Impact of Middle Cretaceous Climatic Change on Sub- Equatorial Pacific Settings (Northern and Central Andes, Peru): a Comparison with the Tethyan Realm

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Impact of Middle Cretaceous climatic change on sub- equatorial Pacific settings (Northern and Central Andes, Peru): A comparison with the Tethyan Realm

Dissertation zur Erlangung des akademischen Grades eines Doktors der Naturwissenschaften an der Fakultät für Geowissenschaften der Ruhr-Universität

Bochum

vorgelegt von Juan-Pablo Navarro-Ramirez geboren am 12.03.1985 in Sullana, Peru Bochum,

November 2015

Die vorliegende Arbeit wurde von der Fakultät für Geowissenschaften der Ruhr-Universität Bochum als Dissertation zur Erlangung des Grades eines Doktors der Naturwissenschaften (Dr. rer. nat.) anerkannt.

Erster Gutachter: Prof. Dr. Adrian Immenhauser

Zweiter Gutachter: Prof. Dr. Jörg Mutterlose fachfremder Gutachter: Prof. Dr. Stephane Wohnlich

Tag der Disputation: 01.02.2016

Erklärung

„Ich erkläre hiermit an Eides statt, dass ich die vorliegende Arbeit selbständig angefertigt sowie die benutzten Quellen und Hilfsmittel vollständig angegeben habe. Ich habe alle Fakten, Textstellen und Abbildungen, die anderen Werken dem Wortlaut oder dem Sinn nach entnommen sind, durch entsprechende Zitate gekennzeichnet. Die vorliegende Dissertation wurde in dieser oder ähnlicher Form bei keiner anderen Fakultät oder Hochschule eingereicht.“

Bochum, November 2015

Juan Pablo Navarro Ramirez

Publication list of the Dissertation

1. Title: Record of Albian to early Cenomanian environmental perturbation in the eastern sub-equatorial Pacific

Authors: Navarro-Ramirez J.P.1, Bodin S.1, Heimhofer U.2, Immenhauser A.1

Published in Palaeogeography, Palaeoclimatology, Palaeoecology in January 2015;

Volume 423; Pages 122–137 (chapter 3 of this thesis)

DOI: //dx.doi.org/10.1016/j.palaeo.2015.01.025

2. Title: Ongoing Cenomanian – Turonian heterozoan carbonate production in the neritic settings of Peru

Authors: Navarro-Ramirez J.P.1,, Bodin S.1, Immenhauser A.1

Published in Sedimentary Geology in October 2015 (chapter 4 of this thesis)

Article in Press

DOI: 10.1016/j.sedgeo.2015.10.011

3. Title: Response of western South American epeiric-neritic heterozoan ecosystem to OAE 1d and 2

Authors: Navarro-Ramirez J.P.1,, Bodin S.1, Immenhauser A.1

In preparation (chapter 5 of this thesis)

1 Ruhr-Universität Bochum, Institut für Geologie, Mineralogie und Geophysik, D-44870 Bochum, Germany

2 Leibniz Universität Hannover, Institut für Geologie, D-30167 Hannover, Germany

The main applicant for the research grant was Prof. Dr. S.Bodin and the co-applicant was Prof. Dr. A. Immenhauser. The research was financed by the Deutsche Forschungsgemeinschaft (DFG, project no. BO-365/2-1) and by the Deutscher Akademischer Austauschdienst (DAAD) through a scholarship to me (PKZ: 91540654). Analytical work in the isotope laboratories at Bochum University was supported by A. Niedermayr. The Geological survey of Peru (INGEMMET) supported with important logistical support.

Author contributions and peer review process

First article:

J.P.N.R., S.B., and A.I. had initial discussions and were involved in the study design; J.P.N.R. wrote the first draft of the manuscript, implemented co-author comments, made the figures, and prepared the samples for stable isotope and total organic carbon analyses; U.H. was involved in the stable isotope analyses on total organic carbon at Leibniz University Hannover; J.P.N.R., S.B., and A.I. were involved in fieldwork data and sample (bulk rock) acquisition; All authors discussed the results and contributed to revisions of the manuscript. The article was first submitted in July 2014, evaluated by the editor F. Surlyk and by two anonymous reviewers. Ending of January 2015 the paper was accepted.

