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Response of western South American epeiric-neritic ecosystem to middle Cretaceous Oceanic Anoxic Events
Accepted Manuscript Response of western South American epeiric-neritic ecosystem to middle Cretaceous Oceanic Anoxic Events J.P. Navarro-Ramirez, S. Bodin, L. Consorti, A. Immenhauser PII:
S0195-6671(16)30203-8
DOI:
10.1016/j.cretres.2017.03.009
Reference:
YCRES 3555
To appear in:
Cretaceous Research
Received Date: 9 September 2016 Accepted Date: 9 March 2017
Please cite this article as: Navarro-Ramirez, J.P., Bodin, S., Consorti, L., Immenhauser, A., Response of western South American epeiric-neritic ecosystem to middle Cretaceous Oceanic Anoxic Events, Cretaceous Research (2017), doi: 10.1016/j.cretres.2017.03.009. This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.
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Response of western South American epeiric-neritic ecosystem to middle Cretaceous Oceanic Anoxic Events
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Navarro-Ramirez J.P.1, *, Bodin S.2, Consorti L.3, Immenhauser A.1
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Germany
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Aarhus University, Department of Geoscience, Høegh-Guldbergs Gade 2, 8000 Aarhus C, Denmark
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Departament de Geologia (Paleontologia), Universitat Autònoma de Barcelona, 08193 Bellaterra,
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Spain
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Ruhr-Universität Bochum, Institut für Geologie, Mineralogie und Geophysik, D-44870 Bochum,
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* Corresponding author. Tel.: +49-234-32-2325.
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E-mail address: [email protected]
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Revised version for the Journal Cretaceous Research
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Keywords: Peru, mid-Cretaceous, Pacific, OAE1d, OAE2, chemostratigraphy, carbonates
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Abstract Little is known about the impact of the mid-Cretaceous Oceanic Anoxic Events (OAEs) on the
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neritic carbonate systems in South America. In order to fill this knowledge gap, the present paper
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reports on the record of environmental changes in the Albian–Turonian neritic carbonates from the
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western South American domain in Peru. Owing to the very expanded and well-exposed sections in
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the Oyon region of central Peru, the OAE 1d and 2 intervals were sampled at high temporal resolution
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for both bulk micrite and bulk organic matter carbon isotopes, allowing us to compare the fingerprint
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of these two events between the northern and central Peruvian regions. This suggests the installation of
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two marked depositional modes: 1) the Albian–Turonian formation of a regional facies belt constituted
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by oyster-rich mixed siliciclastic-carbonate deposition along the western South America platform; 2) a
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restricted oligotrophic environment, characterized by the mass occurrence of Perouvianella peruviana
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and associated miliolids in central Peru during the late Cenomanian–Turonian. These observation
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advocate for the following scenario: Global warming during the late Albian–early Turonian resulted in
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humid climate on the western platform. This in turn caused enhanced chemical weathering rates on the
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Brazilian Shield, resulting in high runoff of nutrients onto the western platform. Nutrient runoff
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promoted the diversification of benthic oyster communities. Due to the uplift of the Marañon Massif
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and the installation of the Huarmey Trough, central Peru was isolated from the Pacific and from
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eastern deltaic influx of the Brazilian continental basement, allowing the local development of
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oligotrophic conditions during OAE 2. Furthermore, an increased influx of argillaceous sediment and
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reduced carbonate production is recorded in northern Peru at the onset of OAE 2, marked by a
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prominent negative shift in δ13C. This negative carbon-isotope excursion has also been identified in
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other sections in the Pacific domain and can be linked to an increase in isotopically light pCO2 induced
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by the formation of the Caribbean large igneous province.
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1. Introduction The mid-Cretaceous (i.e., the Albian–Turonian interval, ca. 110–90 Ma; Ogg and Hinnov, 2012) is
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considered as a time governed by intensification of greenhouse conditions due to elevated emission of
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pCO2 to the atmosphere following massive pulses of ocean crust production and the formation of large
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igneous provinces (LIPs; Arthur et al., 1985; Barron and Washington 1985; Larson, 1991; Leckie et
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al., 2002; Snow et al., 2005; Bodin et al., 2015). The intensified greenhouse conditions enhanced the
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hydrological cycle, which in turn accelerated continental weathering rates (Berner et al., 1983;
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Weissert, 1990; Weissert et al. 1998; Hay, 1998; Menegatti et al., 1998; Wortmann et al., 2004; van
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Helmond et al., 2013; Bodin et al., 2015) and led to increased fluxes of continental nutrients to the
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oceans. This chain of events stimulated biological productivity, leading to widespread organic-carbon
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deposition and enhanced marine anoxia and euxinia (Manabe and Bryan, 1985; Föllmi et al., 1994,
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2006; Turgeon and Brumsack, 2006; Owens et al., 2013; Pogge von Strandmann et al., 2013; Lechler
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et al., 2015). The processes described above probably occurred on some level throughout the
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Cretaceous, but became particularly enhanced during several geologically short intervals (<1 Myr),
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known as oceanic anoxic events (OAEs). These latter are best defined based on carbon isotope (δ13C)
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excursions that reflect a fundamental perturbation in the global carbon cycle (Erbacher et al., 1996;
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Leckie et al., 2002). The best-studied OAEs in Europe, North America, and from Atlantic ODP sites
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include the early Aptian OAE1a, the Albian OAE1b, OAE1c and OAE1d, and the Cenomanian–
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Turonian OAE2. Additional events, such as the Coniacian OAE3 and the mid-Cenomanian event I
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(MCEI) represent related phenomena that lack widespread organic carbon-rich facies and/or a
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significant δ13C excursion (e.g., Jarvis et al., 2006; Jenkyns, 2010).
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OAE1a and OAE2 are conventionally interpreted as the result of massive outgassing of the
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Ontong Java, Manihiki and Hikurangi LIP (Tarduno et al., 1991; Larson and Erba, 1999; Méhay et al.,
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2009; Tejada et al., 2009) and by the Caribbean LIP (Snow et al., 2005; Du Vivier et al., 2014). At the
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onset of OAE1a and OAE2, a negative shift in lithium and calcium isotope ratios is observed, most
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likely associated to weathering of continental crust (Blättler et al., 2011; Pogge von Strandmann et al., 3
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(southeastern France), and Eastbourne (England) sections, suggested that the Caribbean LIP triggered
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a change in the predominant type of biomineralization (i.e., ocean acidification) at the onset of OAE2
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(Du Vivier et al., 2015). As recorded in the geological record, transient periods of high atmospheric
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CO2 led to several mass extinctions, most likely because excess CO2 reduced the ability of carbonate-
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secreting organisms to secrete their carbonate shells (e.g., Glikson, 2010). In a more qualitative way,
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sedimentological studies suggest enhanced weathering at the onset of OAE 1a and OAE2, as reflected
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by the relative abundance of kaolinite (Stein et al., 2012; Gertsch et al., 2010).
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Evidences of a sudden influx of siliciclastic material coeval to OAEs have been reported in the
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North Atlantic and the Tethys domains (Weissert, 1990; Wortmann et al., 2004) and are indicative of a
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change to more humid and warmer conditions (Weissert, 1990; Weissert et al., 1998; Wortmann et al.,
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2004). This concept is exemplified in the evolution of the northern Tethyan platform during the Early
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Cretaceous, where carbon isotope (δ13C) excursions occurred at times of elevated nutrient levels due to
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intensified continental weathering rates (Erba, 1994; Weissert et al., 1998; Immenhauser et al., 2005;
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Bodin et al., 2006; Föllmi et al., 2006; Huck et al., 2010, 2011, 2012, 2013, 2014). This caused a
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change from oligo- to meso- or eutrophic conditions, ecological reorganization, phases of carbonate
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platform development and drowning episodes (Weissert et al., 1998; Immenhauser et al., 2005; Bodin
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et al., 2006; Föllmi et al., 2006; Huck et al., 2011). However, these hypotheses have yet to be applied
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to sections in South America, which record the response of shallow-marine carbonate depositional
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environments of a southern hemisphere continent adjacent to the Pacific Ocean.
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The western South America platform (Fig. 1) was characterised by a large oceanic basin to the
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west and a very large uplifted continental basement area to the east. The enormous dimensions of this
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region make this a challenging natural laboratory to understand southern hemisphere Earth System
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behaviour under extreme climates. The aim of this paper is to document and discuss the Albian to
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Turonian sedimentological and palaeoecological evolution of the epeiric-neritic ecosystem of the
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western platform in central Peru (Oyon region; Jaillard, 1986; Fig. 2). The assessment of potential 4
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published northern Peru reference composite section (Cajamarca region, Fig. 3, Navarro-Ramirez et
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al., 2015, 2016). Specific attention is paid to environmental stressors associated to OAE1d and OAE2
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and the manner in which they affect the main carbonate producers such as oysters, gastropods,
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echinoids, and endemic large benthic foraminifera (Perouvianella peruviana, Steimann) that have so far only been reported from these sections (Jaillard and Arnaud-Vanneau, 1993).
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2. Regional tectonic and stratigraphic setting
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The South America margin was largely influenced by late stages of Gondwana breakup that
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culminated in the separation of Africa and South America and the opening of the Equatorial Atlantic
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Ocean during the Aptian–Albian transition (Moulin et al., 2010; Fig. 1). This allowed for the
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connection of North and South Atlantic Ocean water masses towards the Turonian (Eagles, 2006).
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This tectonic event caused the activation of the western South America volcanic arc in the Early
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Cretaceous (Huarmey-Trough, Atherton and Webb, 1989; Soler and Bonhomme, 1990; Jaillard et al.,
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1999, 2000; Winter et al., 2010) and subduction took place along the western portion of South
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America where sedimentation was controlled by NNW–SSE-trending structures (e.g., Benavides-
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Caceres, 1956; Jaillard, 1986, 1987; Jaillard et al., 1990; Callot et al., 2008; Robert et al., 2009).
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Mid-Cretaceous sedimentation in what today is Peru began with the deposition of siliciclastic
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rocks now forming the Inca Formation, which is built by iron-rich, sandy beds, assigned to an early
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Albian age based on the ammonite Neodeshayesites nicholsoni (Robert and Bulot, 2004; Robert et al.,
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2009). This siliciclastic unit in turn is overlain by Albian carbonates represented by the Chulec, and
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Pariatambo formations, recording a second-order Albian transgression (Benavides-Caceres, 1956;
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Jaillard, 1986; Robert, 2002; Robert and Bulot, 2004, 2005).
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The Chulec Formation unconformably overlies the Inca Formation with a discontinuity surface
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separating the two units (Jaillard, 1987). The Chulec Formation is characterized by marl-limestone 5
ACCEPTED MANUSCRIPT alternations with a very abundant and diversified outer shelf fauna (Jaillard, 1987). Numerous
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ammonites have been reported and were assigned to the Knemiceras raimondii Zone (Robert and
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Bulot, 2004; Robert et al., 2009), indicating an early Albian age. The Pariatambo Formation is
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characterized by fossiliferous, black, bituminous, fetid marly limestones and includes finely laminated
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and siliceous intervals. Ammonites and planktonic foraminifera are common. The presence of
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Oxytropidoceras carbonacrium and Prolyelliceras ulrichi Zones indicate an early middle Albian age
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(Robert et al., 2009). The Pariatambo Formation grades upsection into nodular bioclastic grey and
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marly limestones of the Yumagual Formation in the northern Peru and by the Jumasha Formation in
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central Peru (upper middle Albian– lower Turonian; Fig. 3).
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At Uchucchacua (Oyon region; central Peru; Jailard, 1986) the Jumasha Formation is about 1320
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m thick and made of very massive, thickly-bedded, light yellowish brown to brownish-grey limestones
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(Jaillard, 1986; Fig. 3). The base of the Jumasha Formation was attributed to the upper middle Albian
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due to ammonite associations including Lyelliceras ulrichi Knechtel and Oxytropidoceras douglasi
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(Benavides-Caceres, 1956; Wilson, 1963; von Hillebrandt, 1970; Jaillad, 1986; Fig. 3). At section
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metre 115 (Fig. 3) i.e., the base of the Jumasha Formation near Uchucchacua, Romani (1982) reported
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Exogyra costagyra cf. olisoponensis and Merlingina cretacea of Cenomanian age. Further south, in
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the upper part of the Jumasha Formation microfauna data and particularly Favusella washitensis, H.
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delrioensis, Globigerinelloides bentonensis and Heterohelix seewashitensis, indicate a middle to late
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Cenomanian age (von Hillebrandt, 1970; Jaillard, 1986; Fig. 3). Rotorbinella mesogeensis (Tronchetti)
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is equally present and supports the Cenomanian age of this unit. The presence of Coilopoceras sp. at
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the top of the Jumasha Formation suggests a Turonian age (Romani, 1982; Jaillard, 1987; Fig. 3).
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It must be noted that, due to scares ammonites and planktonic foraminifera findings in the
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Peruvian sections, the biostratigraphic scheme established by the aforementioned authors remains
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incomplete in places. There are uncertainties about the exact position of the stage and substage
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boundaries, although the general scheme has been confirmed by carbon isotope chemostratigraphic
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correlation as well as Sr-isotope chronostratigraphy (Navarro-Ramirez et al., 2015, 2016). 6
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3. Methods and materials
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3.1 Field work and thin-section microscopy Two well-exposed sections (Uchucchacua and Lauricocha localities; Fig. 2) were chosen for this
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study to provide maximum coverage of the Albian–Cenomanian and Cenomanian–Turonian
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transitions (Fig. 3). In total, ca. 620 m of section have been logged and studied for their
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sedimentological, stratigraphic, and chemostratigraphic archive (Figs. 3 and 4). Outcrop observations,
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rock samples, and thin-sections provide the fundament for petrographic and facies interpretations
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presented here and allow us to place the chemostratigraphic results in a facies context. In analogy with
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the nomenclature and approach presented in Navarro-Ramirez et al. (2015, 2016), non-skeletal
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components and skeletal components were analysed semi-quantitatively with numbers indicating their
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relative abundance: 0 = absent, 1 = present, 2 = frequent, 3 = abundant, 4 = dominant; Table 1). Facies
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pattern and discontinuities acted as pining points for sequence stratigraphic interpretations (see also
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Christ et al., 2015).
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3.2. Carbon isotope stratigraphy
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3.2.1.
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Bulk micrite data (δ13CCarb)
Carbon-isotope analysis of 181 micrite samples was performed using a Thermo Finnigan MAT
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delta-S mass spectrometer at the isotope laboratory of the Ruhr-University Bochum (see Appendix A),
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Germany. In order to reveal the variability of carbon-isotope composition within a single rock sample,
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several sub-samples were drilled from 20% of all the hand specimens collected. Carbonate bulk rock
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specimens were sampled by use of tungsten drill bits. While drilling the carbonate powder, areas rich
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in spary cement, large bioclasts and diagenetic calcite vein material were avoided. Refer to
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Immenhauser et al. (2005) for more details of the analytical procedure. Repeated analyses of certified
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carbonate standards (NBS 19, IAEA CO-1 and CO-8) and internal standards show an external 7
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reproducibility of ≤0.06‰ for δ13Ccarb. The δ13Ccarb values are expressed on a per mil (‰) basis relative
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to the Vienna-Pee Dee Belemnite standard (V-PDB).
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Bulk organic matter data (δ13Corg)
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Carbon-isotope analyses were performed on 201 bulk organic matter samples (see Appendix B).
