palaeotemperature, coupled with the lack of evidence for marked diagenesis within these intervals, is consistent with an interpretation of biogenic silica accumulation related to locally warmer water masses and increased nutrient gradients (Jones and Jenkyns, 2001). This interpretation is consistent with the inverse correlation between matrix micrite δ18O and Mn abundance (Fig. 13) and fits a scenario of expected hydrothermal influence.

R. Coimbra et al. / Earth-Science Reviews 139 (2014) 1–32

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Table 3 Carbon and oxygen isotope composition of best preserved belemnite rostra.

C S CH CU

PE

RP CF

a

δ13C (‰) (V-PDB)

δ18O (‰) (V-PDB)

Mg (ppm)

Fe (ppm)

Mn (ppm)

Sr (ppm)

Δδ13Ca (‰) (V-PDB)

Δδ18Oa (‰) (V-PDB)

0.9 xx 0.8 1.3 −0.2 −0.6 −0.2 0.8 0.1 0.8 −0.5 −1.4 0.7 0.4 0.3 0.0 1.2 0.5 0.8 1.0

−0.5 xx −0.3 −0.8 −0.9 −0.5 −0.8 −0.5 −0.5 −0.3 −0.6 −0.1 −0.9 −0.9 −0.6 −0.1 −0.5 −1.0 −0.7 −1.5

3381 xx 2000 3038 2398 2340 2488 2896 2243 2600 2130 2229 2684 3249 2147 1909 2611 2622 3781 2170

126 xx 163 79 107 62 185 465 99 512 89 65 45 53 273 51 187 248 440 239

100 xx 47 30 22 6 27 61 34 70 54 20 1 1 100 65 25 42 75 58

971 xx 926 987 1129 1171 935 821 1213 1180 1085 1266 1286 1187 847 1019 774 906 914 842

−1.0 xx 0.0 0.0 +2.7 +1.8 +1.5 0.0 0.0 +0.5 +0.3 +1.0 +3.0 +3.0 +0.5 +1.0 +1.0 +1.0 +1.0 +1.0

−1.5 xx −1.6 −1.2 −2.7 −2.1 −1.8 −1.8 −2.0 −1.2 −2.5 −3.5 −1.0 −1.0 −2.6 −2.4 −1.2 −2.0 −2.0 −1.5

Offset between C and O isotope composition of belemnite rostra and neighboring encasing micrite matrix.

5.2. The epioceanic record Having excluded significant diagenetic alteration in epioceanic Ammonitico Rosso samples, isotope data can be placed in their palaeogeographical context across the proximal-to-distal transect investigated. The carbon-isotope curves show a similar stratigraphic pattern in all epioceanic sections (Fig. 9A), matching published coeval data (Bartolini et al., 1996; Weissert and Mohr, 1996; Bartolini et al., 1999; Cecca et al., 2001; Rey and Delgado, 2002; Savary et al., 2003; Rais et al., 2007; Coimbra et al., 2009). Epioceanic sections appear to record an at least Tethys-wide chemostratigraphic pattern of seawater δ13CDIC. The overall Tethyan epioceanic δ13C pattern is not overprinted by potential, local palaeoenvironmental deviations. This interpretation is relevant given the irregular bottom palaeo-physiography across the study area characterized as a swell-and-trough relief resulting from an inherited bottom relief and syn-sedimentary tectonics (Fig. 4). This scenario favors spatially complex, local productivity changes, similar to those in modern oceans due to up- and downwelling patterns (Boehlert, 1988; Beckmann and Mohn, 2002; Auster et al., 2005), as well as associated local hydrodynamics (Fig. 4). Nonetheless, the consistency of epioceanic data suggests that any local isotope signals were buffered by the signature of the open, oceanic water masses. Similar to carbon, the matrix micrite δ18O record of all epioceanic sections is relatively consistent. Overall, excursions in matrix micrite δ18 O tend to be more pronounced in the most distal section (Cardador section; Fig. 9B) but are still evident in more proximal areas. Interestingly, matrix micrite δ 18 O seems to follow relative changes in sea level, with values being more positive (1‰ higher) during periods of high sea level, recording thus relatively deeper, hence cooler conditions, especially when major flooding events took place in both epicontinental areas and pelagic platforms. Examples of the previous include the Mid-Oxfordian transgression and the Jurassic eustatic maximum towards the end of the Kimmeridgian (Haq et al., 1988; Hardenbol et al., 1998). Relative differences in absolute oxygen isotope values in different epioceanic sections (Figs. 8 and 9) could relate to local differences in palaeogeographic setting and related hydrodynamics. Palaeotemperature estimates can clarify observed differences in oxygen-isotope values. Acknowledging the fundamental problems of comparative seawater palaeotemperature estimates, the equation of Anderson and Arthur (1983) was also applied to matrix micrite δ18O