Second article:

J.P.N.R., S.B., and A.I. had initial discussions and were involved in the study design; J.P.N.R. wrote the first draft of the manuscript, implemented co-author comments, made the figures, and prepared the samples for stable isotope and total organic carbon analyses; J.P.N.R., S.B., and A.I. were involved in fieldwork data and sample (bulk rock) acquisition; All authors discussed the results and contributed to revisions of the manuscript. The article was first submitted in July 2015, evaluated by the editor Prof. Jones and by the reviewers T. Adatte and B. Sageman. On 23 October 2015 the paper was accepted.

Third article:

J.P.N.R., S.B., and A.I. had initial discussions and were involved in the study design; J.P.N.R. wrote the first draft of the manuscript, implemented co-author comments, made the figures, and prepared the samples for stable isotope and total organic carbon analyses; J.P.N.R., S.B., and A.I. were involved in fieldwork data and sample (bulk rock) acquisition; All authors discussed the results and contributed to revisions of the manuscript. The article is in preparation and intended to be sent to Sedimentology.

Bochum, November 2015

Juan Pablo Navarro Ramirez

Abstract

The Albian–Turonian interval (ca 113–90 Ma; mid-Cretaceous) is characterized by series of global environmental perturbations referred to as “oceanic anoxic events” (OAEs). OAEs are represented by an expanded and intensified oxygen minimum zone in the world oceans and coincide well with periods of disturbed carbon cycle. OAEs are often related to enhanced accumulation of black shale in deeper and basinal settings, anomalously high burial rates of marine organic carbon, sea-level fluctuations and changes in trophic levels. A linkage between massive volcanism and OAEs has been suggested. To date, much of the present knowledge of oceanic anoxia in the mid-Cretaceous world is biased towards data derived from hemipelagic and pelagic sections in Europe (Tethys), North America (Western Interior Seaway) and data from various ocean drilling projects. Less attention has been paid to the impact of mid-Cretaceous oceanic anoxic events to shallow water carbonate systems, and the sub-equatorial eastern Pacific domain is particularly underrepresented. In this thesis I argue that the Cretaceous epeiric depositional environments in western South America (here Peru) arguably bear crucial information to understand OAEs causes and consequences.

In order to close this gap, a field-based project has been performed in Peru. This study is based on a multidisciplinary approach making used of a broad, multi-proxy approach (C, and Sr isotopes, sedimentology, sequence stratigraphy etc.). Geochemical and sedimentological analyses were obtained from 1600 section metres of Albian–Turonian sedimentary rocks, logged in a bed-by-bed approach in six sections in the northern and central part of the Andes Mountain in Peru. The results are compared and discussed to coeval findings from the proto- Atlantic and Tethyan realms in order to understand causal linkages among global geological patterns during the mid-Cretaceous greenhouse world.

The Albian–Turonian interval in Peru (western platform) is characterized by the following ecological and environmental patterns: (i) an late Aptian–early Albian major change from siliciclastic-dominated to carbonate sedimentation coinciding with the impact of the Kilian Level; (ii) an early Albian incipient platform drowning linked to the impact of the Paquier Level; (iii) an early middle Albian major demise of neritic carbonate production that coincides with the Leenhardt Level, followed by middle Albian condensed sedimentation that

13 is associated with prominent negative values in δ Ccarb prior to the onset of OAE1c; (iv) renewed carbonate ramp production during the late Albian associated to the impact of OAE1d; (v) early to late Cenomanian heterozoan ramp recording the pre-OAE2 δ13C excursions, specifically the Mid-Cenomanian Event; (vi) a late Cenomanian interval typified by outer-ramp heterozoan type sedimentation prior to the δ13C-trough of the OAE2 interval. i

(vii) a late Cenomanian to early Turonian δ13C plateau phase characterized by benthonic inner ramp sedimentation during a phase of sea-level highstand; (viii) a recovery of δ13C values at the end of OAE2 associated to increased influx of argillaceous facies and reduced carbonate production; (ix) and finally an early to middle Turonian fluctuating δ13C curve, linked to a maximum flooding phase in the Mammites nodosoides Zone and enhanced carbonate production during the Collignoniceras woollgar Zone.