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Two grams of sample powder were placed in 50mL centrifuge tubes and 6N HCl was added, this
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procedure was repeated until no carbonate reaction was visible any more. The samples are then
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centrifuged and the supernatant removed, the residues were rinsed with distilled water via centrifuge
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and dried at 40 °C. The δ13Corg measurements were performed with an elemental analyser (CE 1110)
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connected online to ThermoFinnigan Delta V Plus masspectrometer. All carbon isotope values are
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reported in the conventional δ‐notation in permil relative to V-PDB (Vienna‐PDB). Accuracy and
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reproducibility of the analyses was checked by replicate analyses of international or laboratory
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standards. Organic carbon analyses were conducted at the stable isotope laboratory of the Friedrich-
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Alexander University of Erlangen-Nuremberg, Germany.
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4. Data description and interpretation
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The Uchucchacua section is here described as the representative case example for the Albian to
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Turonian sedimentary record in the study area (Jaillard, 1986; Jaillard and Arnaud-Vanneau, 1993;
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Figs. 3 and 4). This choice is motivated by the fact that the outcrop conditions are excellent and
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stratigraphically near-complete (Figs. 5A–C). In addition, the Lauricocha section, which is of
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comparably high outcrop quality, serves as a complementary section allowing to separate local from
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regional facies changes during the late Cenomanian to early Turonian transition (Fig. 5D).
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4.1 Facies associations and depositional environments By integrating previous work (e.g., Jaillard, 1986, 1987; Jaillard and Arnaud-Vanneau, 1993;
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Dhondt and Jaillard, 2005; Jaillard et al., 2005) with our field data, four main depositional
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environments, each with a number of standard facies types are established for the intertidal to outer
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ramp domain (Table 1). The documentation of these fundamental sedimentological data is important to
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match the type of carbonate producers versus OAEs.
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Please note that the first depositional environment is represented by facies 1a, 1b and 1c, where
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facies 1a is the most proximal deposits composed of sandy marls with scarce fauna (Fig. 6A).
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Whereas, facies 1b and 1c, studied by Jaillard (1986) for the lower Cenomanian in the Uchucchacua
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section, are generally represented by beds of dolostones with sedimentary structures like tepees,
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birdeyes and mudcracks (facies 1b) and by algal laminites and lenticular structures (facies 1c),
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corresponding all to a intertidal zone setting (please refer to Jaillard (1986)´s work for more details).
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Shallow subtidal inner ramp setting
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The Albian facies types recognized at Uchucchacua display important similarities with those
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previously described from the Cajamarca region (Northern Peru, Navarro-Ramirez et al., 2015, 2016)
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and are only briefly summarized here (Fig. 4 and Table 1). At Uchucchacua, facies 2a is composed of
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grey argillaceous mudstones with scarce fauna, whereas in more distal sub-environments, facies 2b is
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characterized by packstones and argillaceous wackestones, exhibiting a nodular fabric due to an
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increased argillaceous content. Facies 2c is characterized by grainstone and occasionally floatstone
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and is typically rich in mixed and fragmented shell debris, mainly derived from oysters, gastropods,
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and echinoids (Table 1).
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The Cenomanian–Turonian facies types recorded at Uchucchacua and Lauricocha show also strong similarities as represented by the Facies 2d–h and 3b and 3c (Table 1).
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ACCEPTED MANUSCRIPT Facies 2d — Peloid-bearing grainstones (Table 1): This facies is composed of light-grey peloidal
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grainstones, showing cm to dm-thick lamination. The fauna yields some disperse miliolids, ostracods
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and sparse Perouvianella peruviana (Fig. 7A) as well as Charophyte gyrogonids. In view of the
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predominance of peloids, disperse miliolids and the low diversity of microfauna, facies 2d is
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interpreted to reflect a restricted lagoon environment (Gräfe, 2005) with limited water circulation
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(Masse et al., 2003) and low trophic levels (Hüneke et al., 2001).
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Facies 2e — Discorbids and ostracods wakestone/mudstone (Table 1): This facies is primarily
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composed by small debris (layers) of discorbid and ostracod valvae accumulations whilst rare
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miliolids are also present (Fig. 6B). This basic facies type has been described first by Sartoni and
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Crescenti (1962) from the Mesozoic of Italy and was interpreted as an environment characterized by
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restricted water circulation, medium to high trophic levels and low salinity.
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Facies 2f — Dasycladales-bearing pack-grainstones (Table 1): This facies is built by dm- to m-
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thick bedded, bluish-grey packstones and occasionally by grainstones. The rocks are typically
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characterized by large rounded bioclasts composed of dasycladales remains (Figs. 7B and 8A–B). The
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main benthic foraminifera are Perouvianella peruviana and miliolids whose tests are often abraded.
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Occasionally peloids are present. Distribution of dasycladales may occur in shallow shoals, lagoonal,
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semi-restricted very shallow environment and backreef environments (Flügel, 2004).
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Facies 2g — Miliolid-bearing pack-grainstones (Table 1): This facies is made of bluish-grey
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packstones to grainstones, displaying dm-thick lamination in units of two to five metres thick (Fig.
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6C). Bioturbation features are abundant. The fauna is dominated by miliolids, being associated to
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Perouvianella peruviana (Fig. 7C) and numerous dasycladal algae. Peloids also occur in certain levels.
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Associations of coated grains, skeletal elements such as miliolids, Perouvianella peruviana, and
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dasycladales indicate deposition in the nearshore shallow zone (Flügel, 2004; Lézin et al., 2012).
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Modern miliolids live mainly in warm, clear water lacking significant fresh-water influx and in some
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places they are associated with reefal settings (Fang, 2003).
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is characterized by grainstones, rudstones, and floatstones containing numerous Perouvianella
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peruviana (Fig. 7D). Facies 2h builds thickening upwards sequences, each one being 20 to 80 m in
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thickness. Typical biota include miliolids, Rotorbinella mesogeensis (Figs. 8C–D), agglutinated (Figs.
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8E and G), porcelaneous foraminifera (Fig. 8F) and Perouvianella peruviana (Figs. 8H–I). According
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to previous workers, this facies may indicate deposition in the external part of the inner platform (<20
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m; Stein et al., 2012), located in the platform margin (Flügel, 2004) with specifically clear and well-
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illuminated waters and low trophic levels (Lézin et al., 2012).
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Open marine middle ramp setting
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In the Central Peru upper Albian interval, facies types 3a, 3b, 3c, and 4a are near-identical to those
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identified at the Piedra Parada location (Northern Peru) as previously describe in detail in Navarro-
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Ramirez et al. (2016; Table 1). For the sake of brevity, we only report the most important
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characteristics here. Facies 3a is mainly characterized by oyster bioherms, interbedded by rare
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grainstone units yielding various skeletal elements and intraclasts. Facies 3b (diverse fauna low
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energy) in the lower Jumasha Formation is characterized by 5 to 10 cm-thick layers with variably
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sized fragments of bryozoans, echinoids, and planktonic foraminifera. Whereas, this facies in the
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upper Jumasha Formation, occurs as dm to m-thick bedded grey pack to grainstones and occasionally
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by floatstones (Fig. 6D). This here includes mixed and fragmented shell debris mainly consisting of
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Perouvianella peruviana, miliolids, gastropods, spicules and rare charophyte gyrogonites (Fig. 7E).
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Dominance of intraclasts appear also in this facies, which are usually interpreted to be formed by
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storm wave erosion occurring in shallow-marine environments (Flügel, 2004). Following previous
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workers, this type of deposits indicate deposition in high-energy, shallow subtidal shoals possibly also
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in the form of storm event beds (Elrick et al., 2009), just beneath the fair-weather wave base (Amini et
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al., 2004; Stein et al., 2012; Lézin et al., 2012).
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organized in higher order units of two to five metres in thickness and is rich in echinoid debris (Figs.
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6E and 7F). These features suggest a middle ramp environment and sediment deposition beneath the
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fair-weather wave base (Flügel, 2004; Stein et al., 2012). This facies 3c appears also in the upper
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Cenomanian sections at Uchucchacua and Lauricocha.
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Facies 4a (planktonic foraminifera-bearing mudstones) is characterized by dark grey mudstones
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with mm to cm-thick lamination and a generally undisturbed horizontal bedding. The fauna includes
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planktonic foraminifera and small echinoid fragments. These evidence an open marine setting below
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the reach of the effective storm wave base (Lézin et al., 2012; Stein et al., 2012).
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4.2 Sequence stratigraphic interpretation
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A basic interpretation of the sequence stratigraphic scheme of Central Peru was presented in
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Jaillard (1986, 1987), Jaillard and Arnaud-Vanneau (1993) and Jaillard et al. (2005) and for the
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Northern Peru by Navarro-Ramirez et al. (2015, 2016), respectively. Following these approaches, we
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tie the sequence stratigraphic for both regional settings as indicated in figure 3, where the Jumasha
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Formation in Uchucchacua is divided in 5 large-scale sequences (Jumasha 1–5; Fig. 3). Here, the
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upper Albian to lower Cenomanian is located in the first Jumasha sequence and the upper Cenomanian
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to lower Turonian in the Jumasha 4 (Figs. 9 and 10). This is feasible as based on biostratigraphic data
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from the northern and central carbonate ramp of Peru (Benavides-Caceres, 1956; Wilson 1963; von
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Hillebrandt, 1970; Romani, 1982; Jaillard 1986, 1987; Jaillard and Arnaud-Vanneau, 1993; Robert
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2002; Robert and Bulot, 2004, 2005; Dhondt and Jaillard, 2005; Jaillard et al., 2005; Robert et al.,
289
2009).
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The sequence stratigraphy approach used here is based on the work of Embry (2009) and Andrieu
291
et al. (2015) for shallow marine environments. Essentially, depositional sequences recognized are
292
bounded by sequences boundaries (SB), which represent either maximum regressive surfaces or 12
ACCEPTED MANUSCRIPT transgressive surfaces (van Wagoner et al., 1988). These surfaces indicate a significant shift from a
294
shallowing-upward to a deepening-upward pattern (Embry, 2009; Andrieu et al., 2015). Moreover, a
295
depositional sequence is a cycle that comprises: transgressive deposits (TD) marking a change to
296
upwards deepening facies (retrograding structures); a maximum flooding interval (MFI), a portion of
297
the section that marks the transition from deepening-upward to shallowing-upward patterns; and a
298
highstand deposit (HD), representing shallowing facies grading upwards into more proximal facies
299
(prograding structures).
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The lowermost medium-scale sequence at Jumasha 1 (Fig. 9) consists of thickening-upward units
301
of argillaceous, nodular limestone built by facies 3a rich in oysters. This pattern may represent a
302
highstand system tract. These units were attributed to the upper middle Albian due to ammonite
303
associations such as Lyelliceras ulrichi, Oxytropidoceras douglasi (Benavides-Caceres, 1956; Wilson,
304
1963; von Hillebrandt, 1970; Jaillad, 1986). The two overlying medium-scale sequences (Fig. 9) show,
305
in their transgressive deposits, thickening-upward transitions ranging from facies types 2c, 3a, to
306
facies 3c. These units are followed by deepening trends to outer ramp sedimentation typified by facies
307
4a with abundant planktonic foraminifera, defining the maximum flooding interval in these two
308
sequences. The highstand deposits are represented by thickening-upward successions from facies 3c to
309
3a, increasingly enriched in quartz grains (facies 1a), oysters and serpulids. At section metre 115 m,
310
measured from the base of the Jumasha Formation at Uchucchacua, Romani (1982) reported Exogyra
311
costagyra cf olisoponensis and Merlingina cretacea of Cenomanian age. Echinoderm and oyster
312
limestones characterize the last sequence (122 to 156 m; Fig. 9).
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The upper Cenomanian–lower Turonian portion at Uchucchacua and Lauricocha is comprised in
314
the Jumasha 4 (Fig. 10), where is showed three medium-scale sequences (MsS1–3). These are part of
315
the three depositional sequences Ro1–3 sensu Jaillard and Arnaud-Vanneau (1993). The medium-scale
316
sequences 1–2 with thickness of 60 to 140 metres (MsS1–2; Fig. 10) build the Jumasha 4. The
317
sequence boundary that limits the top of the Jumasha 4 (Fig. 6F), displays a prominent stratigraphic
13
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feature, defining a major change in depositional style from a regressive to a major transgression event
319
in the early Turonian (Jaillard and Arnaud-Vanneau, 1993; Fig. 10). At the base of both sections, an abrupt deepening trend from nearshore restricted (facies 2d–h) to
321
middle ramp facies (3b) is observed through medium-scale sequence 1 and may define a transgressive
322
system tract in terms of the Jumasha 4. The maximum of this deepening interval, situated at around
323
section metre 35 in both sequences, is marked by an increase in facies associations 3c. This is
324
followed by a shallowing-upward trend recorded in the medium-scale sequences 1, indicated by a
325
thickening-upward transition facies 3c upwards 2d and presence of rotaliid foraminifera Rotorbinella
326
mesogeensis, representing the regressive trend of the medium-scale sequence 1. This is followed by a
327
deepening trend represented from facies 1a, 2d to facies 3b, which form the transgressive deposits of
328
the medium-scale sequence 2. The regressive deposits of the medium-scale sequence 2 are upsection
329
indicated by the appearance of facies 2h to 2d. Medium-scale sequence 3 (MsS3, Fig. 10) is at least 40
330
m thick (but not fully exposed). This sequence includes the lower Turonian and the Jumasha 5. Facies
331
are dominated by 3c, which is well-defined in the lower part of both sections, showing a middle ramp
332
environment rich in echinoids. The transition from facies 3c into the overlaying thickly-bedded
333
limestones of facies 2d is gradual, marking the highstand system tract.
335
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4.3 Carbon-isotope stratigraphy
The upper Albian to lower Cenomanian bulk micrite and organic carbon isotope stratigraphy at
337
Uchucchacua is documented in figure 9. At the lowermost part of the curve, the δ13Ccarb and δ13Corg
338
values vary between –0.3 and +2.7‰ and between –27.4 and –25.3‰. These are followed by a
339
significant positive shift from –0.2 to +3.1‰ (δ13Ccarb) and from –27.4 to –24.2‰ (δ13Corg), showing a
340
prominent positive chemostratigraphic feature. From about section metre 90 onwards (Fig. 9),
341
values decrease reaching 1.3‰ in δ13Ccarb and –26.8‰ in δ13Corg, followed by increasingly
342
enriched values reaching +2.9‰ and –24.2‰, respectively. The top of the prominent positive isotope
343
interval ends with a declining trend to a base level of +1.1‰ in δ13Ccarb and –26.7‰ in δ13Corg. 14
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13
C-
13
C
ACCEPTED MANUSCRIPT Based on existing biostratigraphic data in Central Peru (von Hillebrandt, 1970; Romani, 1982;
345
Jaillard and Arnaud-Vanneau, 1993) and the chemostratigraphic terminology used for OAE2 interval
346
(onset, trough, plateau, and recovery, c.f., Paul et al., 1999; Tsikos et al., 2004), and datum levels A,
347
B, C and D as used in Tethyan and North Atlantic sections (Meyers et al., 2012; Du Vivier et al., 2014,
348
2015), we characterize specific sub-intervals of the Cenomanian–Turonian isotope excursion for the
349
Uchucchacua and Lauricocha sections (Fig. 10).