ratios from epioceanic sections. In agreement with the proposed conservative character of matrix micrite, the record of oxygen isotope values, CL and major and trace elements (Figs. 7, 8, 9, 11 and Tables 2 and 3) suggest that the dominant Ammonitico Rosso lithofacies within the epioceanic realm underwent a very early lithification process. This is consistent with previous interpretations (Coimbra et al., 2009; Coimbra and Olóriz, 2012c) and agrees with the new data obtained. Palaeotemperature estimations based on matrix micrite mean δ18O values are interpreted as corresponding to porewater conditions in near-equilibrium with bottom water masses. Proximal to distal averaged δ18Omicrite ratios from epioceanic sections suggest a bottom seawater palaeotemperature range of 15–22 °C for the Cañada del Hornillo and Puerto Escaño sections, 20–27 °C for Cuber and Cala Fornells sections, and 11–17 °C for the more distal Cardador and Salcedo sections (Figs. 2 through 4). In this way, less distal epioceanic sections seem to reflect slightly warmer pore water palaeotemperatures of 15–22 °C. Hence, the trend is consistent with their palaeogeographic and relative location along the studied transect (Figs. 3A and 4) and the concept of increasingly cooler temperatures seawards in the epioceanic fringe, irrespective of potential slight difference in depth within a typically structured Tethyan margin. The sections from the Majorca Island (Figs. 2 to 4) are interpreted to record the overall warmest palaeotemperatures of 20–27 °C. Considering that δ18O is a diagenesis-sensitive proxy, complementary screening methods as elemental crossplots (Sr/Ca and Mn/Ca versus δ18O, see Coimbra and Olóriz, 2012c) attested for the good preservation state of matrix micrite in both sections. This validates interpretations of matrix micrite oxygen isotope values in these sections within the context of the transect under scope. The overall higher palaeotemperatures might either place these sections at an overall less distal, shallower setting within the epioceanic fringe, or reveal local differences in palaeoceanographic parameters (e.g., bottom physiography and related current patterns, see Fig. 4) relative to sections from southeastern Iberia. A differential palaeotemperature for the Majorca sections could agree with palaeogeographic, and therefore palaeoenvironmental, signatures of the eastern versus the southeastern epioceanic segment of the Iberian paleaomargin (see Fig. 4) in a context of episodic hydrothermal influence. A statistically significant linear correlation (R = 0.6; p b 0.0001) between oxygen-isotope composition of matrix micrite and the estimated

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abundance of bioclasts was obtained for all epioceanic sections (Fig. 14A). Plotted in a stratigraphic context (Fig. 14B), maxima in δ18O and skeletal abundance are roughly coincident (Fig. 14B). In these data, a potential influence of isotope fractionation is unlikely since no particular recurrence in microfossil assemblages occurs throughout the sections. The general pattern pointing to the coupling of δ18O signals and skeletal abundance cannot be unequivocally interpreted in palaeoecological terms, especially in epioceanic waters, and further research is needed. A particular case of both δ18O and skeletal maxima (Fig. 14B) is that of calcareous, less nodular, grayish Ammonitico Rosso marker beds, which were used for lateral correlation of time-equivalent deposits (Kimmeridgian/Tithonian boundary, Fig. 5A). Coimbra et al. (2009) suggested that the δ18O record of the epioceanic sections follows Late Jurassic sea level fluctuations. Specifically, cooler seawater temperatures during higher sea levels were interpreted to be recorded by more positive δ18O values (Fig. 9B), placing the studied areas in relatively deeper, hence cooler settings during higher sea-level times. According to previous studies dealing with Ammonitico Rosso pigmentation (Berner, 1969; Jenkyns, 1971; Mamet and Préat, 2003, 2006a; Préat et al., 2006; van der Kooij et al., 2007; Préat et al., 2008), grayish facies relate to high energy, oxygenated conditions that also could apply to microbially mediated mounds (Boulvain et al., 2001 and references above). Nevertheless, grayish Ammonitico Rosso horizons can be commonly intercalated among the characteristic red Ammonitico Rosso facies, or even recorded as evidence of lateral change of lithofacies pigmentation. This color variation reveals a heterogeneous distribution of redox conditions within the substrate (e.g., Mamet and Préat, 2003, 2006a). Overall, markedly suboxic porewater is unlikely for the Jurassic epioceanic setting studied here. Active bioturbation favoring nodulation early during diagenesis within the taphonomically active zone (see Coimbra et al., 2009 for further details) reinforces this interpretation. Thus, oxygenation levels that were high enough for Thalassinoides makers and the corresponding irrigation of the colonized sedimentary column seem likely. In a context of sea-level change, initial transgressive phases commonly produce ravinement surfaces — i.e., bioclastic-rich horizons related to erosion and most usually interpreted as resulting from transgressive forcing. Later during transgression, due to persistent rising of sea level, bottom-current energy commonly decreases, resulting in reduced winnowing, sorting and changes in net accumulation rates on epioceanic swells (raised sea bottoms). The case of less nodular, grayish marker beds (Figs. 3B and 5A) that correlate with the eustatic maximum near the Kimmeridgian/Tithonian boundary exemplifies particular sedimentary conditions. In these marker beds the occurrence of subtle laminated fabrics (Saccocoma-rich horizons) supports the interpretation of a higher than averaged environmental energy and accumulation rate for typical Ammonitico Rosso facies. In the context of the eustatic maximum for Jurassic times, these Saccocoma-rich horizons were associated with winnowing and reworking (see Appendix 2 Figs. A.1 to A.3 for illustration), the latter commonly most probably related to storms and/or bottom instability. The recorded 1‰ increase in δ18O values during the highest sea level (Fig. 9) is interpreted as a result of 4 °C decrease in bottom temperature. Assuming normal oceanic seawater salinity (35 psu) for the epioceanic setting, this decrease would result in an increase of about 10% in the concentration of dissolved oxygen in seawater (Kester, 1975). 5.3. The epicontinental record In terms of its geochemistry, the epicontinental Rocha Poço section (Figs. 2 to 4) stands in sharp contrast with epioceanic sections. Differences are present mainly in the stratigraphically lower portions of the section. There, matrix micrite δ 13 C and δ 18 O values from the silty, lower portion of the Rocha Poço section (Figs. 2C, 3B for facies distribution) are depleted with respect to the overlying