Results compiled in this thesis are relevant as they come from very expanded neritic sections in the sub-equatorial eastern Pacific domain of the South America Platform. In summary, the present thesis underlines the importance of transient environmental perturbations as a possible mechanism to affect neritic carbonate systems during the Albian– Turonian. Changes in depositional style and carbon isotope signatures are possibly affected by the very specific regional palaeogeographic setting of these sections, i.e., the vast but topographically compartmentalized carbonate ramp of the western platform. Regional features also include high rates of basement subsidence and related relative sea-level change, and transient intervals of increased continental run-off under humid climate conditions from the exposed continental shield to the east (present-day Brazil). Despite these regionally important features, however, the sections measured in Peru record clear evidence of global environmental patterns during the mid-Cretaceous. The very expanded nature of these sections allows for a uncommonly detailed temporal resolution of chemostratigraphic features.

Kurzfassung

Das Albium-Turonium Intervall (ca. 113-90; mittlere Kreide) ist charakterisiert durch eine Reihe von globalen Umweltereignissen, bekannt als „oceanic anoxic events“ (OAEs). OAEs sind gekennzeichnet durch eine ausgedehntere und intensivere Sauerstoff Minimum Zone, die zeitlich gut mit Perioden eines gestörten Kohlenstoffzyklus‘ zusammenfällt. OAEs sind oft verbunden mit einer erhöhten Ablagerung von Schwarzschiefern, anormal hoher Einlagerungsraten von marinem ozeanischen Kohlenstoff, Meeresspiegelschwankungen und Veränderungen im Nährstoffgehalt. Eine Verbindung zwischen hohen Vulkanismus und OAEs wird vermutet. Bisher ist vieles des aktuellen Kenntnisstandes über ozeanische Anoxia in der mittleren kreidezeitlichen Welt einseitig über Daten gewonnen worden, die aus hemipelagischen und pelagischen Abschnitten in Europa (Tethys), Nord Amerika (Western Interior Seaway) und von verschiedenen Meeresbohrungen stammen. Weniger Aufmerksamkeit wurde dem Einfluss von kreidezeitlichen ozeanischen anoxischen Events auf Karbonatsysteme des Flachwassers geschenkt und die sub-äquatorial, ostpazifischen Gebiete sind unterrepräsentiert. In dieser Arbeit lege ich dar, dass kreidezeiltiche Randmeerablagerungsbedingungen im westlichen Südamerika (hier Peru) wohl entscheidende Informationen liefern um Ursachen und Konsequenzen von OAEs zu verstehen.

Um diese Lücke zu schließen, wurde ein feldbasiertes Projekt in Peru durchgeführt. Diese Studie basiert auf einer multidisziplinären Vorgehensweise, die eine weite, multi-Proxy Methode (C und Sr Isotope, Sedimentologie, Sequenzstratigraphie etc.) nutzt. Geochemische und sedimentologische Untersuchungen wurden an einer 1600 Meter Sektion von Alb-Turon Sedimentgesteinen vorgenommen, die über eine Schicht-für-Schicht Methode an sechs Lokalitäten in den nördlichen und zentralen Anden Perus aufgenommen wurden. Die Ergebnisse wurden mit Funden der gleichen Zeit des Proto-Atlantiks und des Tethysschelfs verglichen und diskutiert um kausale Zusammenhänge in globalen geologischen Modellen während der kreidezeitlichen Treibhauswelt zu verstehen.

Das Albium-Turonium Interval in Peru (westliche Plattform) ist charakterisiert durch folgende ökologisches Modell: (i) Ein deutlicher Wechsel von siliziklatisch dominierter zu karbonatischer Sedimentation im späten Apt bis frühen Alb übereinstimmend mit der Einwirkung des Kilian Level; (ii) ein einsetzendes Ertrinken der Plattform im früheren Alb verbunden mit der Einwirkung des Paquier Levels; (iii) ein bedeutender Rückgang der Karbonatproduktion im Flachwasser im frühen Mittelalb übereinstimmend mit dem Leenhardt Level, gefolgt von einer kondensierten Sedimentation im Mittelalb, die mit deutlich negativen Werten im d13Ccarb verbunden ist, welche dem Einsetzten des OAE1c iii