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In terms of their δ13C amplitudes and trends, the Uchucchacua and Lauricocha isotope curves are
351
directly comparable. In both sections, the lower part of the prominent carbon isotope excursion related
352
to OAE2 was not covered due to a gap in exposure. However, the OAE2 interval at Uchucchacua
353
displays a significant rise from mean values of +3.7‰ to a first peak of +4.6‰ in δ13Ccarb and a high
354
ratio –22.6‰ in δ13Corg, whereas at Lauricocha the δ13Ccarb curve shows
355
specific segment is here interpreted to represent the upper part of the onset of OAE2. Upsection at
356
both locations, both bulk micrite and organic carbon ratios are characterized by decreasing 13C-values
357
of 3.3‰ in δ13Ccarb and –25‰ in δ13Corg, followed by a rapid trend to
358
+4.2‰ and –22.1‰, respectively. This pattern corresponds with a `trough´ according to the Tethyan
359
scheme (Du Vivier et al., 2014; Fig. 10). Thereafter, a plateau interval of δ13Ccarb and δ13Corg values
360
ranging between +1.78 to +3.3‰ and –23 to –20.9‰ is recorded. The recovery interval is
361
characterized by a rapid decline to a base level of +2.5‰ in δ13Ccarb and –25.9‰ in δ13Corg.
13
13
C depleted values. This
C-enriched values reaching
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363
5
Discussion
364
5.1 Environmental evolution of the Western Platform from the Albian towards the Turonian
365
Jaillard et al. (1990) suggested that the opening of the North Atlantic (Western Tethys) in the
366
Early Cretaceous lead to the formation of a Tethyan oceanic seaway that extended through the Palaeo-
367
Pacific realm. This interoceanic communication may explain the Tethyan affinities of some Pacific
368
provinces during the Aptian to Turonian (e.g., Iba and Sano, 2007; Robert, 2002; Robert and Bulot, 15
ACCEPTED MANUSCRIPT 2004; Dhondt and Jaillard, 2005). According to Robert (2002) and Robert and Bulot (2004), the
370
earliest Albian was dominated by migration of Tethyan Engonoceratidae specimens to the Western
371
Platform due to an Albian transgression (Robert, 2002). Cretaceous Tethys taxa have also been
372
recorded in the Berriasian to early Albian of the north-western Pacific, perhaps evidencing a Tethyan
373
biotic realm in the Pacific in that time (Iba and Sano, 2007; Robert and Bulot, 2004). In addition to
374
this palaeontological observation, bivalve taxa of the Albian-Cenomanian, recorded at Pongo de
375
Rentema in Peru, are characterized by taxa known from Tethyan faunas, North Africa, southwestern
376
Europe, and Texas (Dhondt and Jaillard, 2005). It is thus conceivable that water masses from the
377
North Atlantic (Western Tethys) reached the Western Platform during the Albian-Cenomanian
378
(Robert, 2002; Jaillard et al., 1990; Dhondt and Jaillard, 2005).
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Moreover, in the earliest middle Albian, the South America margin witnessed the onset of the
380
Huarmey-Trough volcanism, leading the depositions of large volumes of basaltic pillow lavas
381
(Atherton and Aguirre, 1992; Winter et al., 2010) . This volcanic activity has been temporally
382
associated to an increased convergence rate of the rifting of South America and Africa (Soler and
383
Bonhomme, 1990; Moulin et al., 2010; Winter et al., 2010). The uplift of the Huarmey-Trough may
384
have also isolated the Western Patform from vigorous exchange and mixing with water masses of the
385
Pacific Ocean in the middle Albian (Fig. 11A). The rifting of South America and Africa continued
386
through the Cenomanian, forming continuous deep-water connection between the North and South
387
Atlantic Ocean (Wagner and Pletsch, 1999; Poulsen et al., 2003). This tectonic event triggered a major
388
reorganization of the global shallow and deep oceanic circulations and regional climate changes
389
(Poulsen et al., 2003). Another feature related to this pattern is the presence of new species of
390
ammonites, bivalves, and benthic foraminifers, which prospered in shallow-marine environments in
391
the Pacific (e.g., Chinzei, 1986; Toshimitsu et al., 1990; Jaillard and Arnaud-Vanneau, 1993; Robert,
392
2002; Robert and Bulot, 2004; Dhondt and Jaillard, 2005; Iba and Sano, 2007).
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These new Peruvian species are exemplified by Glottoceras moorei and Engonoceras gr. stolleyi-
394
hilli Böhm from the middle Albian (Robert and Bulot, 2004). Bivalve communities also prospered in 16
ACCEPTED MANUSCRIPT the Western Platform in the late middle Albian to late Cenomanian (Dhondt and Jaillard, 2005). The
396
lack of fossil remains of reef constructers (Jaillard, 1987) suggests the establishment of a heterozoan
397
ramp devoid of any morphological features such as a rimmed platform morphology (Fig. 11A). The
398
absence of a rim or barrier may have caused rather uniform facies belts across wide portions of the
399
Western Platform, where sediment and organic material were transported and re-distributed by storm
400
action and currents (Navarro-Ramirez et al., 2015). The presence of argillaceous material in facies 1a
401
and 2a-b points to continent-derived sediments probably transported by rivers. According to Jaillard
402
(1987), during the late Albian to late Cenomanian, a diachronic progradation of terrigenous supratidal
403
facies was accentuated, resulting in a NE-oriented deltaic system fed by runoff from the Brazilian
404
continental basement (Fig. 11A). This may indicate the installation of high trophic levels and more
405
humid climate with all corresponding changes in sediment runoff and type on the Western Platform
406
(Fig. 11A).
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In the late Cenomanian towards the early Turonian, a key feature of Jumasha 4 recorded at
408
Uchucchacua and Lauricocha (central Peru) is the ubiquitous presence of Perouvianella peruviana-
409
dominated carbonate facies (Fig. 10). During this period, central Peru was isolated from the Pacific
410
and from eastern deltaic influx of the Brazilian continental basement due to the uplift of the Marañon
411
Massif during the early Cenomanian (Jaillard, 1986; Fig. 11B). Evidences of this tectonic event is
412
given by a shallowing maximum, associated with tectonic breccias and synsedinemtary normal faults,
413
occurring from the lower to the middle Cenomanian (Jaillard, 1986). Furthermore, the presence of a
414
morphological barrier is given by the lack of an argillaceous material in these sections (Jaillard, 1986,
415
1987), probably implying the installation of an oligotrophic environment (Fig. 11B). This later
416
palaeoceanographic feature might have triggered – or at least favoured – the mass occurrence of
417
Perouvianella peruviana and associated foraminifera, being only known from these Central Peruvian
418
rocks (Jaillard, 1986 and references therein). In contrast to northern Peru (Pongo de Rentema), benthic
419
taxa are represented by inoceramids such as Mytiloides mytiloides, M. opalensis, and Sergipia sp.
420
(Dhondt and Jaillard, 2005). The genus Sergipia has a wide distribution, being reported in Brazil,
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17
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Nigeria, Mexico, Japan, and France (Hessel, 1988; Jolet et al., 1997; Dhondt and Jaillard, 2005), and
422
indicates the installation of an open marine setting in northern Peru (Fig. 11B). Perouvianella peruviana is a porcelaneous benthic foraminifera with a complex shell architecture
424
and flattened shape. The divisions of chamber lumen below the lateral wall may indicate the loci for
425
endosymbionts. In recent oceans, the large discoidal porcelaneous Archaias hosts clorophycean algae
426
symbionts in their shells which restrict their habitat to the first 20 m of the water column. Besides,
427
Marginopora, Amphisorus and Sorites possess dinophycean algae as symbionts allowing them to
428
reach a life depth of about 50 m. (Leutenegger, 1984). On the level of a working hypothesis, similar
429
environmental conditions are inferred for Perouvianella peruviana. Concerning the reproduction cycle
430
of Perouvianella peruviana, in the fourth Jumasha sequence, two A-forms have been found (see Fig.
431
8I) but no B-forms are present. This may be compared with the population of Amphisorus hemprichii
432
from the Red Sea (Zohary et al., 1980). Dasylcadacean and Codiacean algae (Figs. 8A–B) in
433
combination with miliolids, hyaline and agglutinated foraminifera (Figs. 8E–G) associated with
434
Perouvianella peruviana support a local oligotrophic environment within the photic zone in which
435
these foraminifera thrived. A foraminifer that is perhaps similar to Perouvianella peruviana has been
436
found the Cenomanian of Mexico (Jaliscella sigali, Fourcade et al., 1990). The presence of
437
Perouvianella and Jaliscella on the American continent, and the lack of alveolinoideans (the genus
438
Sellialveolina indicated by Jaillard and Arnaud-Vanneau (1993) is not found in our samples) may
439
confirm an endemic Central Peruvian palaeobioprovince that is in contrast with the Tethyan biotic
440
realm for the late Cenomanian–early Turonian time interval.
442
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5.2 Evidence for a disturbed carbon cycle
443
Shallow water carbonates have been shown to record, under favourable conditions, first order
444
patterns of the global marine carbon isotope signature (Weissert et al., 1998; Immenhauser et al.,
445
2005; Föllmi et al., 2006; Huck et al., 2010, 2011, 2012, 2013, 2014; Andrieu et al., 2015). The global
446
seawater δ13CDIC preserved in the rock record are related to changes in the ratio of marine carbonate 18
ACCEPTED MANUSCRIPT 447
carbon and organic carbon burial (Scholle and Arthur, 1980; Weissert, 1989; Immenhauser et al.,
448
2008). With reference to the Peruvian sections documented here, we argue that the mid-Cretaceous low-
450
amplitude sea-level fall was insufficient to expose significant portions of the carbonate ramp studied.
451
Reasons for that might include rapid basement subsidence, a feature that is supported by the
452
stratigraphically thick successions found for the OAE2 interval (Ucchucchacua = 210 m and
453
Cajamarca composite section = 52 m; Fig. 12). Further evidence for rapid basement subsidence comes
454
from the remarkable lack of karst-related features, bleaching, isotopic shifts or other patterns assigned
455
to meteoric diagenesis (Allan and Matthews, 1977; Christ et al., 2012; Huck et al., 2014) beneath
456
discontinuity surfaces recorded in the sections studied. This may imply only weak or no meteoric
457
diagenesis and suggest that the carbonates studied here experienced mainly marine and subsequent
458
burial diagenesis. Judging from the direct comparability of chemostratigraphic sections in Peru with
459
such in many other locations worldwide (Sageman et al., 2006; Jarvis et al., 2011), we tentatively
460
suggest that most of these carbonates stabilized in the presence of marine pore waters. This is turn
461
implies that at least the chemostratigraphic patterns, but perhaps not absolute isotope values, represent
462
palaeoceanographic features as opposed to diagenetic patterns. The question to which degree these
463
patterns reflect global features versus regional water mass properties merits discussion.
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In the Tethyan Ocean, OAE1d is associated to carbon isotope shifts with amplitudes between 0.5
465
and 1.5‰ (e.g., Bornemann et al., 2005; Gambacorta et al., 2015 and references therein; Fig. 9).
466
Despite the limitations of the biostratigraphic control used here, the δ13Ccarb and δ13Corg records at
467
Uchucchacua indicated shifts of 3‰ during the late Albian being more pronounced than those
468
recorded in the Cajamarca composite section (1.5‰ δ13Ccarb and 2.5‰ for δ13Corg; Fig. 12). The
469
carbonates facies, within the limitations of facies characterization, are near-identical in both areas (Fig.
470
11A). This implies that facies-control on chemostratigraphy is weak to absent. It is likely that the
471
regional offset in δ13Ccarb ratio amplitude may have been induced by differences in the water mass
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19
ACCEPTED MANUSCRIPT 472
residence time or differential
473
Immenhauser et al., 2008).
13
C sources (Lloyd, 1964; Holmden et al., 1998; Saltzman, 2003;
For OAE2, the contrast-comparison of δ13Ccarb and δ13Corg records from the Uchucchacua and
475
Lauricocha sections (situated 36 km apart) agrees with the notion of mainly palaeoceanographic
476
patterns and a weak diagenetic overprint (Fig. 10). However, discrepancies exist in the amplitudes of
477
the δ13Ccarb and δ13Corg chemostratigraphic curves when these are compared with the reference section
478
of the Cajamarca region situated approximately 500 km to the north (Fig. 12; Navarro-Ramirez et al.,
479
2016). The differential amplitudes of δ13Ccarb and δ13Corg records are more pronounced in Cajamarca
480
location than in sections at the Uchucchacua and Lauricocha sites. Probably, differences in
481
palaeogeographic settings significantly affected or overprinted global patterns at Uchucchacua and
482
Lauricocha as indicated by mesotrophic and oligotrophic environments in northern and central Peru,
483
respectively (Fig. 11B).
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474
484 5.2.1
Western Platform patterns during the mid-Cretaceous OAEs
TE D
485
In the Western Platform domain, the uppermost Aptian Jacob Level (1st black-shale level of the
487
OAE1b set), which is associated to relative sea-level fall near the Aptian–Albian boundary, caused
488
erosional surfaces, disconformities, basal conglomerates, and ferruginous deposits as reported at the
489
base of the Inca Formation (Jaillard, 1987; Fig. 12). This relative sea-level fall had an amplitude of at
490
least 50m (Maurer et al., 2013) and is probably linked to global cooling and ice sheet dynamics at
491
high-latitudes (Frakes and Francis, 1988; Weissert and Lini, 1991; Frakes et al., 1995; Bodin et al.,
492
2015) responding to global scale and prolonged episodes of organic carbon (Menegatti et al., 1998;
493
Gröcke, 1998; Bralower et al., 1999; Herrle et al., 2004; Westermann et al., 2010; McAnena et al.,
494
2013; Bodin et al., 2015). For this time interval, decreased rates of continental weathering as
495
suggested by low 87Sr/86Sr values and the coeval positive shift of δ18O have been proposed (Bodin et
496
al., 2015).
AC C
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20
ACCEPTED MANUSCRIPT The Kilian and Paquier levels (2nd and 3rd black-shale levels of the OAE1b set) led to transition
498
stages from a siliciclastic-dominated sedimentation to a heterozoan-neritic carbonate factory, followed
499
by incipient phases of platform demise of the Tethyan biotic realm in the Pacific (Robert and Bulot,
500
2004; Iba and Sano, 2007). The earliest Albian corresponds to a change from low to high sea-level,
501
marked by a transition between coldhouse to greenhouse climate mode (Bodin et al., 2015). This
502
climatic pattern is perhaps due to the activity of the South Kerguelen plateau around 112–110 Ma
503
(Coffin et al., 2002; Duncan, 2002; Frey et al., 2003; Bryan and Ferrari, 2013). Early Albian global
504
warming led to a transgressive event that transported water masses from the North Atlantic (Western
505
Tethys) onto the neritic domain of the Western Platform. Moreover, global warming also shifted the
506
Inter-Tropical Convergence Zone (ITCZ) towards the equator, installing more humid climate in the
507
Western Platform domain (Hay and Flögel, 2012; Hasegawa et al., 2012). The enhanced hydrological
508
cycle may have increased continental weathering of the Brazilian continental shield, delivering
509
nutrients and clay-sized material to the ramp, promoting high productivity of benthic organisms
510
(oysters, gastropods, and echinoids). High productivity and an abrupt change in lithology and biota
511
have been associated to the OAE1b set around the globe, indicating the significance of this event
512
during the Cretaceous (Erbacher et al., 1996, 1998, 1999; 2001; Leckie et al., 2002).