facies and the other studied sections (Figs. 8, 9 and 10). Possible mechanisms include (1) differential diagenetic alteration, (2) changes in local water mass properties, (3) an admixture of carbonate material from different source areas, or (4) a combination of these factors. This epicontinental section is the only site among all the sampled localities where a major, most probably combined, influx of sediments derived from emerged areas and inner-to-mid platform settings, was recorded. A final consideration concerning the relative depletion of matrix micrite δ13C in the lower part of the Rocha Poço section points to potentially long residence time of seawater on platforms. This effect results in “aging” of the affected water mass, during which the input of 12C from remineralized organic carbon may lower the δ13C values up to 4.0‰ (Immenhauser et al., 2002, 2008), as well as from higher nutrients and organic matter levels derived from lower exportation of organic matter to open seas. Moreover, depleted carbon isotope values may also result from primary differences, with silty to fine sandy limestone and marls favoring meteoric (phreatic) diagenetic imprint due to grain size and related porosity. The development of sponge bioherms above the silty facies with ammonites indicates changing ecological conditions. These relate to carbonate margin progradation, counteracting second order sea-level rise and favoring redistribution of detrital input and promoting its local decrease, together with bottom stabilization favoring sponge growth. Accordingly, the third order peak transgression during the Early Kimmeridgian (Divisum Chron) is identified by an ammonitebearing marker limestone bed interrupting spongiolithic facies (Fig. 9). Furthermore the characteristic carbon isotope background signature for Jurassic marine water of 2‰ recorded from the epioceanic sections is also identified in bulk samples from sponge bioherms and stratigraphically overlying limestones (Fig. 9A). In the lowermost part of the spongiolithic facies, fine detrital carbonate material with more depleted δ13C values (about 5‰ lower) fills interstices and cavities in the buildup frame. Interestingly, these depleted values are not found further up-section where spongiolithic facies are fully established (Fig. 9A), thus revealing prevailing mesotrophic conditions favoring sponge–microbial–algal consortia. Within the spongiolitic facies, a single bed shows a pronounced negative carbon isotope shift (Fig. 9A). With respect to encasing horizons, microfacies analysis of this bed revealed lower bioclast content and lesser development of microbial–algal structures. In the absence of evidence for subaerial exposure, this particular record is noteworthy and will be interpreted as resulting from a palaeoenvironmental event. The single-bed occurrence recording a sharp decrease in δ 13 C, coupled with lower Mg, Sr and Fe content and no changes in Mn concentration most likely points to rapid change in seawater conditions. The alternative of bed-specific differential diagenesis is not favored. Declining productivity and related, combined effects (alkalinity, pH, CO2) forced an episode of carbonate crisis, in agreement with corroded surface and distinctive reddish coloration in this particular bed. This single-bed scenario is compatible with the influence of freshwater (e.g., intense rainfall event), excluding direct continental input (river discharge) and meteoric diagenesis, the latter usually affecting a more extended sedimentary package (e.g., Allan and Matthews, 1982). The higher complexity of the obtained epicontinental record contrasts with the epioceanic data set, encouraging comparisons of both spatial as well as temporal changes. The outcome was a better understanding of forcing processes underlying these complex archives. 6. The wider context In marine geochemistry, spatial versus single site approaches for interpreting geochemical data represents a serious challenge for a better understanding of complexity in environmental dynamics based on