vorausgehen; (iv) erneute Karbonatrampenproduktion während des späten Albs übereinstimmend mit der Einwirkung des OAE1d; (v) heterozoische Rampen des frühen bis späten Cenoman, insbesondere mittel cenomanische Events, erfassen den Prä-OAE2-δ13C- Trend; (vi) ein Intervall des späten Cenoman wird typisiert durch heterozoiche Sedimentation außerhalb der Rampe, welche dem d13C-Tief des OAE2 Intervalls vorausgeht; (vii) eine δ13C Plateau-Phase im späten Cenoman bis frühen Turon, die durch benthonische Sedimentation während einer Phase eines Meeresspiegelhochstandes innerhalb der Plattform charakterisiert ist; (viii) ein Wiederanstieg der δ13C Werte am Ende des OAE2 verbunden im einem ansteigenden Zustrom tonhaltiger Fazies und reduzierter Karbonatproduktion; (ix) und schließlich einer fluktuierende d13C Kurve im frühen bis mittel Turon, verbunden mit einer „maximum flooding phase“ in der M. nodosoides Zone und erhöhte Karbonatproduktion während der C. woollgari Zone.

Die zusammengestellten Ergebnisse in dieser Arbeit sind relevant, weil sie von sehr ausgedehnten Flachwasserabschnitten im sub-äquatorialen ostpazifischen Gebiet der Südamerikaplattform stammen. Zusammengefasst unterstreicht diese Arbeit die Wichtigkeit von kurzzeitigen ökologischen Perturbationen als mögliche Mechanismen, die Flachwasserkarbonatsysteme während des Alb bis Turons beeinflussen. Veränderungen in der Ablagerungsart und der Kohlenstoffisotopensignatur sind möglicherweise durch eine sehr spezifische regionale paläogeographische Umgebung in diesen Bereichen beeinflusst, z.B. die enorme aber topographisch aufgeteilte Karbonatrampe der westlichen Plattform. Regionale Besonderheiten beinhalten auch hohe Raten in Meeresbodenabsenkungen und damit verbundener relativen Meeresspiegelschwankungen und kurzzeitige Intervalle erhöhtem kontinentalen Abflusses unter humiden klimatischen Bedingungen des herausgehobenen kontinentalen Schildes im Osten (heutiges Brasilien). Trotz dieser regional wichtigen Besonderheiten, zeigen die gemessenen Sektionen in Peru dennoch klare Anhaltpunkte für globale ökologische Muster während der Mittleren Kreide. Die sehr ausgedehnte Beschaffenheit dieser Sektion erlaubt eine außerordentlich detaillierte zeitliche Auflösung von chemostratigraphischen Eigenschaften.

Acknowledgements

I would like to start this with the IAS meeting in Mendoza 2010, because this event changed my life forever. Here, I met for the first time to Prof. Dr. Adrian Immenhauser and the idea for a Peruvian Project was created during this meeting. Therefore, I would like to thank you Prof. Dr. Adrian Immenhauser for having believed in me since and for having given the opportunity to continue my geological career in Bochum. I appreciated your supervision and the time that you spend in clarifying my research with important constructive discussions and correcting my English. The successful of this thesis is largely due to your critical and justified comments. Thank so much also for having shared with me your knowledge in the field, as well.

Next, I would like to thank Prof. Dr. Stéphane Bodin who was always able at any time to answer my questions regarding to my thesis. This would not have been possible without your guidance and supervision, Maestro!. I would like to thank you so much for your time reading and implementing constructive comments to the manuscripts and correcting my Spanglish. Your critical discussion always helped me a lot with the success in publishing in peer-reviewed journal.

The team of the laboratory of Bochum are thanked for their assistance in the lab as well, conformed by Dr. Dieter Buhl, Dr. Andrea Niedermayr, Beate Gehnen, and Kathrin Schauerte. I would like also to thank Jan Danisch and Andreas Hollweg for their precious help in the lab.

Thanks so much for my co-author Prof. Dr. Ulrich Heimhofer for his help, critical and constructive discussions in my manuscripts.

In Peru, I would like to thank Dr. Jean Noel Martinez, Ing. Arturo Cordova, Dr. Victor Carlotto, Dr. Mirian Mamani and Dr. Aldo Alvan who supported my application for a DAAD scholarship. As well as my colleagues from the “Geología Regional” department of INGEMMET who helped me in the fieldwork in Peru, especially for Pedro Navarro, Wilson Gomez, Elvis Sanchez, Claudia Fabian and Elizabeth Ordoñez. To my friends from the Universidad Cajamarca are strongly thanked, as well.