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The Leenhardt level (i.e., the last black-shale level of the OAE1b set) is contemporaneous to a
514
major demise of the Cretaceous Tethyan taxa in the Pacific and neritic carbonate production. In the
515
field in Peru, this feature is marked by a regional discontinuity and the disappearance of all lower
516
Albian ammonite species with the exception Glottoceras crassinodosum Sommermeier (Robert, and
517
Bulot, 2004). High atmospheric CO2 output by the South Kerguelen plateau reduced carbonate
518
availability and seawater pH, thereby causing demise episodes of the neritic carbonate production, as
519
well as the loss of calcareous plankton during the broad interval of the OAE1b set (Leckie et al.,
520
2002).
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521
In the middle Albian, the increased flux of isotopically light CO2 at times of the activation of the
522
western South America volcanic arc, shifted the δ13Ccarb towards more negative values as perhaps 21
ACCEPTED MANUSCRIPT 523
documented in the Cajamarca composite section (Navarro-Ramirez et al., 2015, 2016; Fig. 12). This
524
feature is evidenced by reduced carbonate production in the Western Platform domain. The Albian global warming continued until the early Turonian, reaching a mid-Cretaceous climate
526
maximum (Pagani et al., 2014). This led to a rapid diversification and a pronounced increase in the
527
degree of calcification of the planktic foraminifera, increased upper water column stratification,
528
initiation of widespread chalk deposition and oligotrophic conditions in (hemi-) pelagic settings in the
529
Atlantic and Tethys (Leckie, 1989; Premoli Silva and Sliter, 1999; Leckie et al., 2002). In contrast,
530
global warming resulted in an enhanced hydrodynamic circle with erosion delivering nutrients and
531
promoting diversification of benthic communities dominated by oyster bioherms and large endemic
532
benthonic foraminifera in the Western Platform (Fig. 11A). On the other hand, in the Tethyan realm
533
and elsewhere, OAE1d, MCEI, and OAE2 have been associated to high turnover rates affecting the
534
planktic foraminifera and radiolarians, anoxia, organic-rich deposition, and demise of the carbonate
535
platform (e.g., Erbacher et al., 1996, 2001; Gale et al., 2000; Wilson and Norris, 2001; Leckie et al.,
536
2002; Bornemann et al., 2005; Gertsch et al., 2010; Schröder-Adams et al., 2012; Scott et al., 2013;
537
Owens et al., 2013; Giraud et al. 2013; Melinte-Dobrinescu et al., 2015; Gambacorta et al., 2015;
538
Andrieu et al., 2015). These patterns are in clear contrast to phases of enhanced carbonate deposition
539
documented for the sections in Peru discussed here. Persistent deposition of benthic communities in
540
these sections supports the notion of well-oxygenated water masses throughout the middle Albian–
541
early Turonian interval.
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525
542
In the Tethyan realm, several groups of foraminifera such as the alveolinids, rhapydionids,
543
pseudorahpydioninids, and nezzazatids characterize the Cenomanian (see Schroeder and Neumann,
544
1985; Calonge et al., 2002; Vicedo et al., 2011; Consorti et al., 2015 and the work cited therein). All of
545
these foraminifera disappear at the Cenomanian–Turonian boundary. The finding of Perouvianella
546
peruviana in Jumasha 4 (Fig. 10) in the central Peru sections is here considered evidence for the
547
continuation of a stable, albeit endemic, marine ecosystem during and after OAE2 contrasting the
548
Tethyan biotic patterns (Parente et al., 2008; Fig. 12). In combination with similar findings in neritic 22
ACCEPTED MANUSCRIPT 549
carbonates in the Pacific - lacking evidence of anoxic conditions during OAE2 (Elrick et al., 2009;
550
Takashima et al., 2011) - this suggests that anoxic conditions did not establish in major portions of the
551
Pacific Ocean during OAE2 (Takashima et al., 2011). In the Cajamarca composite section, a prominent negative shift is observed at the onset of OAE2
553
(Fig. 12). This negative δ13C excursion has also been identified in other sections in the Pacific realm
554
(California, USA; Hokkaido, Japan; Takashima et al., 2011). According to Takashima et al. (2011),
555
this negative δ13C excursion is linked to an increase in pCO2 induced by the formation of the
556
Caribbean large igneous province (LIP; ~95.1–92.2 Ma, Snow et al., 2005). Arguments for the claim
557
that the Caribbean massive volcanism triggered OAE2 come from a negative 34Ssulphate shift ca. 500 kyr
558
before OAE2 (Adams et al., 2010) as well as from a negative
559
2014) and increased pCO2 levels (e.g., Barclay et al., 2010).
SC
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552
Os/188Os shift (e.g., Du Vivier et al.,
M AN U
187
This significant increase in pCO2 drove the climate to a hothouse mode (Kidder and Worsley
561
2010), perhaps leading to an acidified ocean (Du Vivier et al., 2015) and a perturbed hydrological
562
cycle, that enhanced weathering rates at the beginning of the onset of OAE2 (Frijia and Parente, 2008;
563
Blätter et al., 2011; van Helmond et al., 2013; Pogge von Strandmann et al., 2013). The beginning of
564
the onset of OAE2 in the Western Platform is coeval to an increased influx of argillaceous material
565
and reduced carbonate production in northern Peru (Fig. 11B). However, continuous deposition of
566
floatstones and grainstones containing Perouvianella peruviana is observed in the central Peru
567
sections (Fig. 10) possibly due to the restriction of continental nutrient influx (Fig. 11B). Consumption
568
of significant amount of pCO2 by the enhanced global organic carbon burial and by the continental
569
weathering after the onset of the Caribbean LIP (Sinninghe Damsté et al., 2010; Jarvis et al., 2011; van
570
Helmond et al., 2013; Pogge von Strandmann et al., 2013) conditioned the oceans to gradually be
571
more oxic (Pogge von Strandmann et al., 2013) and might have caused the reduction of benthic
572
communities in the Western Platform of Peru after OAE2 (Figs. 10 and 12).
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Conclusions The mid-Cretaceous sections of the Western Platform of Peru document a predominance of
576
oyster-rich, mixed siliciclastic-carbonate facies in the middle Albian–late Cenomanian. Regional
577
evidence suggests that this rather uniform facies belt expanded across wide portions of the spatially
578
very extensive carbonate ramp in northern and central Peru. In the late Cenomanian and the early
579
Turonian, the deposition of an endemic large benthic foraminifera facies (Perouvianella peruviana) is
580
a key facies element that continues throughout the OAE2 interval without discernible adaptation to
581
environmental change. During this time, central Peru was isolated from the Pacific water masses and
582
from deltaic influx from the emerged shield of Brazil in the east. This specific palaeoceanographic
583
setting might have promoted the mass occurrence of Perouvianella peruviana. Probably due to the
584
complex setting of the Western Platform in the mid-Cretaceous, clear evidence for a carbonate crisis
585
associated with OAE 1d and OAE 2 in the sections of the central Peru is lacking.
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Global warming during the Albian–early Turonian resulted in a more humid climate, high trophic
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levels, and a diversification of benthic communities in Western Platform carbonates, influenced by
588
palaeogeographic modifications. Even though the benthic communities thrived during extremely warm
589
conditions in the late middle Albian and towards the early Turonian, they were exposed to high
590
nutrient levels caused by a perturbed hydrological cycle at the beginning of the onset OAE2. The
591
intensified chemical weathering on the Brazilian shield during the mid-Cretaceous may have
592
contributed to trigger the OAE in the North Atlantic nutrient trap. The work documented here is a
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clear case example for the superposition of regional and global patterns in palaeoceanography and
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complements scarce existing data from South America for this critical interval in Earth’s history.
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Acknowledgements
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ACCEPTED MANUSCRIPT This project was supported by the Deutsche Forschungsgemeinschaft (DFG, project n° BO-365/2-1)
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and by the Deutscher Akademischer Austauschdienst (DAAD) through a scholarship to J.P.N. (PKZ:
600
91540654). We thank the Geological survey of Peru (INGEMMET) for important logistical support.
601
Esmeralda Caus (Barcelona) is gratefully acknowledged for a critical reading of the text in its early
602
version. Analytical work was performed in the isotope laboratories at Bochum and Erlangen-
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Nuremberg. We greatly acknowledge the editorial advice of E. Koutsoukos and the comments of two
604
anonymous journal reviewers.
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Figure Captions
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Figure 1: Palaeogeographic reconstruction for the mid-Cretaceous (modified after Blakey, 2011)
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indicating the position of what today is the Western Platform of South America (yellow arrow). Figure 2: Geological map of the Oyon region (modified after INGEMMET, 1980) showing the location of the Uchucchacua and Lauricocha sections documented in this study. Figure 3: Regional correlation of the sequence stratigraphic interpretation between the Cajamarca
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composite section (northern Peru; Navarro-Ramirez et al., 2016) and the Uchucchacua section (central
1003
Peru; this study). The Uchucchacua section is juxtaposed with our sections (Jumasha 1 and 4) and by
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the Jaillard (1986)´s work. Main biostratigraphic data are indicated (Benavides-Caceres, 1956; Wilson
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1963; von Hillebrandt, 1970; Romani, 1982; Jaillard 1986, 1987; Jaillard and Arnaud-Vanneau, 1993;
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Robert 2002; Robert and Bulot, 2004, 2005; Dhondt and Jaillard, 2005; Jaillard et al., 2005; Robert et
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al., 2009).
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Figure 4: Legend for figs. 3, 9 and 10, denoting colour codes for different facies types.
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Figure 5: Landscape images showing A): The Jumasha Formation of the upper Albian–lower
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Turonian at Uchucchacua, Oyon region, Central Peru, built by massive, thickly-bedded, light-grey
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limestones. B) Lower part of the Jumasha Formation at Uchucchacua, including the upper Albian and
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the lower Cenomanian rocks. C) The upper part of the Jumasha Formation at Uchucchacua, including 42
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the middle Cenomanian to lower Turonian rocks. D) The upper part of the Jumasha Formation at
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Lauricocha, including the late Cenomanian to lower Turonian rocks. Figure 6: Field images taken at the Uchucchacua section. A) Thin lamination of facies 1a,
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represented by grey and argillaceous limestones intercalations. B) Nodular limestones of the facies 2e.
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C) Thickly-bedded limestones of facies 2g. D) Thickly-bedded limestones of medium-scale sequence
1018
10 (facies 2b). C) Floatstones of facies 3d. F) Bioturbation at the level of regressive deposits in the top
1019
of Jumasha 4.
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Figure 7A): Thin section images documenting grainstones of facies 2d, represented by peloids,
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miliolids, and isolated specimen of Perouvianella peruviana. B) Dasycladales-bearing grainstones of
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facies 2f. C) Miliolids and Perouvianella peruviana of facies 2g. D) Grain-rudstones dominated by
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Peruvianella peruviana in facies 2h. E) Diverse fauna floatstones of facies 3b, characterized by a
1024
chaotic fabric. F) Packstone rich in echinoids of facies 3c.
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Figure 8: Upper Cenomanian‒lower Turonian microfossil from shallow-water deposits of the
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Jumasha Formation in central Peru. Scale bar in images C and D is 0.5 mm, in all other images the
1027
scale bar is 1 mm. A) and B) show dasycladales, closely related to the genus Trinocladus. A.
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Transversal section. B. Longitudinal section (central part) and transversal section (upper right part). C)
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and D) Rotorbinella mesogeensis; C. axial section, note the presence of central plug. D. equatorial
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section. E) Agglutinated foraminifera in equatorial section. F) Porcelaneous foraminifera in equatorial
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section. G) Agglutinated foraminifera with spiral outline closely related to the genus Merlingina
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cretacea Hamaoui and Saint Marc, 1965. H) Perouvianella peroviana in perfect equatorial section. I)
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Perouvianella peroviana proloculi of two specimens, difference in diameter suggest the existence of
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multiple cycles of reproduction within the population.
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Figure 9: Sequence stratigraphic interpretation and carbon-isotope stratigraphy of the
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Uchucchacua section for the upper Albian (Jumasha 1). Correlation with the Tethyan composite curve
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of Herrle et al. (2015) is shown. Biostratigraphic data for the late Albian have been obtained from
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Jaillard (1986, 1987 and references therein). 43
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Uchucchacua and Lauricocha sections for the late Cenomanian (Jumasha 4). Data points shaded grey
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correspond to OAE2 based on bio- and chemostratigraphic data. Biostratigraphic data of the
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Uchucchacua section has been taken from Jaillard (1986) and Jaillard and Arnaud-Vanneau (1993).
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The chemostratigraphic markers of OAE2 (A‒B) include Navarro-Ramirez et al. (2016), Eastbourne
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(Jarvis et al., 2011), and the Portland # 1 core (Sageman et al., 2006).
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Figure 11A): Palaeogeographic interpretation of the western South America platform during the
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late Albian, characterized by oyster-rich mixed siliciclastic-carbonate deposition. Note regional oyster
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facies deposition, diachronic progradation of the NE deltaic system fed from the Brazilian continental
1048
basement, as well as the Huarmey Trough isolating the western Platform from the Pacific Ocean. In
1049
contrast to the central Peru, northern Peru was localized more proximal to the southern margin of the
1050
proto-Atlantic, resulting in a potentially stratified water column setting. B) During the late
1051
Cenomanian, the Western Platform was characterized by two distinctive palaeogeographic settings: (i)
1052
northern Peru characterized by continental deltaic influx from the emerged shield of Brazil in the east;
1053
(ii) central Peru representing an oligotrophic environment, leading to the mass occurrence of
1054
Perouvianella peruviana and associated miliolids. The four main depositional environments, each
1055
with a number of standard facies types (1a through 4a), are established for the inner to outer ramp.
1056
Interpretations are based on Benavides-Caceres (1956), Jaillard (1986, 1987), Atherton and Webb
1057
(1989), Soler and Bonhomme (1990), Jaillard et al. (2000), Callot et al. (2008), Robert et al. (2009)
1058
and Winter et al. (2010).
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Figure 12: Regional correlation of the carbon-isotope stratigraphy between the Cajamarca
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composite section (northern Peru; Navarro-Ramirez et al., 2016) and by the Uchucchacua section
1061
(central Peru), exhibiting the Jumasha 1 and 4 carbon-isotope data. Main OAEs and biostratigraphic
1062
data are indicated (Benavides-Caceres, 1956; Jaillard, 1986, 1987; Jaillard and Arnaud-Vanneau,
1063
1993; Dhondt and Jaillard, 2005; Robert et al., 2009). The chemostratigraphic terminology for OAE2
1064
interval (onset, trough, plateau, and recovery, c.f., Paul et al., 1999; Tsikos et al., 2004), and datum 44
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levels A, B, C and D as is used in Tethyan and North Atlantic sections (Meyers et al., 2012; Du Vivier
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et al., 2014, 2015). The red data points in the Cajamarca composite section form part of this study. Table 1: Overview of facies classification and interpretation. Numbers indicate the relative
1068
abundance of non-skeletal and skeletal components: 0 = absent, 1 = present, 2 = frequent, 3 =
1069
abundant, 4 = dominant.