R. Coimbra et al. / Earth-Science Reviews 139 (2014) 1–32

the analysis of its identifiable traces in both, the sea-water column and the sediments, skeletals included. Information on present-day oceans resulted from the pioneer contributions in the seventies (e.g., GEOSECS, for used acronyms see Appendix 3) and more or less global programs from the late 80s onwards. The progress under regionalized subprograms has been mainly conducted on particular ocean basins to unravel interactions among biological, chemical and physical processes involved in the carbon cycle dynamics and the ocean role in the global climate system (e.g., JGOFS, WOCE, SCICEX, and more recently GEOTRACES). All these initiatives have been providing encouraging progress pointing to global climate forcing, coupled ocean/atmosphere interactions, and the role of biological systems as regulation filters within the concept of “the living ocean” (e.g., Demina and Lisitzin, 2013). Concerning water masses characterization, the present-day context reveals that the spatial distribution of trace-elements and isotopes is highly intricate, indicating that many details about three-dimensional distributions, the involved processes and, especially, their interactions are still insufficiently understood (e.g., Sohrin and Bruland, 2011). Transferring this challenge to the information retrieved from the near or close past has been considered (e.g., Paleo-JGOFS and later GEOTRACES). The application to increasingly older rocks poses additional and progressively greater difficulties related to the nature of the stratigraphic record, among which postdepositional processes (diagenesis) are of special relevance for interpreting site effects. From DSSP (1966–1983) to the ODP (starting on 1985) and the present IODP (beginning in 2003), transects provided with examples from offshore cores, but reported Jurassic cases were very limited, as was their correlation potential within a biostratigraphic resolution far below that obtained from land-based outcrops. The Jurassic case study investigated here aims to reveal geochemical signatures of different water masses, as evidenced by bio-geochemically different marine domains. Specifically, these water masses refer to (1) an epicontinental middle shelf, submitted to the influence of the marginal filter, and (2) a structured epioceanic fringe corresponding to the pelagic zone filter, locally exposed to episodes of changing currents and affected by hydrothermal activity. This study reveals the possibility that valuable information can be obtained on two separate levels: (1) a high variability in sedimentological details, combined with stratigraphic and geochemical records emerging from spatial versus single-site, particular or combined chemostratigraphy; and (2) the recognition of major trends in chemostratigraphic patterns from spatially distant sites revealing their high temporal correlation potential with wider than regional chemostratigraphic trends. Despite local particularities, chemostratigraphic compatibility with well established trends at larger scale points to “global” forcing determining relevant features of the Jurassic world ocean at intertropical latitudes. The present study will be followed by comparison of data retrieved from sections of the same age and outcropping in different basins close to both extremes of the proto-Atlantic seaway — the Hispanic Corridor. 7. Conclusions The geochemical analysis of Upper Jurassic epicontinental to epioceanic carbonates from a several-hundred-kilometer long proximalto-distal transect corresponding to the S-E Iberian subplate palaeomargin context, reveals: (1) Among epioceanic carbonates, CL patterns show preservation of matrix micrites and belemnite rostra, as well as of shallow marine burial precipitation of secondary carbonate cements in veinlets. In comparison, a rather complex diagenetic history occurs in ammonite shells, experiencing variable stages

(2)

(3)

(4)

(5)

(6)

23

of neomorphism and local cementation. Ammonitico Rosso geochemistry facilitated discrimination of global to wide regional signals forced by Tethys-wide palaeoceanographic– palaeoenvironmental patterns from local as well as postdepositional deviations. Stratigraphic trends in carbon and oxygen isotope composition across all epioceanic sections are consistent with major Late Jurassic palaeoceanographic events and trends. The Middle Oxfordian Tethyan transgression is recorded as a carbon-isotope maximum, related to high productivity levels. Sea-level highs are reflected in 18O-enriched values, probably indicative of relatively cooler water masses. Spatial variations in δ18O reflect a cooling trend towards the open ocean, while local differences in current patterns might be driven by relative bottom physiography and related oceanographic features. Carbon isotopic composition of belemnite rostra suggests influences of vital effects. Nonetheless, belemnite oxygenisotope ratios apparently are not affected by metabolic processes and in agreement with proposed seawater values. Of special interest are geochemical data from Ammonitico Rosso matrix micrite that are also in good agreement with expected seawater values. This despite the fact that matrix micrite is commonly considered as less reliable relative to those from belemnite rostra. Whereas epioceanic geochemical patterns are consistent with patterns previously recognized across epioceanic Tethyan margins, local/regional features are recorded in the epicontinental, inner platform domain under continental influence. Two stratigraphic features require attention: (1) maximum abundance of skeletal components coincides with positive peak values in oxygen isotope values. This pattern points to a complex relation between relative sea level, productivity, sediment entrainment and the geochemical record. (2) Silica nodule-bearing horizons refer to increased nutrient levels potentially triggered by hydrothermal pulses, as evidenced by the inverse correlation between manganese elemental abundances and oxygen isotope ratios. The data obtained represent a strong motivation for assessment of spatial patterns in chemostratigraphy. This research claims for moving the investigation of past geochemical records towards recent improvements applying a geologically more reasonable 3D dynamic palaeoceanography, which reveals intricate relationships between lithofacies and isotope signature, and promotes combined interpretations of modern marine and ancient settings.