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New Albian–Cenomanian and Cenomanian–Turonian carbon isotope curves based on epeiric-neritic successions from Central Peru Evidence for the impact of transient carbon cycle perturbations (OAE1d and OAE2) is provided
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Clear evidence for a carbonate neritic crisis associated with OAE 1d and OAE 2 is lacking
S0195-6671(16)30203-8
DOI:
10.1016/j.cretres.2017.03.009
Reference:
YCRES 3555
To appear in:
Cretaceous Research
Received Date: 9 September 2016 Accepted Date: 9 March 2017
Please cite this article as: Navarro-Ramirez, J.P., Bodin, S., Consorti, L., Immenhauser, A., Response of western South American epeiric-neritic ecosystem to middle Cretaceous Oceanic Anoxic Events, Cretaceous Research (2017), doi: 10.1016/j.cretres.2017.03.009. This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.
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Response of western South American epeiric-neritic ecosystem to middle Cretaceous Oceanic Anoxic Events
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Navarro-Ramirez J.P.1, *, Bodin S.2, Consorti L.3, Immenhauser A.1
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Germany
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Aarhus University, Department of Geoscience, Høegh-Guldbergs Gade 2, 8000 Aarhus C, Denmark
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Departament de Geologia (Paleontologia), Universitat Autònoma de Barcelona, 08193 Bellaterra,
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Spain
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Ruhr-Universität Bochum, Institut für Geologie, Mineralogie und Geophysik, D-44870 Bochum,
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* Corresponding author. Tel.: +49-234-32-2325.
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E-mail address: [email protected]
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Revised version for the Journal Cretaceous Research
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Keywords: Peru, mid-Cretaceous, Pacific, OAE1d, OAE2, chemostratigraphy, carbonates
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Abstract Little is known about the impact of the mid-Cretaceous Oceanic Anoxic Events (OAEs) on the
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neritic carbonate systems in South America. In order to fill this knowledge gap, the present paper
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reports on the record of environmental changes in the Albian–Turonian neritic carbonates from the
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western South American domain in Peru. Owing to the very expanded and well-exposed sections in
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the Oyon region of central Peru, the OAE 1d and 2 intervals were sampled at high temporal resolution
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for both bulk micrite and bulk organic matter carbon isotopes, allowing us to compare the fingerprint
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of these two events between the northern and central Peruvian regions. This suggests the installation of
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two marked depositional modes: 1) the Albian–Turonian formation of a regional facies belt constituted
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by oyster-rich mixed siliciclastic-carbonate deposition along the western South America platform; 2) a
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restricted oligotrophic environment, characterized by the mass occurrence of Perouvianella peruviana
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and associated miliolids in central Peru during the late Cenomanian–Turonian. These observation
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advocate for the following scenario: Global warming during the late Albian–early Turonian resulted in
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humid climate on the western platform. This in turn caused enhanced chemical weathering rates on the
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Brazilian Shield, resulting in high runoff of nutrients onto the western platform. Nutrient runoff
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promoted the diversification of benthic oyster communities. Due to the uplift of the Marañon Massif
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and the installation of the Huarmey Trough, central Peru was isolated from the Pacific and from
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eastern deltaic influx of the Brazilian continental basement, allowing the local development of
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oligotrophic conditions during OAE 2. Furthermore, an increased influx of argillaceous sediment and
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reduced carbonate production is recorded in northern Peru at the onset of OAE 2, marked by a
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prominent negative shift in δ13C. This negative carbon-isotope excursion has also been identified in
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other sections in the Pacific domain and can be linked to an increase in isotopically light pCO2 induced
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by the formation of the Caribbean large igneous province.
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1. Introduction The mid-Cretaceous (i.e., the Albian–Turonian interval, ca. 110–90 Ma; Ogg and Hinnov, 2012) is
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considered as a time governed by intensification of greenhouse conditions due to elevated emission of
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pCO2 to the atmosphere following massive pulses of ocean crust production and the formation of large
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igneous provinces (LIPs; Arthur et al., 1985; Barron and Washington 1985; Larson, 1991; Leckie et
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al., 2002; Snow et al., 2005; Bodin et al., 2015). The intensified greenhouse conditions enhanced the
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hydrological cycle, which in turn accelerated continental weathering rates (Berner et al., 1983;
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Weissert, 1990; Weissert et al. 1998; Hay, 1998; Menegatti et al., 1998; Wortmann et al., 2004; van
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Helmond et al., 2013; Bodin et al., 2015) and led to increased fluxes of continental nutrients to the
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oceans. This chain of events stimulated biological productivity, leading to widespread organic-carbon
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deposition and enhanced marine anoxia and euxinia (Manabe and Bryan, 1985; Föllmi et al., 1994,
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2006; Turgeon and Brumsack, 2006; Owens et al., 2013; Pogge von Strandmann et al., 2013; Lechler
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et al., 2015). The processes described above probably occurred on some level throughout the
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Cretaceous, but became particularly enhanced during several geologically short intervals (<1 Myr),
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known as oceanic anoxic events (OAEs). These latter are best defined based on carbon isotope (δ13C)
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excursions that reflect a fundamental perturbation in the global carbon cycle (Erbacher et al., 1996;
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Leckie et al., 2002). The best-studied OAEs in Europe, North America, and from Atlantic ODP sites
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include the early Aptian OAE1a, the Albian OAE1b, OAE1c and OAE1d, and the Cenomanian–
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Turonian OAE2. Additional events, such as the Coniacian OAE3 and the mid-Cenomanian event I
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(MCEI) represent related phenomena that lack widespread organic carbon-rich facies and/or a
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significant δ13C excursion (e.g., Jarvis et al., 2006; Jenkyns, 2010).
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OAE1a and OAE2 are conventionally interpreted as the result of massive outgassing of the
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Ontong Java, Manihiki and Hikurangi LIP (Tarduno et al., 1991; Larson and Erba, 1999; Méhay et al.,
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2009; Tejada et al., 2009) and by the Caribbean LIP (Snow et al., 2005; Du Vivier et al., 2014). At the
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onset of OAE1a and OAE2, a negative shift in lithium and calcium isotope ratios is observed, most
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likely associated to weathering of continental crust (Blättler et al., 2011; Pogge von Strandmann et al., 3
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(southeastern France), and Eastbourne (England) sections, suggested that the Caribbean LIP triggered
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a change in the predominant type of biomineralization (i.e., ocean acidification) at the onset of OAE2
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(Du Vivier et al., 2015). As recorded in the geological record, transient periods of high atmospheric
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CO2 led to several mass extinctions, most likely because excess CO2 reduced the ability of carbonate-
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secreting organisms to secrete their carbonate shells (e.g., Glikson, 2010). In a more qualitative way,
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sedimentological studies suggest enhanced weathering at the onset of OAE 1a and OAE2, as reflected
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by the relative abundance of kaolinite (Stein et al., 2012; Gertsch et al., 2010).
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Evidences of a sudden influx of siliciclastic material coeval to OAEs have been reported in the
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North Atlantic and the Tethys domains (Weissert, 1990; Wortmann et al., 2004) and are indicative of a
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change to more humid and warmer conditions (Weissert, 1990; Weissert et al., 1998; Wortmann et al.,
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2004). This concept is exemplified in the evolution of the northern Tethyan platform during the Early
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Cretaceous, where carbon isotope (δ13C) excursions occurred at times of elevated nutrient levels due to
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intensified continental weathering rates (Erba, 1994; Weissert et al., 1998; Immenhauser et al., 2005;
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Bodin et al., 2006; Föllmi et al., 2006; Huck et al., 2010, 2011, 2012, 2013, 2014). This caused a
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change from oligo- to meso- or eutrophic conditions, ecological reorganization, phases of carbonate
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platform development and drowning episodes (Weissert et al., 1998; Immenhauser et al., 2005; Bodin
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et al., 2006; Föllmi et al., 2006; Huck et al., 2011). However, these hypotheses have yet to be applied
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to sections in South America, which record the response of shallow-marine carbonate depositional
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environments of a southern hemisphere continent adjacent to the Pacific Ocean.
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The western South America platform (Fig. 1) was characterised by a large oceanic basin to the
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west and a very large uplifted continental basement area to the east. The enormous dimensions of this
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region make this a challenging natural laboratory to understand southern hemisphere Earth System
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behaviour under extreme climates. The aim of this paper is to document and discuss the Albian to
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Turonian sedimentological and palaeoecological evolution of the epeiric-neritic ecosystem of the
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western platform in central Peru (Oyon region; Jaillard, 1986; Fig. 2). The assessment of potential 4
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published northern Peru reference composite section (Cajamarca region, Fig. 3, Navarro-Ramirez et
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al., 2015, 2016). Specific attention is paid to environmental stressors associated to OAE1d and OAE2
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and the manner in which they affect the main carbonate producers such as oysters, gastropods,
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echinoids, and endemic large benthic foraminifera (Perouvianella peruviana, Steimann) that have so far only been reported from these sections (Jaillard and Arnaud-Vanneau, 1993).
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2. Regional tectonic and stratigraphic setting
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The South America margin was largely influenced by late stages of Gondwana breakup that
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culminated in the separation of Africa and South America and the opening of the Equatorial Atlantic
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Ocean during the Aptian–Albian transition (Moulin et al., 2010; Fig. 1). This allowed for the
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connection of North and South Atlantic Ocean water masses towards the Turonian (Eagles, 2006).
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This tectonic event caused the activation of the western South America volcanic arc in the Early
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Cretaceous (Huarmey-Trough, Atherton and Webb, 1989; Soler and Bonhomme, 1990; Jaillard et al.,
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1999, 2000; Winter et al., 2010) and subduction took place along the western portion of South
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America where sedimentation was controlled by NNW–SSE-trending structures (e.g., Benavides-
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Caceres, 1956; Jaillard, 1986, 1987; Jaillard et al., 1990; Callot et al., 2008; Robert et al., 2009).
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Mid-Cretaceous sedimentation in what today is Peru began with the deposition of siliciclastic
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rocks now forming the Inca Formation, which is built by iron-rich, sandy beds, assigned to an early
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Albian age based on the ammonite Neodeshayesites nicholsoni (Robert and Bulot, 2004; Robert et al.,
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2009). This siliciclastic unit in turn is overlain by Albian carbonates represented by the Chulec, and
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Pariatambo formations, recording a second-order Albian transgression (Benavides-Caceres, 1956;
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Jaillard, 1986; Robert, 2002; Robert and Bulot, 2004, 2005).
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The Chulec Formation unconformably overlies the Inca Formation with a discontinuity surface
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separating the two units (Jaillard, 1987). The Chulec Formation is characterized by marl-limestone 5
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ammonites have been reported and were assigned to the Knemiceras raimondii Zone (Robert and
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Bulot, 2004; Robert et al., 2009), indicating an early Albian age. The Pariatambo Formation is
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characterized by fossiliferous, black, bituminous, fetid marly limestones and includes finely laminated
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and siliceous intervals. Ammonites and planktonic foraminifera are common. The presence of
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Oxytropidoceras carbonacrium and Prolyelliceras ulrichi Zones indicate an early middle Albian age
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(Robert et al., 2009). The Pariatambo Formation grades upsection into nodular bioclastic grey and
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marly limestones of the Yumagual Formation in the northern Peru and by the Jumasha Formation in
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central Peru (upper middle Albian– lower Turonian; Fig. 3).
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At Uchucchacua (Oyon region; central Peru; Jailard, 1986) the Jumasha Formation is about 1320
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m thick and made of very massive, thickly-bedded, light yellowish brown to brownish-grey limestones
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(Jaillard, 1986; Fig. 3). The base of the Jumasha Formation was attributed to the upper middle Albian
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due to ammonite associations including Lyelliceras ulrichi Knechtel and Oxytropidoceras douglasi
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(Benavides-Caceres, 1956; Wilson, 1963; von Hillebrandt, 1970; Jaillad, 1986; Fig. 3). At section
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metre 115 (Fig. 3) i.e., the base of the Jumasha Formation near Uchucchacua, Romani (1982) reported
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Exogyra costagyra cf. olisoponensis and Merlingina cretacea of Cenomanian age. Further south, in
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the upper part of the Jumasha Formation microfauna data and particularly Favusella washitensis, H.
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delrioensis, Globigerinelloides bentonensis and Heterohelix seewashitensis, indicate a middle to late
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Cenomanian age (von Hillebrandt, 1970; Jaillard, 1986; Fig. 3). Rotorbinella mesogeensis (Tronchetti)
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is equally present and supports the Cenomanian age of this unit. The presence of Coilopoceras sp. at
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the top of the Jumasha Formation suggests a Turonian age (Romani, 1982; Jaillard, 1987; Fig. 3).
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It must be noted that, due to scares ammonites and planktonic foraminifera findings in the
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Peruvian sections, the biostratigraphic scheme established by the aforementioned authors remains
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incomplete in places. There are uncertainties about the exact position of the stage and substage
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boundaries, although the general scheme has been confirmed by carbon isotope chemostratigraphic
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correlation as well as Sr-isotope chronostratigraphy (Navarro-Ramirez et al., 2015, 2016). 6
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3. Methods and materials
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3.1 Field work and thin-section microscopy Two well-exposed sections (Uchucchacua and Lauricocha localities; Fig. 2) were chosen for this
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study to provide maximum coverage of the Albian–Cenomanian and Cenomanian–Turonian
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transitions (Fig. 3). In total, ca. 620 m of section have been logged and studied for their
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sedimentological, stratigraphic, and chemostratigraphic archive (Figs. 3 and 4). Outcrop observations,
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rock samples, and thin-sections provide the fundament for petrographic and facies interpretations
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presented here and allow us to place the chemostratigraphic results in a facies context. In analogy with
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the nomenclature and approach presented in Navarro-Ramirez et al. (2015, 2016), non-skeletal
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components and skeletal components were analysed semi-quantitatively with numbers indicating their
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relative abundance: 0 = absent, 1 = present, 2 = frequent, 3 = abundant, 4 = dominant; Table 1). Facies
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pattern and discontinuities acted as pining points for sequence stratigraphic interpretations (see also
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Christ et al., 2015).
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3.2. Carbon isotope stratigraphy
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3.2.1.
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Bulk micrite data (δ13CCarb)
Carbon-isotope analysis of 181 micrite samples was performed using a Thermo Finnigan MAT
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delta-S mass spectrometer at the isotope laboratory of the Ruhr-University Bochum (see Appendix A),
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Germany. In order to reveal the variability of carbon-isotope composition within a single rock sample,
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several sub-samples were drilled from 20% of all the hand specimens collected. Carbonate bulk rock
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specimens were sampled by use of tungsten drill bits. While drilling the carbonate powder, areas rich
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in spary cement, large bioclasts and diagenetic calcite vein material were avoided. Refer to
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Immenhauser et al. (2005) for more details of the analytical procedure. Repeated analyses of certified
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carbonate standards (NBS 19, IAEA CO-1 and CO-8) and internal standards show an external 7
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reproducibility of ≤0.06‰ for δ13Ccarb. The δ13Ccarb values are expressed on a per mil (‰) basis relative
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to the Vienna-Pee Dee Belemnite standard (V-PDB).
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Bulk organic matter data (δ13Corg)
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Carbon-isotope analyses were performed on 201 bulk organic matter samples (see Appendix B).