Acknowledgments We wish to thank Dieter Buhl, Ulrike Schulte, Beate Gehnen and Noushin Arshadi (Ruhr Universität Bochum — RUB) for their support during laboratory measurements; Rolf Neuser (RUB) for CL and SEM analysis; Matthias Born (RUB) and Alberto Montes (University of Granada — UGR) for the preparation of thin sections; and Alicia González, Concepción Hernandez and Hoda Khaldy (CIC-UGR) for collaboration during SEM analysis and preparation of laboratory material. We also thank Beatriz Marques (University of Lisbon) for background information and field support in S Portugal. Editorial guidance by Susan Marriott and comments by two anonymous reviewers are acknowledged. The authors also acknowledge fruitful discussions with Eugene Rankey, Elias Samankassou and Benoit Vincent. This research was supported by Projects CGL2008-05251-E and CGL2012-39835 (MINECO) and the Research Group RNM-178 Junta de Andalucía, Spain.

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Appendix 1 Table A.1 supports Section 1.3 by providing an overview of main items, contribution of geochemistry and related topics to the interpretation of ARFC deposits. Table A1 Overview of information reported on geochemical studies performed to outcrops encasing Ammonitico Rosso deposits. Site

Time frame

Geochemistry

Main items & contribution of geochemistry and related topics to the interpretation of ARFC deposits

Selected references

Trento Plateau, N. Italy

Jurassic–earliest Cretaceous

Co, Cr, Cu, Fe, Mn, Ni, Pb, Ti

Drittenbass (1979)

Italy, Hungary and Switzerland

Early Jurassic

C-isotopes

Bakony Mts., Hungary

Pliensbachian, Toarcian, Bajocian

Al, Co, Cu, Fe, Mg, Mn, Ni, Pb, Zn. Clay, Ca and Mn mineralogy

N. Italy

Kimmeridgian– Berriasian.

C-isotopes

Veneto, N. Italy

Callovian– Kimmeridgian.

Corg, amino acids

Hungary

Pliesbachian– Bajocian

Major, minor elements

Trento Plateau, N. Italy

Late Jurassic

Biochemistry, biosedimentation, Fe, Mn,

Majorca Island, SE Spain

Kimmeridgian– Early Tithonian (?)

C- and O-isotopes, Mg, Fe, Sr, Mn

S Spain

Late Jurassic

Major & trace elements and REE

Lavarone, N Italy

Late Jurassic

C- and O-isotopes from belemnites

Aragonite dissolution forced by increasing pelagization in ARFC deposits with typical occurrence of hardgrounds, mineralized crusts and stromatolitic structures. Hardgrounds and mineralized crusts similar to recent ones. Persistent high Fe/Mn ratios, and Fe–Mn and Fe–Ti positive correlations. Nodules with Fe–Co and Cr–Pb positive correlation. Ni/Pb and Cu/Pb ratios equivalent to those in recent nodules from intermediate water depths. No reliable depth inference from element geochemistry. Improbable diagenetic origin for nodulation under oxidizing context in red ARFC limestones. Metal precipitation mediated by biotic or abiotic activity. Geochemical signatures from Lower Jurassic ARFC and black shale deposits under biozone-level biostratigraphic control. ARFC related to comparatively “normal” but diagenetically over imprinted records. Fe–Mn nodules and crust in ARFC and basinal deposits. Similarity in chemical composition in ARFC and basinal samples. Lacks relevant diagenetic effects. Comparison with modern analogs points to continental marging palaeoenvironments for analyzed fossil nodules and crusts. Oceanic seamount and abyssal modern analogs differ in element geochemistry. Late Jurassic/earliest Cretaceous δ13C curve in ARFC and basinal deposits under magneto-biostratigraphic control. Forcing from global carbon and hydrological cycles. Stromatolitic ARFC deposits. Calcium carbonate phases and aminoacid contents. Sedimentary and early diagenetic processes. Cementation and condensation relationships. Early–Middle Jurassic trace element contents in Mn-rich nodules and encrustations from seamounts (ARFC deposits) and basinal deposits. Element composition similar to modern examples. No evidence for post-depositional changes other than cementation. Organic matter content of stromatolitic growths in ARFC deposits. Amino acids known from modern cyanobacteria. Occurrence of microbial ecosystems promoting centimetric, granular, laminated fabrics (cyanobacteria) and micrometric bands of Fe–Mn concretions on bio- and lithoclasts (bacterial– fungal consortia). Cyclic growth relating episodes of soft- to hardground conditions in sediment/water interface and across their interphase. Microbial origin for metal deposition. Amino acids promoting early cementation and indirectly dissolution. Fluctuations in oxygenation and quasi-in situ deposition of Fe-oxydroxides. Dissolution forcing strong condensation in the RAV (a particular type of ARFC deposits). Skeletal geochemistry from belemnite rostra as the most reliable palaeotemperature record in Upper Jurassic ARCF deposits. Palaeotemperature links to palaeolatitude. Bulk-element geochemistry and micro-biogenic activity. Conditions for Fe–Mn crust and stromatolitic–oncolitic growths in a particular subfacies of Late Jurassic ARFC deposits. Geochemical signatures and relative sea level under biozone-level biostratigraphic control in Upper Jurassic ARFC deposits. Marine current energy, productivity, winnowing, and depositional conditions. Bulk element geochemistry in Jurassic Fe–Mn crusts related with ARFC deposits. Quantitative and qualitative analyses. Depth and oxygenation, detrital and hydrothermal inflows, and metal adsorption rates on raised sea-bottoms forced difference in Fe–Mn crusts and nodule formation. Middle Jurassic C-isotopes from raised bottoms (ARFC with, gray to pink limestones, nodular limestones and red marls on epioceanic swells) and troughs (darker gray, siliceous micritic limestones and marls, cherty limestones) belonging to distant, Tethyan areas. C-isotope curves supporting long distance correlation. Aalenian–Bathonian biozone-level control biostratigraphy. C-isotope values complement the correlation potential of combined ammonite and radiolarian biostratigraphy.

Betic Cordillera, S Spain

Latest–Early to Late Jurassic

Major and minor elements, REE, transition metals

Betic Cordillera, S Spain and Umbria–Marche– Sabina, Apennines, Italy

Middle Jurassic

C-isotopes

Jenkyns and Clayton (1986)

Mindszenty et al. (1986)

Weissert and Channell (1989) Scudeler Bacelle and Nardi (1991) Cronan et al. (1991)

Ballarini et al. (1994)

Price and Sellwood (1994) Martín-Algarra and Sánchez-Navas (1995)

Caracuel et al. (1997)

Jiménez-Espinosa et al. (1997)

Bartolini et al. (1998)

R. Coimbra et al. / Earth-Science Reviews 139 (2014) 1–32

25

Table A1 (continued) Site

Time frame

Geochemistry

Main items & contribution of geochemistry and related topics to the interpretation of ARFC deposits

Selected references

S. France

Devonian

Biogeochemistry, Fe, Mn

Préat et al. (1999)

Mont Inici, Sicily

Late Jurassic to Early Cretaceous

C- and O-isotopes

Global

Jurassic

C-, Nd., Norg-, O-, Os-, S-, Sr-isotopes, trace elements, TOC

Betic Cordillera, S Spain

Jurassic–Early Cretaceous

C- and O-isotopes

Trento Plateau, N. Italy

Callovian–Tithonian

Biogeochemistry, Fe, Mn

Trento Plateau, N. Italy

Callovian–Tithonian

Biosedimentology

Western Europe

Devonian, Jurassic

Biogeochemistry, Fe, Mn

Betic Cordillera, S Spain

Jurassic–Early Cretaceous

C- and O-isotopes

N Italy (among others)

Phanerozoic

Trace elements, Fe-isotopes

Subbetic Zone, S. Spain; Trento Plateau, N. Italy Subbetic Zone, S Spain

late-Early to Late Jurassic

Biogeochemistry, Fe, Mn

Aalenian to Early Callovian

C- and O-isotopes

Trento Plateau, N. Italy

Middle to Late Jurassic

Bioprecipitation C- and O-isotopes, Fe, Mg, Mn, Sr

Red pigmentation in limestones with low Fe & Mn contents. Biogenic activity of Fe-bacteria during the early diagenesis producing submicronic hematite. Diagenetic remobilization of hematite. Biogenic, bacterial Fe-oxidation forces red coloration irrespective of age and palaeogeography, hence predicted for ARFC deposits. Geochemical contexts for palaeoenvironmental evolution and the early diagenesis. ARFC subfacies differentiation. Bottom topography forcing of deposition. Biozone level integrated stratigraphy. Revision of applications of chemostratigraphy to Jurassic materials, including ARFC deposits. Use of multi-parameter based chemostratigraphic patterns under biozone-level biostratigraphic control. Global cycle-processes to regional interpretation of marine palaeoenvironments. Diagenesis. Isotope values from Middle Jurassic to earliest Cretaceous sediments in a pelagic trough (siliceous deposits with meager biotratigraphic control) and a swell (ARFC deposits under biostratigraphic control). Chemostratigraphic correlation. Identification of global and local geochemical signatures. Differential but typical early and burial diagenesis imprints. Red pigmentation in ARFC deposits irrespective of depth. Red pigmentation in Paleozoic and Mesozoic limestones due to Fe-oxidizing bacteria. Petrography and biosedimentology of ARFC deposits. Biodiagenessis of matrix carbonate. SEM observed microfilaments attributed to Fe-bacteria, fungi and nannobaceria. Bioprecipitation of Fe-hydroxides at O2deficient sediment–water interfaces. Microbial–fungal decay liberating submicron Fe-hydroxides later transformed in hematite forcing red pigmentation. Pigmentation relates to micrite texture, with red for subhaedral micrite grains and pink-gray for anhedral ones. Red pigmentation in sediments with Fe and oxygen contents relatively low. Neutrophile bacteria and associated microbes biomineralizing Fe and/or Mn. Releasing of Fe and Mn to encasing micrite after death. No bathymetric influence for iron bacterial activity. Comparison of Middle Jurassic to earliest Cretaceous ARFC (upper slope of swells) and more basinal, marly Berriasian deposits (troughs). Correlation of sea level and C-isotope curves differentiates regional from global environmental forcing. Synsedimentary tectonics controls palaeoenvironmental context, depositional conditions, and diagenesis (preservation of geochemical signatures of early marine diagenesis by local tectonic uplift while trough conditions those of burial diagenesis). Fe-isotopes revealing bacterial–fungal activity early during diagenesis in ARFC red subfacies. Recognition of paragenetic phases in calcite. Early diagenetic transformation of (hydro) oxides to hematite and dispersion in microaerophilitic environments. Processes occurring at centimeter depth within the sediment and below the photic zone. Fe-remobilization early during diagenesis in Condensed ARFC deposits. Petrographic analysis of early diagenetic processes. C-isotope curves from swells (ARFC deposits) vs. troughs (rhytmites). O-isotopes reflecting changes in burial diagenesis. C-isotope curves reflecting changes of the palaeocarbon cycle. Biozone level controlled main faunal turnovers and radiations follow changes in the carbon isotope record. Long distance chemostratigraphic correlation. Potential minor discrepancy in some biostratigraphic boundaries. Petrography and biosedimentology of red, pink, gray coloration in ARFC deposits. Diagenetic stratigraphy for cementation. Elemental analysis revealing early diagenesis courses and HMC to LMC change. Diagenetic speed affecting red and gray ARFC subfacies. Differential micrite texture linked to different coloration. Microbial–fungal forcing for reddish coloration. Diagenetic and biological forcing of carbonate matrix formation and alteration.