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Two grams of sample powder were placed in 50mL centrifuge tubes and 6N HCl was added, this
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procedure was repeated until no carbonate reaction was visible any more. The samples are then
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centrifuged and the supernatant removed, the residues were rinsed with distilled water via centrifuge
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and dried at 40 °C. The δ13Corg measurements were performed with an elemental analyser (CE 1110)
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connected online to ThermoFinnigan Delta V Plus masspectrometer. All carbon isotope values are
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reported in the conventional δ‐notation in permil relative to V-PDB (Vienna‐PDB). Accuracy and
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reproducibility of the analyses was checked by replicate analyses of international or laboratory
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standards. Organic carbon analyses were conducted at the stable isotope laboratory of the Friedrich-
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Alexander University of Erlangen-Nuremberg, Germany.
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4. Data description and interpretation
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The Uchucchacua section is here described as the representative case example for the Albian to
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Turonian sedimentary record in the study area (Jaillard, 1986; Jaillard and Arnaud-Vanneau, 1993;
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Figs. 3 and 4). This choice is motivated by the fact that the outcrop conditions are excellent and
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stratigraphically near-complete (Figs. 5A–C). In addition, the Lauricocha section, which is of
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comparably high outcrop quality, serves as a complementary section allowing to separate local from
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regional facies changes during the late Cenomanian to early Turonian transition (Fig. 5D).
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4.1 Facies associations and depositional environments By integrating previous work (e.g., Jaillard, 1986, 1987; Jaillard and Arnaud-Vanneau, 1993;
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Dhondt and Jaillard, 2005; Jaillard et al., 2005) with our field data, four main depositional
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environments, each with a number of standard facies types are established for the intertidal to outer
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ramp domain (Table 1). The documentation of these fundamental sedimentological data is important to
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match the type of carbonate producers versus OAEs.
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Please note that the first depositional environment is represented by facies 1a, 1b and 1c, where
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facies 1a is the most proximal deposits composed of sandy marls with scarce fauna (Fig. 6A).
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Whereas, facies 1b and 1c, studied by Jaillard (1986) for the lower Cenomanian in the Uchucchacua
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section, are generally represented by beds of dolostones with sedimentary structures like tepees,
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birdeyes and mudcracks (facies 1b) and by algal laminites and lenticular structures (facies 1c),
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corresponding all to a intertidal zone setting (please refer to Jaillard (1986)´s work for more details).
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Shallow subtidal inner ramp setting
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The Albian facies types recognized at Uchucchacua display important similarities with those
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previously described from the Cajamarca region (Northern Peru, Navarro-Ramirez et al., 2015, 2016)
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and are only briefly summarized here (Fig. 4 and Table 1). At Uchucchacua, facies 2a is composed of
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grey argillaceous mudstones with scarce fauna, whereas in more distal sub-environments, facies 2b is
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characterized by packstones and argillaceous wackestones, exhibiting a nodular fabric due to an
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increased argillaceous content. Facies 2c is characterized by grainstone and occasionally floatstone
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and is typically rich in mixed and fragmented shell debris, mainly derived from oysters, gastropods,
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and echinoids (Table 1).
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The Cenomanian–Turonian facies types recorded at Uchucchacua and Lauricocha show also strong similarities as represented by the Facies 2d–h and 3b and 3c (Table 1).
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ACCEPTED MANUSCRIPT Facies 2d — Peloid-bearing grainstones (Table 1): This facies is composed of light-grey peloidal
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grainstones, showing cm to dm-thick lamination. The fauna yields some disperse miliolids, ostracods
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and sparse Perouvianella peruviana (Fig. 7A) as well as Charophyte gyrogonids. In view of the
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predominance of peloids, disperse miliolids and the low diversity of microfauna, facies 2d is
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interpreted to reflect a restricted lagoon environment (Gräfe, 2005) with limited water circulation
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(Masse et al., 2003) and low trophic levels (Hüneke et al., 2001).
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Facies 2e — Discorbids and ostracods wakestone/mudstone (Table 1): This facies is primarily
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composed by small debris (layers) of discorbid and ostracod valvae accumulations whilst rare
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miliolids are also present (Fig. 6B). This basic facies type has been described first by Sartoni and
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Crescenti (1962) from the Mesozoic of Italy and was interpreted as an environment characterized by
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restricted water circulation, medium to high trophic levels and low salinity.
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Facies 2f — Dasycladales-bearing pack-grainstones (Table 1): This facies is built by dm- to m-
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thick bedded, bluish-grey packstones and occasionally by grainstones. The rocks are typically
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characterized by large rounded bioclasts composed of dasycladales remains (Figs. 7B and 8A–B). The
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main benthic foraminifera are Perouvianella peruviana and miliolids whose tests are often abraded.
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Occasionally peloids are present. Distribution of dasycladales may occur in shallow shoals, lagoonal,
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semi-restricted very shallow environment and backreef environments (Flügel, 2004).
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Facies 2g — Miliolid-bearing pack-grainstones (Table 1): This facies is made of bluish-grey
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packstones to grainstones, displaying dm-thick lamination in units of two to five metres thick (Fig.
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6C). Bioturbation features are abundant. The fauna is dominated by miliolids, being associated to
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Perouvianella peruviana (Fig. 7C) and numerous dasycladal algae. Peloids also occur in certain levels.
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Associations of coated grains, skeletal elements such as miliolids, Perouvianella peruviana, and
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dasycladales indicate deposition in the nearshore shallow zone (Flügel, 2004; Lézin et al., 2012).
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Modern miliolids live mainly in warm, clear water lacking significant fresh-water influx and in some
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places they are associated with reefal settings (Fang, 2003).
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is characterized by grainstones, rudstones, and floatstones containing numerous Perouvianella
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peruviana (Fig. 7D). Facies 2h builds thickening upwards sequences, each one being 20 to 80 m in
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thickness. Typical biota include miliolids, Rotorbinella mesogeensis (Figs. 8C–D), agglutinated (Figs.
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8E and G), porcelaneous foraminifera (Fig. 8F) and Perouvianella peruviana (Figs. 8H–I). According
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to previous workers, this facies may indicate deposition in the external part of the inner platform (<20
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m; Stein et al., 2012), located in the platform margin (Flügel, 2004) with specifically clear and well-
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illuminated waters and low trophic levels (Lézin et al., 2012).
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Open marine middle ramp setting
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In the Central Peru upper Albian interval, facies types 3a, 3b, 3c, and 4a are near-identical to those
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identified at the Piedra Parada location (Northern Peru) as previously describe in detail in Navarro-
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Ramirez et al. (2016; Table 1). For the sake of brevity, we only report the most important
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characteristics here. Facies 3a is mainly characterized by oyster bioherms, interbedded by rare
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grainstone units yielding various skeletal elements and intraclasts. Facies 3b (diverse fauna low
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energy) in the lower Jumasha Formation is characterized by 5 to 10 cm-thick layers with variably
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sized fragments of bryozoans, echinoids, and planktonic foraminifera. Whereas, this facies in the
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upper Jumasha Formation, occurs as dm to m-thick bedded grey pack to grainstones and occasionally
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by floatstones (Fig. 6D). This here includes mixed and fragmented shell debris mainly consisting of
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Perouvianella peruviana, miliolids, gastropods, spicules and rare charophyte gyrogonites (Fig. 7E).
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Dominance of intraclasts appear also in this facies, which are usually interpreted to be formed by
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storm wave erosion occurring in shallow-marine environments (Flügel, 2004). Following previous
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workers, this type of deposits indicate deposition in high-energy, shallow subtidal shoals possibly also
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in the form of storm event beds (Elrick et al., 2009), just beneath the fair-weather wave base (Amini et
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al., 2004; Stein et al., 2012; Lézin et al., 2012).
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organized in higher order units of two to five metres in thickness and is rich in echinoid debris (Figs.
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6E and 7F). These features suggest a middle ramp environment and sediment deposition beneath the
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fair-weather wave base (Flügel, 2004; Stein et al., 2012). This facies 3c appears also in the upper
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Cenomanian sections at Uchucchacua and Lauricocha.
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Facies 4a (planktonic foraminifera-bearing mudstones) is characterized by dark grey mudstones
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with mm to cm-thick lamination and a generally undisturbed horizontal bedding. The fauna includes
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planktonic foraminifera and small echinoid fragments. These evidence an open marine setting below
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the reach of the effective storm wave base (Lézin et al., 2012; Stein et al., 2012).
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4.2 Sequence stratigraphic interpretation
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A basic interpretation of the sequence stratigraphic scheme of Central Peru was presented in
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Jaillard (1986, 1987), Jaillard and Arnaud-Vanneau (1993) and Jaillard et al. (2005) and for the
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Northern Peru by Navarro-Ramirez et al. (2015, 2016), respectively. Following these approaches, we
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tie the sequence stratigraphic for both regional settings as indicated in figure 3, where the Jumasha
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Formation in Uchucchacua is divided in 5 large-scale sequences (Jumasha 1–5; Fig. 3). Here, the
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upper Albian to lower Cenomanian is located in the first Jumasha sequence and the upper Cenomanian
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to lower Turonian in the Jumasha 4 (Figs. 9 and 10). This is feasible as based on biostratigraphic data
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from the northern and central carbonate ramp of Peru (Benavides-Caceres, 1956; Wilson 1963; von
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Hillebrandt, 1970; Romani, 1982; Jaillard 1986, 1987; Jaillard and Arnaud-Vanneau, 1993; Robert
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2002; Robert and Bulot, 2004, 2005; Dhondt and Jaillard, 2005; Jaillard et al., 2005; Robert et al.,
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2009).
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The sequence stratigraphy approach used here is based on the work of Embry (2009) and Andrieu
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et al. (2015) for shallow marine environments. Essentially, depositional sequences recognized are
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bounded by sequences boundaries (SB), which represent either maximum regressive surfaces or 12
ACCEPTED MANUSCRIPT transgressive surfaces (van Wagoner et al., 1988). These surfaces indicate a significant shift from a
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shallowing-upward to a deepening-upward pattern (Embry, 2009; Andrieu et al., 2015). Moreover, a
295
depositional sequence is a cycle that comprises: transgressive deposits (TD) marking a change to
296
upwards deepening facies (retrograding structures); a maximum flooding interval (MFI), a portion of
297
the section that marks the transition from deepening-upward to shallowing-upward patterns; and a
298
highstand deposit (HD), representing shallowing facies grading upwards into more proximal facies
299
(prograding structures).
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The lowermost medium-scale sequence at Jumasha 1 (Fig. 9) consists of thickening-upward units
301
of argillaceous, nodular limestone built by facies 3a rich in oysters. This pattern may represent a
302
highstand system tract. These units were attributed to the upper middle Albian due to ammonite
303
associations such as Lyelliceras ulrichi, Oxytropidoceras douglasi (Benavides-Caceres, 1956; Wilson,
304
1963; von Hillebrandt, 1970; Jaillad, 1986). The two overlying medium-scale sequences (Fig. 9) show,
305
in their transgressive deposits, thickening-upward transitions ranging from facies types 2c, 3a, to
306
facies 3c. These units are followed by deepening trends to outer ramp sedimentation typified by facies
307
4a with abundant planktonic foraminifera, defining the maximum flooding interval in these two
308
sequences. The highstand deposits are represented by thickening-upward successions from facies 3c to
309
3a, increasingly enriched in quartz grains (facies 1a), oysters and serpulids. At section metre 115 m,
310
measured from the base of the Jumasha Formation at Uchucchacua, Romani (1982) reported Exogyra
311
costagyra cf olisoponensis and Merlingina cretacea of Cenomanian age. Echinoderm and oyster
312
limestones characterize the last sequence (122 to 156 m; Fig. 9).
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The upper Cenomanian–lower Turonian portion at Uchucchacua and Lauricocha is comprised in
314
the Jumasha 4 (Fig. 10), where is showed three medium-scale sequences (MsS1–3). These are part of
315
the three depositional sequences Ro1–3 sensu Jaillard and Arnaud-Vanneau (1993). The medium-scale
316
sequences 1–2 with thickness of 60 to 140 metres (MsS1–2; Fig. 10) build the Jumasha 4. The
317
sequence boundary that limits the top of the Jumasha 4 (Fig. 6F), displays a prominent stratigraphic
13
ACCEPTED MANUSCRIPT 318
feature, defining a major change in depositional style from a regressive to a major transgression event
319
in the early Turonian (Jaillard and Arnaud-Vanneau, 1993; Fig. 10). At the base of both sections, an abrupt deepening trend from nearshore restricted (facies 2d–h) to
321
middle ramp facies (3b) is observed through medium-scale sequence 1 and may define a transgressive
322
system tract in terms of the Jumasha 4. The maximum of this deepening interval, situated at around
323
section metre 35 in both sequences, is marked by an increase in facies associations 3c. This is
324
followed by a shallowing-upward trend recorded in the medium-scale sequences 1, indicated by a
325
thickening-upward transition facies 3c upwards 2d and presence of rotaliid foraminifera Rotorbinella
326
mesogeensis, representing the regressive trend of the medium-scale sequence 1. This is followed by a
327
deepening trend represented from facies 1a, 2d to facies 3b, which form the transgressive deposits of
328
the medium-scale sequence 2. The regressive deposits of the medium-scale sequence 2 are upsection
329
indicated by the appearance of facies 2h to 2d. Medium-scale sequence 3 (MsS3, Fig. 10) is at least 40
330
m thick (but not fully exposed). This sequence includes the lower Turonian and the Jumasha 5. Facies
331
are dominated by 3c, which is well-defined in the lower part of both sections, showing a middle ramp
332
environment rich in echinoids. The transition from facies 3c into the overlaying thickly-bedded
333
limestones of facies 2d is gradual, marking the highstand system tract.
335
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4.3 Carbon-isotope stratigraphy
The upper Albian to lower Cenomanian bulk micrite and organic carbon isotope stratigraphy at
337
Uchucchacua is documented in figure 9. At the lowermost part of the curve, the δ13Ccarb and δ13Corg
338
values vary between –0.3 and +2.7‰ and between –27.4 and –25.3‰. These are followed by a
339
significant positive shift from –0.2 to +3.1‰ (δ13Ccarb) and from –27.4 to –24.2‰ (δ13Corg), showing a
340
prominent positive chemostratigraphic feature. From about section metre 90 onwards (Fig. 9),
341
values decrease reaching 1.3‰ in δ13Ccarb and –26.8‰ in δ13Corg, followed by increasingly
342
enriched values reaching +2.9‰ and –24.2‰, respectively. The top of the prominent positive isotope
343
interval ends with a declining trend to a base level of +1.1‰ in δ13Ccarb and –26.7‰ in δ13Corg. 14
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13
C-
13
C
ACCEPTED MANUSCRIPT Based on existing biostratigraphic data in Central Peru (von Hillebrandt, 1970; Romani, 1982;
345
Jaillard and Arnaud-Vanneau, 1993) and the chemostratigraphic terminology used for OAE2 interval
346
(onset, trough, plateau, and recovery, c.f., Paul et al., 1999; Tsikos et al., 2004), and datum levels A,
347
B, C and D as used in Tethyan and North Atlantic sections (Meyers et al., 2012; Du Vivier et al., 2014,
348
2015), we characterize specific sub-intervals of the Cenomanian–Turonian isotope excursion for the
349
Uchucchacua and Lauricocha sections (Fig. 10).