Cecca et al. (2001)

Jenkyns et al. (2002)

Rey and Delgado (2002)

Mamet and Préat (2003) Morano et al. (2003)

Mamet and Préat (2005)

Rey and Delgado (2005)

Mamet and Préat (2006a)

Mamet and Préat (2006b) O'Dogherty et al. (2006)

Préat et al. (2006)

(continued on next page)

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R. Coimbra et al. / Earth-Science Reviews 139 (2014) 1–32

Table A1 (continued) Site

Time frame

Geochemistry

Main items & contribution of geochemistry and related topics to the interpretation of ARFC deposits

Selected references

Trento Plateau, N. Italy

Middle to Late Jurassic

Fe-isotopes, Ca CO3, Fe, Mg, Mn, Sr

Préat et al. (2008)

Betic Cordillera, External Subbetic, S. Spain

Jurassic– Cretaceous

Elemental geochemistry, Al, Ce, Co, Cu, Fe, K, Mg, Mn, Na, Ni, Si, Ti, REE

Betic Cordillera, S Spain

Late Jurassic

C- and O-isotopes

Subbetic Zone, S Spain

Middle to Late Jurassic

Fe/Mn, REE

Lessini Veronese, N. Italy and W. Sicily

Middle to Late Jurassic

Biogeochemistry, Fe, Mn, Ni

S Spain and others in a Tethys Boreal correlation context

J/K boundary

Bulk vs. belemnite C- and O-isotopes, trace elements

Betic Cordillera, S Spain

Late Jurassic

C- and O-isotopes, Fe, Mg, Mn, Sr

Majorca Island, SE Spain

Late Jurassic

C- and O-isotopes, Mg, Sr, Fe, Mn

Betic Cordillera, S Spain

Late Jurassic

C- and O-isotopes, Ca, Fe, Mg, Mn, Sr

Betic Cordillera, S Spain

Late Jurassic

C- and O-isotopes, Ca, Fe, Mg, Mn, Sr, skeletals abundance

Fe-isotopic fractionation is sensitive biomarker in red dysaerobic and gray ARFC sediments. Fe-content higher in red than in gray facies. Red facies result from oxidizing Fe-bacteria and fungi. Microenvironmental factors are redox potential, pH, oxygenation degree, and hiatuses. Processes are light independent. Fe–Mn crusts. Jurassic discontinuities. Mostly ARFC deposition during sea level falls. Element geochemistry differentiating thin, deeper crusts from thicker, shallower, stromatolitic/oncolitic, and REE-rich with Ce+ anomaly, and igher Fe/Mn ratios. Crust's formation involving difference in depth, oxygenation, and mainly abiotic but locally biogenic precipitation of Fe–Mn oxydroxydes intercalating episode of fine siliciclastic sedimentation. Metals supplied by detrital and hydrothermal sources. Isotope evidence for rather conservative matrix micrite, early lithification, nodulation, and marine porewaters cementation in ARFC deposits,. Freezing of isotopic signals favored by nodulation related to current swept, discontinuous sedimentation, burrowing and irrigation. Chemostratigraphic correlation in epioceanic transect under biozone level control biostratigraphy. Ferromanganesiferous macro-oncoids in ARFC deposits. Authigenic Fe–Mn mineralization forced by chemoorganotrophic behavior in benthic microbial communities. Deep euphotic zone processes in epioceanic swells. No particular depth range for ARFC deposits. Fe and Mn as potential palaeogeographical markers for ARFC deposits. Coloration in ARFC deposits relates to O-gradients. Dysoxia and low sedimentation rates favors microbial activity and red pigmentation. Diagenetic sequence, 40% of which under biogenic influence. A Cretaceous analog for ARFC deposits. Approaching “global” correlations of the J/K boundary through combined bio-, magneto-, and chemostratigraphy. Comparison of mainly bulk geochemistry of Tethyan ARFC deposits, with bulk geochemistry of a mid-latitude hemipelagic carbonate rhythmite, and belemnite geochemistry from a Boreal shally, organic matter-rich and carbonate poor section. Possibilities for gradual warming decrease in sea water δ18O, or temperature difference in Tethys and Boreal realms if δ18O was constant. Carbon isotopes from belemnite rostra could be affected by vital effects and/or belemnite ecology. Geochemical data from carbonate skeletals vs. matrix micrite in normal epioceanic swell and volcanic depositional settings for ARFC deposits. Slightly warmer bottom waters along with increased abundance of Mg, Fe, and especially Mn, interpreted as resulting from hydrothermal influence from nearby submarine volcanic ridge. Belemnite and ammonite ecology, and belemnite vital