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In terms of their δ13C amplitudes and trends, the Uchucchacua and Lauricocha isotope curves are
351
directly comparable. In both sections, the lower part of the prominent carbon isotope excursion related
352
to OAE2 was not covered due to a gap in exposure. However, the OAE2 interval at Uchucchacua
353
displays a significant rise from mean values of +3.7‰ to a first peak of +4.6‰ in δ13Ccarb and a high
354
ratio –22.6‰ in δ13Corg, whereas at Lauricocha the δ13Ccarb curve shows
355
specific segment is here interpreted to represent the upper part of the onset of OAE2. Upsection at
356
both locations, both bulk micrite and organic carbon ratios are characterized by decreasing 13C-values
357
of 3.3‰ in δ13Ccarb and –25‰ in δ13Corg, followed by a rapid trend to
358
+4.2‰ and –22.1‰, respectively. This pattern corresponds with a `trough´ according to the Tethyan
359
scheme (Du Vivier et al., 2014; Fig. 10). Thereafter, a plateau interval of δ13Ccarb and δ13Corg values
360
ranging between +1.78 to +3.3‰ and –23 to –20.9‰ is recorded. The recovery interval is
361
characterized by a rapid decline to a base level of +2.5‰ in δ13Ccarb and –25.9‰ in δ13Corg.
13
13
C depleted values. This
C-enriched values reaching
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363
5
Discussion
364
5.1 Environmental evolution of the Western Platform from the Albian towards the Turonian
365
Jaillard et al. (1990) suggested that the opening of the North Atlantic (Western Tethys) in the
366
Early Cretaceous lead to the formation of a Tethyan oceanic seaway that extended through the Palaeo-
367
Pacific realm. This interoceanic communication may explain the Tethyan affinities of some Pacific
368
provinces during the Aptian to Turonian (e.g., Iba and Sano, 2007; Robert, 2002; Robert and Bulot, 15
ACCEPTED MANUSCRIPT 2004; Dhondt and Jaillard, 2005). According to Robert (2002) and Robert and Bulot (2004), the
370
earliest Albian was dominated by migration of Tethyan Engonoceratidae specimens to the Western
371
Platform due to an Albian transgression (Robert, 2002). Cretaceous Tethys taxa have also been
372
recorded in the Berriasian to early Albian of the north-western Pacific, perhaps evidencing a Tethyan
373
biotic realm in the Pacific in that time (Iba and Sano, 2007; Robert and Bulot, 2004). In addition to
374
this palaeontological observation, bivalve taxa of the Albian-Cenomanian, recorded at Pongo de
375
Rentema in Peru, are characterized by taxa known from Tethyan faunas, North Africa, southwestern
376
Europe, and Texas (Dhondt and Jaillard, 2005). It is thus conceivable that water masses from the
377
North Atlantic (Western Tethys) reached the Western Platform during the Albian-Cenomanian
378
(Robert, 2002; Jaillard et al., 1990; Dhondt and Jaillard, 2005).
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Moreover, in the earliest middle Albian, the South America margin witnessed the onset of the
380
Huarmey-Trough volcanism, leading the depositions of large volumes of basaltic pillow lavas
381
(Atherton and Aguirre, 1992; Winter et al., 2010) . This volcanic activity has been temporally
382
associated to an increased convergence rate of the rifting of South America and Africa (Soler and
383
Bonhomme, 1990; Moulin et al., 2010; Winter et al., 2010). The uplift of the Huarmey-Trough may
384
have also isolated the Western Patform from vigorous exchange and mixing with water masses of the
385
Pacific Ocean in the middle Albian (Fig. 11A). The rifting of South America and Africa continued
386
through the Cenomanian, forming continuous deep-water connection between the North and South
387
Atlantic Ocean (Wagner and Pletsch, 1999; Poulsen et al., 2003). This tectonic event triggered a major
388
reorganization of the global shallow and deep oceanic circulations and regional climate changes
389
(Poulsen et al., 2003). Another feature related to this pattern is the presence of new species of
390
ammonites, bivalves, and benthic foraminifers, which prospered in shallow-marine environments in
391
the Pacific (e.g., Chinzei, 1986; Toshimitsu et al., 1990; Jaillard and Arnaud-Vanneau, 1993; Robert,
392
2002; Robert and Bulot, 2004; Dhondt and Jaillard, 2005; Iba and Sano, 2007).
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These new Peruvian species are exemplified by Glottoceras moorei and Engonoceras gr. stolleyi-
394
hilli Böhm from the middle Albian (Robert and Bulot, 2004). Bivalve communities also prospered in 16
ACCEPTED MANUSCRIPT the Western Platform in the late middle Albian to late Cenomanian (Dhondt and Jaillard, 2005). The
396
lack of fossil remains of reef constructers (Jaillard, 1987) suggests the establishment of a heterozoan
397
ramp devoid of any morphological features such as a rimmed platform morphology (Fig. 11A). The
398
absence of a rim or barrier may have caused rather uniform facies belts across wide portions of the
399
Western Platform, where sediment and organic material were transported and re-distributed by storm
400
action and currents (Navarro-Ramirez et al., 2015). The presence of argillaceous material in facies 1a
401
and 2a-b points to continent-derived sediments probably transported by rivers. According to Jaillard
402
(1987), during the late Albian to late Cenomanian, a diachronic progradation of terrigenous supratidal
403
facies was accentuated, resulting in a NE-oriented deltaic system fed by runoff from the Brazilian
404
continental basement (Fig. 11A). This may indicate the installation of high trophic levels and more
405
humid climate with all corresponding changes in sediment runoff and type on the Western Platform
406
(Fig. 11A).
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In the late Cenomanian towards the early Turonian, a key feature of Jumasha 4 recorded at
408
Uchucchacua and Lauricocha (central Peru) is the ubiquitous presence of Perouvianella peruviana-
409
dominated carbonate facies (Fig. 10). During this period, central Peru was isolated from the Pacific
410
and from eastern deltaic influx of the Brazilian continental basement due to the uplift of the Marañon
411
Massif during the early Cenomanian (Jaillard, 1986; Fig. 11B). Evidences of this tectonic event is
412
given by a shallowing maximum, associated with tectonic breccias and synsedinemtary normal faults,
413
occurring from the lower to the middle Cenomanian (Jaillard, 1986). Furthermore, the presence of a
414
morphological barrier is given by the lack of an argillaceous material in these sections (Jaillard, 1986,
415
1987), probably implying the installation of an oligotrophic environment (Fig. 11B). This later
416
palaeoceanographic feature might have triggered – or at least favoured – the mass occurrence of
417
Perouvianella peruviana and associated foraminifera, being only known from these Central Peruvian
418
rocks (Jaillard, 1986 and references therein). In contrast to northern Peru (Pongo de Rentema), benthic
419
taxa are represented by inoceramids such as Mytiloides mytiloides, M. opalensis, and Sergipia sp.
420
(Dhondt and Jaillard, 2005). The genus Sergipia has a wide distribution, being reported in Brazil,
AC C
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17
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Nigeria, Mexico, Japan, and France (Hessel, 1988; Jolet et al., 1997; Dhondt and Jaillard, 2005), and
422
indicates the installation of an open marine setting in northern Peru (Fig. 11B). Perouvianella peruviana is a porcelaneous benthic foraminifera with a complex shell architecture
424
and flattened shape. The divisions of chamber lumen below the lateral wall may indicate the loci for
425
endosymbionts. In recent oceans, the large discoidal porcelaneous Archaias hosts clorophycean algae
426
symbionts in their shells which restrict their habitat to the first 20 m of the water column. Besides,
427
Marginopora, Amphisorus and Sorites possess dinophycean algae as symbionts allowing them to
428
reach a life depth of about 50 m. (Leutenegger, 1984). On the level of a working hypothesis, similar
429
environmental conditions are inferred for Perouvianella peruviana. Concerning the reproduction cycle
430
of Perouvianella peruviana, in the fourth Jumasha sequence, two A-forms have been found (see Fig.
431
8I) but no B-forms are present. This may be compared with the population of Amphisorus hemprichii
432
from the Red Sea (Zohary et al., 1980). Dasylcadacean and Codiacean algae (Figs. 8A–B) in
433
combination with miliolids, hyaline and agglutinated foraminifera (Figs. 8E–G) associated with
434
Perouvianella peruviana support a local oligotrophic environment within the photic zone in which
435
these foraminifera thrived. A foraminifer that is perhaps similar to Perouvianella peruviana has been
436
found the Cenomanian of Mexico (Jaliscella sigali, Fourcade et al., 1990). The presence of
437
Perouvianella and Jaliscella on the American continent, and the lack of alveolinoideans (the genus
438
Sellialveolina indicated by Jaillard and Arnaud-Vanneau (1993) is not found in our samples) may
439
confirm an endemic Central Peruvian palaeobioprovince that is in contrast with the Tethyan biotic
440
realm for the late Cenomanian–early Turonian time interval.
442
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5.2 Evidence for a disturbed carbon cycle
443
Shallow water carbonates have been shown to record, under favourable conditions, first order
444
patterns of the global marine carbon isotope signature (Weissert et al., 1998; Immenhauser et al.,
445
2005; Föllmi et al., 2006; Huck et al., 2010, 2011, 2012, 2013, 2014; Andrieu et al., 2015). The global
446
seawater δ13CDIC preserved in the rock record are related to changes in the ratio of marine carbonate 18
ACCEPTED MANUSCRIPT 447
carbon and organic carbon burial (Scholle and Arthur, 1980; Weissert, 1989; Immenhauser et al.,
448
2008). With reference to the Peruvian sections documented here, we argue that the mid-Cretaceous low-
450
amplitude sea-level fall was insufficient to expose significant portions of the carbonate ramp studied.
451
Reasons for that might include rapid basement subsidence, a feature that is supported by the
452
stratigraphically thick successions found for the OAE2 interval (Ucchucchacua = 210 m and
453
Cajamarca composite section = 52 m; Fig. 12). Further evidence for rapid basement subsidence comes
454
from the remarkable lack of karst-related features, bleaching, isotopic shifts or other patterns assigned
455
to meteoric diagenesis (Allan and Matthews, 1977; Christ et al., 2012; Huck et al., 2014) beneath
456
discontinuity surfaces recorded in the sections studied. This may imply only weak or no meteoric
457
diagenesis and suggest that the carbonates studied here experienced mainly marine and subsequent
458
burial diagenesis. Judging from the direct comparability of chemostratigraphic sections in Peru with
459
such in many other locations worldwide (Sageman et al., 2006; Jarvis et al., 2011), we tentatively
460
suggest that most of these carbonates stabilized in the presence of marine pore waters. This is turn
461
implies that at least the chemostratigraphic patterns, but perhaps not absolute isotope values, represent
462
palaeoceanographic features as opposed to diagenetic patterns. The question to which degree these
463
patterns reflect global features versus regional water mass properties merits discussion.
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In the Tethyan Ocean, OAE1d is associated to carbon isotope shifts with amplitudes between 0.5
465
and 1.5‰ (e.g., Bornemann et al., 2005; Gambacorta et al., 2015 and references therein; Fig. 9).
466
Despite the limitations of the biostratigraphic control used here, the δ13Ccarb and δ13Corg records at
467
Uchucchacua indicated shifts of 3‰ during the late Albian being more pronounced than those
468
recorded in the Cajamarca composite section (1.5‰ δ13Ccarb and 2.5‰ for δ13Corg; Fig. 12). The
469
carbonates facies, within the limitations of facies characterization, are near-identical in both areas (Fig.
470
11A). This implies that facies-control on chemostratigraphy is weak to absent. It is likely that the
471
regional offset in δ13Ccarb ratio amplitude may have been induced by differences in the water mass
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19
ACCEPTED MANUSCRIPT 472
residence time or differential
473
Immenhauser et al., 2008).
13
C sources (Lloyd, 1964; Holmden et al., 1998; Saltzman, 2003;
For OAE2, the contrast-comparison of δ13Ccarb and δ13Corg records from the Uchucchacua and
475
Lauricocha sections (situated 36 km apart) agrees with the notion of mainly palaeoceanographic
476
patterns and a weak diagenetic overprint (Fig. 10). However, discrepancies exist in the amplitudes of
477
the δ13Ccarb and δ13Corg chemostratigraphic curves when these are compared with the reference section
478
of the Cajamarca region situated approximately 500 km to the north (Fig. 12; Navarro-Ramirez et al.,
479
2016). The differential amplitudes of δ13Ccarb and δ13Corg records are more pronounced in Cajamarca
480
location than in sections at the Uchucchacua and Lauricocha sites. Probably, differences in
481
palaeogeographic settings significantly affected or overprinted global patterns at Uchucchacua and
482
Lauricocha as indicated by mesotrophic and oligotrophic environments in northern and central Peru,
483
respectively (Fig. 11B).
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484 5.2.1
Western Platform patterns during the mid-Cretaceous OAEs
TE D
485
In the Western Platform domain, the uppermost Aptian Jacob Level (1st black-shale level of the
487
OAE1b set), which is associated to relative sea-level fall near the Aptian–Albian boundary, caused
488
erosional surfaces, disconformities, basal conglomerates, and ferruginous deposits as reported at the
489
base of the Inca Formation (Jaillard, 1987; Fig. 12). This relative sea-level fall had an amplitude of at
490
least 50m (Maurer et al., 2013) and is probably linked to global cooling and ice sheet dynamics at
491
high-latitudes (Frakes and Francis, 1988; Weissert and Lini, 1991; Frakes et al., 1995; Bodin et al.,
492
2015) responding to global scale and prolonged episodes of organic carbon (Menegatti et al., 1998;
493
Gröcke, 1998; Bralower et al., 1999; Herrle et al., 2004; Westermann et al., 2010; McAnena et al.,
494
2013; Bodin et al., 2015). For this time interval, decreased rates of continental weathering as
495
suggested by low 87Sr/86Sr values and the coeval positive shift of δ18O have been proposed (Bodin et
496
al., 2015).
AC C
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ACCEPTED MANUSCRIPT The Kilian and Paquier levels (2nd and 3rd black-shale levels of the OAE1b set) led to transition
498
stages from a siliciclastic-dominated sedimentation to a heterozoan-neritic carbonate factory, followed
499
by incipient phases of platform demise of the Tethyan biotic realm in the Pacific (Robert and Bulot,
500
2004; Iba and Sano, 2007). The earliest Albian corresponds to a change from low to high sea-level,
501
marked by a transition between coldhouse to greenhouse climate mode (Bodin et al., 2015). This
502
climatic pattern is perhaps due to the activity of the South Kerguelen plateau around 112–110 Ma
503
(Coffin et al., 2002; Duncan, 2002; Frey et al., 2003; Bryan and Ferrari, 2013). Early Albian global
504
warming led to a transgressive event that transported water masses from the North Atlantic (Western
505
Tethys) onto the neritic domain of the Western Platform. Moreover, global warming also shifted the
506
Inter-Tropical Convergence Zone (ITCZ) towards the equator, installing more humid climate in the
507
Western Platform domain (Hay and Flögel, 2012; Hasegawa et al., 2012). The enhanced hydrological
508
cycle may have increased continental weathering of the Brazilian continental shield, delivering
509
nutrients and clay-sized material to the ramp, promoting high productivity of benthic organisms
510
(oysters, gastropods, and echinoids). High productivity and an abrupt change in lithology and biota
511
have been associated to the OAE1b set around the globe, indicating the significance of this event
512
during the Cretaceous (Erbacher et al., 1996, 1998, 1999; 2001; Leckie et al., 2002).