effects assumed. Biozone-level biostratigraphic control. Rapid, thick deposition of grayish–brownish, macrofossil poor fine carbonates encased by ARCF deposits on an epioceanic swell. Geochemical signatures of matrix micrite, belemnites and cements under biozone-level biostratigraphic control. Higher than expected C and O isotope signatures, and Sr enrichment along with low Fe and Mn concentration, point to a shallower, carbonate shelf aragonitic source. Stratigraphic and spatial variations in carbonate elemental chemostratigraphy across a proximal-to-distal transect, from an epiocontinental rhythmite to epioceanic ARFC deposits. PCA and hierarchical cluster tests revealing palaeoenvironmental significance of elemental associations. Mn-decoupling related to hydrothermal activity. Potential time-fluctuation of geochemical patterns approached through variogram computation under biozone-level biostratigraphic control. Patterns of stratigraphic, temporal behavior most probably forced by tectono-eustasy and inner planets interactions. Automatic modeling of geochemical behavior in ARFC deposits along a Late Jurassic palaeomargin. Application of a machine learning method (Regression Trees) on skeletal–geochemical, multi-data sets with different statistical distributions. O-isotopes revealing waters masses traits. Positive correlation of skeletals abundance with O-isotopes points to particular sedimentary episodes in cooler waters.

Jiménez-Millán and Nieto (2008)

Coimbra et al. (2009)

Reolid and Nieto (2010)

Préat et al. (2011)

Žák et al. (2011)

Coimbra and Olóriz (2012a)

Coimbra and Olóriz (2012c)

Coimbra et al. (2014a)

Coimbra et al. (2014b)

R. Coimbra et al. / Earth-Science Reviews 139 (2014) 1–32

Appendix 2 Microfacies details providing complementary support to Section 4.1.1 are summarized as text focused on macrofossil remains, microfossils (mainly foraminifera) and grains biogenically mediated by microbial/ algal growths for the Rocha Poço section. Illustrations of selected biosedimentary fabrics (Figs. A.1 to A.3) are also given for epioceanic and epicontinental deposits in accordance to features and interpretation in the main text. Foraminifera from the lower siliciclastic interval at Rocha Poço are dominated by agglutinated and calcareous forms such as lituolaceans and nodosariaceans. In the overlying spongiolitic horizons of this section, sponge spicules frequently occur as oncoids' nuclei, and there are cases in which the nuclei correspond to relatively large recrystallized voids, hard to differentiate from surrounding pale-gray matrix typical of interstitial spaces. Microbial/algal growths include benthic foraminifera, nubeculariids, or are encrusted by them. Tubyphites and tuberoids exist, and cortoids occurs at the top of the section (Jordana Formation; Marques, 1985). Among bioclasts are common sponge frames (dictyonals) and spicules, plates, spines and ossicles of echinoderms; fragments of benthic rarely articulated and micrite coated bivalves, as well as pelagic ones (“filaments”); gastropods; rare and incomplete phragmocones of ammonites, one preserving ammonitella; and a single case of broken brachiopod showing an incomplete brachidium. Less abundant are neomorphized worm tubes of serpulids, rarely in local patches, and remains of probable filamentous algae. Scattered benthic foraminifera are calcareous nodosariaceans, Lenticulina; robertinaceans, epistominins?; rare involutinids, Trocholina; porcelanous miliolaceans such as nubeculariids, and miliolids including ecuatorial sections of Nautiloculina or Meandrospira; agglutinated forms referable to non-septate ammodiscids?; and uniserial and more scarce biserial, planispiral and trochopiral lituolaceans such as Reofax, Textularia, and possible ataxophragmiids. Appendix 3 Acronyms used in Section 6: GEOSECS Geochemical Ocean Sections Study; GEOTRACES Global marine biogeochemical cycles of Trace Elements and their isotopes; JGOFS Joint Global Ocean Flux Study; PALEO-JGOFS Task Team JGOFS-PAGES — The Joint Global Ocean Flux Study; SCICEX Submarine Arctic Science Program; WOCE World Ocean Circulation Experiment.

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