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The Leenhardt level (i.e., the last black-shale level of the OAE1b set) is contemporaneous to a
514
major demise of the Cretaceous Tethyan taxa in the Pacific and neritic carbonate production. In the
515
field in Peru, this feature is marked by a regional discontinuity and the disappearance of all lower
516
Albian ammonite species with the exception Glottoceras crassinodosum Sommermeier (Robert, and
517
Bulot, 2004). High atmospheric CO2 output by the South Kerguelen plateau reduced carbonate
518
availability and seawater pH, thereby causing demise episodes of the neritic carbonate production, as
519
well as the loss of calcareous plankton during the broad interval of the OAE1b set (Leckie et al.,
520
2002).
AC C
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521
In the middle Albian, the increased flux of isotopically light CO2 at times of the activation of the
522
western South America volcanic arc, shifted the δ13Ccarb towards more negative values as perhaps 21
ACCEPTED MANUSCRIPT 523
documented in the Cajamarca composite section (Navarro-Ramirez et al., 2015, 2016; Fig. 12). This
524
feature is evidenced by reduced carbonate production in the Western Platform domain. The Albian global warming continued until the early Turonian, reaching a mid-Cretaceous climate
526
maximum (Pagani et al., 2014). This led to a rapid diversification and a pronounced increase in the
527
degree of calcification of the planktic foraminifera, increased upper water column stratification,
528
initiation of widespread chalk deposition and oligotrophic conditions in (hemi-) pelagic settings in the
529
Atlantic and Tethys (Leckie, 1989; Premoli Silva and Sliter, 1999; Leckie et al., 2002). In contrast,
530
global warming resulted in an enhanced hydrodynamic circle with erosion delivering nutrients and
531
promoting diversification of benthic communities dominated by oyster bioherms and large endemic
532
benthonic foraminifera in the Western Platform (Fig. 11A). On the other hand, in the Tethyan realm
533
and elsewhere, OAE1d, MCEI, and OAE2 have been associated to high turnover rates affecting the
534
planktic foraminifera and radiolarians, anoxia, organic-rich deposition, and demise of the carbonate
535
platform (e.g., Erbacher et al., 1996, 2001; Gale et al., 2000; Wilson and Norris, 2001; Leckie et al.,
536
2002; Bornemann et al., 2005; Gertsch et al., 2010; Schröder-Adams et al., 2012; Scott et al., 2013;
537
Owens et al., 2013; Giraud et al. 2013; Melinte-Dobrinescu et al., 2015; Gambacorta et al., 2015;
538
Andrieu et al., 2015). These patterns are in clear contrast to phases of enhanced carbonate deposition
539
documented for the sections in Peru discussed here. Persistent deposition of benthic communities in
540
these sections supports the notion of well-oxygenated water masses throughout the middle Albian–
541
early Turonian interval.
AC C
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542
In the Tethyan realm, several groups of foraminifera such as the alveolinids, rhapydionids,
543
pseudorahpydioninids, and nezzazatids characterize the Cenomanian (see Schroeder and Neumann,
544
1985; Calonge et al., 2002; Vicedo et al., 2011; Consorti et al., 2015 and the work cited therein). All of
545
these foraminifera disappear at the Cenomanian–Turonian boundary. The finding of Perouvianella
546
peruviana in Jumasha 4 (Fig. 10) in the central Peru sections is here considered evidence for the
547
continuation of a stable, albeit endemic, marine ecosystem during and after OAE2 contrasting the
548
Tethyan biotic patterns (Parente et al., 2008; Fig. 12). In combination with similar findings in neritic 22
ACCEPTED MANUSCRIPT 549
carbonates in the Pacific - lacking evidence of anoxic conditions during OAE2 (Elrick et al., 2009;
550
Takashima et al., 2011) - this suggests that anoxic conditions did not establish in major portions of the
551
Pacific Ocean during OAE2 (Takashima et al., 2011). In the Cajamarca composite section, a prominent negative shift is observed at the onset of OAE2
553
(Fig. 12). This negative δ13C excursion has also been identified in other sections in the Pacific realm
554
(California, USA; Hokkaido, Japan; Takashima et al., 2011). According to Takashima et al. (2011),
555
this negative δ13C excursion is linked to an increase in pCO2 induced by the formation of the
556
Caribbean large igneous province (LIP; ~95.1–92.2 Ma, Snow et al., 2005). Arguments for the claim
557
that the Caribbean massive volcanism triggered OAE2 come from a negative 34Ssulphate shift ca. 500 kyr
558
before OAE2 (Adams et al., 2010) as well as from a negative
559
2014) and increased pCO2 levels (e.g., Barclay et al., 2010).
SC
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552
Os/188Os shift (e.g., Du Vivier et al.,
M AN U
187
This significant increase in pCO2 drove the climate to a hothouse mode (Kidder and Worsley
561
2010), perhaps leading to an acidified ocean (Du Vivier et al., 2015) and a perturbed hydrological
562
cycle, that enhanced weathering rates at the beginning of the onset of OAE2 (Frijia and Parente, 2008;
563
Blätter et al., 2011; van Helmond et al., 2013; Pogge von Strandmann et al., 2013). The beginning of
564
the onset of OAE2 in the Western Platform is coeval to an increased influx of argillaceous material
565
and reduced carbonate production in northern Peru (Fig. 11B). However, continuous deposition of
566
floatstones and grainstones containing Perouvianella peruviana is observed in the central Peru
567
sections (Fig. 10) possibly due to the restriction of continental nutrient influx (Fig. 11B). Consumption
568
of significant amount of pCO2 by the enhanced global organic carbon burial and by the continental
569
weathering after the onset of the Caribbean LIP (Sinninghe Damsté et al., 2010; Jarvis et al., 2011; van
570
Helmond et al., 2013; Pogge von Strandmann et al., 2013) conditioned the oceans to gradually be
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more oxic (Pogge von Strandmann et al., 2013) and might have caused the reduction of benthic
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communities in the Western Platform of Peru after OAE2 (Figs. 10 and 12).
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Conclusions The mid-Cretaceous sections of the Western Platform of Peru document a predominance of
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oyster-rich, mixed siliciclastic-carbonate facies in the middle Albian–late Cenomanian. Regional
577
evidence suggests that this rather uniform facies belt expanded across wide portions of the spatially
578
very extensive carbonate ramp in northern and central Peru. In the late Cenomanian and the early
579
Turonian, the deposition of an endemic large benthic foraminifera facies (Perouvianella peruviana) is
580
a key facies element that continues throughout the OAE2 interval without discernible adaptation to
581
environmental change. During this time, central Peru was isolated from the Pacific water masses and
582
from deltaic influx from the emerged shield of Brazil in the east. This specific palaeoceanographic
583
setting might have promoted the mass occurrence of Perouvianella peruviana. Probably due to the
584
complex setting of the Western Platform in the mid-Cretaceous, clear evidence for a carbonate crisis
585
associated with OAE 1d and OAE 2 in the sections of the central Peru is lacking.
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Global warming during the Albian–early Turonian resulted in a more humid climate, high trophic
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levels, and a diversification of benthic communities in Western Platform carbonates, influenced by
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palaeogeographic modifications. Even though the benthic communities thrived during extremely warm
589
conditions in the late middle Albian and towards the early Turonian, they were exposed to high
590
nutrient levels caused by a perturbed hydrological cycle at the beginning of the onset OAE2. The
591
intensified chemical weathering on the Brazilian shield during the mid-Cretaceous may have
592
contributed to trigger the OAE in the North Atlantic nutrient trap. The work documented here is a
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clear case example for the superposition of regional and global patterns in palaeoceanography and
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complements scarce existing data from South America for this critical interval in Earth’s history.
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Acknowledgements
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ACCEPTED MANUSCRIPT This project was supported by the Deutsche Forschungsgemeinschaft (DFG, project n° BO-365/2-1)
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and by the Deutscher Akademischer Austauschdienst (DAAD) through a scholarship to J.P.N. (PKZ:
600
91540654). We thank the Geological survey of Peru (INGEMMET) for important logistical support.
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Esmeralda Caus (Barcelona) is gratefully acknowledged for a critical reading of the text in its early
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version. Analytical work was performed in the isotope laboratories at Bochum and Erlangen-
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Nuremberg. We greatly acknowledge the editorial advice of E. Koutsoukos and the comments of two
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anonymous journal reviewers.
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Figure Captions
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Figure 1: Palaeogeographic reconstruction for the mid-Cretaceous (modified after Blakey, 2011)
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indicating the position of what today is the Western Platform of South America (yellow arrow). Figure 2: Geological map of the Oyon region (modified after INGEMMET, 1980) showing the location of the Uchucchacua and Lauricocha sections documented in this study. Figure 3: Regional correlation of the sequence stratigraphic interpretation between the Cajamarca
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composite section (northern Peru; Navarro-Ramirez et al., 2016) and the Uchucchacua section (central
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Peru; this study). The Uchucchacua section is juxtaposed with our sections (Jumasha 1 and 4) and by
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the Jaillard (1986)´s work. Main biostratigraphic data are indicated (Benavides-Caceres, 1956; Wilson
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1963; von Hillebrandt, 1970; Romani, 1982; Jaillard 1986, 1987; Jaillard and Arnaud-Vanneau, 1993;
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Robert 2002; Robert and Bulot, 2004, 2005; Dhondt and Jaillard, 2005; Jaillard et al., 2005; Robert et
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al., 2009).
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Figure 4: Legend for figs. 3, 9 and 10, denoting colour codes for different facies types.
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Figure 5: Landscape images showing A): The Jumasha Formation of the upper Albian–lower
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Turonian at Uchucchacua, Oyon region, Central Peru, built by massive, thickly-bedded, light-grey
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limestones. B) Lower part of the Jumasha Formation at Uchucchacua, including the upper Albian and
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the lower Cenomanian rocks. C) The upper part of the Jumasha Formation at Uchucchacua, including 42
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the middle Cenomanian to lower Turonian rocks. D) The upper part of the Jumasha Formation at
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Lauricocha, including the late Cenomanian to lower Turonian rocks. Figure 6: Field images taken at the Uchucchacua section. A) Thin lamination of facies 1a,
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represented by grey and argillaceous limestones intercalations. B) Nodular limestones of the facies 2e.
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C) Thickly-bedded limestones of facies 2g. D) Thickly-bedded limestones of medium-scale sequence
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10 (facies 2b). C) Floatstones of facies 3d. F) Bioturbation at the level of regressive deposits in the top
1019
of Jumasha 4.
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Figure 7A): Thin section images documenting grainstones of facies 2d, represented by peloids,
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miliolids, and isolated specimen of Perouvianella peruviana. B) Dasycladales-bearing grainstones of
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facies 2f. C) Miliolids and Perouvianella peruviana of facies 2g. D) Grain-rudstones dominated by
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Peruvianella peruviana in facies 2h. E) Diverse fauna floatstones of facies 3b, characterized by a
1024
chaotic fabric. F) Packstone rich in echinoids of facies 3c.
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Figure 8: Upper Cenomanian‒lower Turonian microfossil from shallow-water deposits of the
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Jumasha Formation in central Peru. Scale bar in images C and D is 0.5 mm, in all other images the
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scale bar is 1 mm. A) and B) show dasycladales, closely related to the genus Trinocladus. A.
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Transversal section. B. Longitudinal section (central part) and transversal section (upper right part). C)
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and D) Rotorbinella mesogeensis; C. axial section, note the presence of central plug. D. equatorial
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section. E) Agglutinated foraminifera in equatorial section. F) Porcelaneous foraminifera in equatorial
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section. G) Agglutinated foraminifera with spiral outline closely related to the genus Merlingina
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cretacea Hamaoui and Saint Marc, 1965. H) Perouvianella peroviana in perfect equatorial section. I)
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Perouvianella peroviana proloculi of two specimens, difference in diameter suggest the existence of
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multiple cycles of reproduction within the population.
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Figure 9: Sequence stratigraphic interpretation and carbon-isotope stratigraphy of the
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Uchucchacua section for the upper Albian (Jumasha 1). Correlation with the Tethyan composite curve
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of Herrle et al. (2015) is shown. Biostratigraphic data for the late Albian have been obtained from
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Jaillard (1986, 1987 and references therein). 43
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Uchucchacua and Lauricocha sections for the late Cenomanian (Jumasha 4). Data points shaded grey
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correspond to OAE2 based on bio- and chemostratigraphic data. Biostratigraphic data of the
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Uchucchacua section has been taken from Jaillard (1986) and Jaillard and Arnaud-Vanneau (1993).
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The chemostratigraphic markers of OAE2 (A‒B) include Navarro-Ramirez et al. (2016), Eastbourne
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(Jarvis et al., 2011), and the Portland # 1 core (Sageman et al., 2006).
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Figure 11A): Palaeogeographic interpretation of the western South America platform during the
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late Albian, characterized by oyster-rich mixed siliciclastic-carbonate deposition. Note regional oyster
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facies deposition, diachronic progradation of the NE deltaic system fed from the Brazilian continental
1048
basement, as well as the Huarmey Trough isolating the western Platform from the Pacific Ocean. In
1049
contrast to the central Peru, northern Peru was localized more proximal to the southern margin of the
1050
proto-Atlantic, resulting in a potentially stratified water column setting. B) During the late
1051
Cenomanian, the Western Platform was characterized by two distinctive palaeogeographic settings: (i)
1052
northern Peru characterized by continental deltaic influx from the emerged shield of Brazil in the east;
1053
(ii) central Peru representing an oligotrophic environment, leading to the mass occurrence of
1054
Perouvianella peruviana and associated miliolids. The four main depositional environments, each
1055
with a number of standard facies types (1a through 4a), are established for the inner to outer ramp.
1056
Interpretations are based on Benavides-Caceres (1956), Jaillard (1986, 1987), Atherton and Webb
1057
(1989), Soler and Bonhomme (1990), Jaillard et al. (2000), Callot et al. (2008), Robert et al. (2009)
1058
and Winter et al. (2010).
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Figure 12: Regional correlation of the carbon-isotope stratigraphy between the Cajamarca
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composite section (northern Peru; Navarro-Ramirez et al., 2016) and by the Uchucchacua section
1061
(central Peru), exhibiting the Jumasha 1 and 4 carbon-isotope data. Main OAEs and biostratigraphic
1062
data are indicated (Benavides-Caceres, 1956; Jaillard, 1986, 1987; Jaillard and Arnaud-Vanneau,
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1993; Dhondt and Jaillard, 2005; Robert et al., 2009). The chemostratigraphic terminology for OAE2
1064
interval (onset, trough, plateau, and recovery, c.f., Paul et al., 1999; Tsikos et al., 2004), and datum 44
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levels A, B, C and D as is used in Tethyan and North Atlantic sections (Meyers et al., 2012; Du Vivier
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et al., 2014, 2015). The red data points in the Cajamarca composite section form part of this study. Table 1: Overview of facies classification and interpretation. Numbers indicate the relative
1068
abundance of non-skeletal and skeletal components: 0 = absent, 1 = present, 2 = frequent, 3 =
1069
abundant, 4 = dominant.
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New Albian–Cenomanian and Cenomanian–Turonian carbon isotope curves based on epeiric-neritic successions from Central Peru Evidence for the impact of transient carbon cycle perturbations (OAE1d and OAE2) is provided
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Clear evidence for a carbonate neritic crisis associated with OAE 1d and OAE 2 is lacking