Thank you our secretaries Cornelia Mell and Sabine Feige for making the German bureaucracy, a relaxing task.

I would like to thank the Deutsche Forschungsgemeinschaft (DFG, project n° BO- 365/2-1) for financing my project, as well as the Deutscher Akademischer Austauschdienst

(DAAD) for having given me the opportunity to come to Germany through a scholarship to continue my researcher career.

All my friends in the University Bochum who made a pleasant and funny stay in Germany, particularly for Jasper, Francois, Nico, Sylvia, Ann-Christine, Ana, Tillman, Jeremy, René, Robert, Ola, Sebastian, Sabine, Carla and Kevin. For those I forgot, I would like to thank you too.

I would like also to thank my friends of the foosball team SK Bochum 11 who helped me to integrate myself quickly into the German culture.

To conclude I would like to dedicate some words in Spanish to my parents and my wife.

Está presente tesis está dedicada a mi padre Felipe Navarro y a mi madre Martha Ramirez quienes a pesar de la distancia que nos separa siempre velan por mí y mi bienestar. Esto no hubiera sido posible si ustedes no fueran como son. Gracias por haberme motivado siempre a seguir con mis metas y ambiciones y por orar siempre por mí. A mis hermanos que siempre confían en mí también les estoy agradecido.

Gracias a la persona más importante de mi vida, mi esposa Vivien Rodriguez que hace que mi vida en cualquier rincón del planeta sea placentera y feliz. Mi amor gracias por apoyarme y ayudarme en hacer este trabajo difícil en algo más fácil, tú sabes que esto es nuestro y que ambos hemos trabajado mucho para llegar a esta meta. Eres mi soporte y mi motivación del día a día. En verdad nunca encuentro las palabras correctas para expresarte lo afortunado que soy al estar a tu lado. Te amo!!!

Table of contents

Abstract i

Kurzfassung iii

Acknowledgements v

Table of contents vii

CHAPTER 1

INTRODUTION 1

1 Rationale 1

2 The Albian–Turonian interval 4

2.1 Palaeogeography 6

2.2 Palaeoclimate 6

2.3 Palaeoceanography 7

2.4 Transgressive-regressive cycles and sea-level fluctuations 9

3 Oceanic anoxic events (OAEs) and periods of disturbed Carbon cycle 13

4 Tectonic of Peru 15

5 Open research questions and motivation of this study 17

6 Aims, scientific approach, and outline of this thesis 19

References 20

CHAPTER 2

METHODS AND MATERIALS 26

1 Field work and thin-section microscopy 26

2 Assessing the diagenetic alteration of shell material 26

2.1 Cathodoluminescence microscopy 26

2.2 Trace and major element analysis 27

3 Carbon isotope stratigraphy 27

13 3.1 Bulk micrite (δ CCarb) 27

13 3.2 Bulk organic matter (δ Corg) 27

4 Strontium-isotope analysis 28

vii

References 28

CHAPTER 3

RECORD OF ALBIAN TO EARLY CENOMANIAN ENVIRONMENTAL PERTURBATION IN THE EASTERN SUB-EQUATORIAL PACIFIC 30

1 Introduction 31

2 Regional tectonic and stratigraphic setting 33

3 Methods and materials 35

3.1 Field work and thin-section microscopy 35

3.2 Assessing the diagenetic alteration of shell material 36

3.2.1 Cathodoluminescence microscopy 36

3.2.2 Trace and major element analysis 36

3.3 Carbon isotope stratigraphy 37

13 3.3.1 Bulk micrite (δ CCarb) 37

13 3.3.2 Bulk organic matter (δ Corg) 37

3.4 Strontium-isotope analysis 37

4 Data description and interpretation 37

4.1 Facies associations and depositional environments 37

4.1.1 Shallow subtidal inner ramp setting 39

4.1.2 Open marine middle ramp setting 42

4.1.3 Outer ramp setting 42

4.2 Sequence stratigraphic interpretation 43

5 Chemostratigraphy 44

5.1 Radiogenic strontium isotope stratigraphy 44

5.2 Carbon-isotope stratigraphy 47

6 Discussion 48

6.1 Large-scale depositional setting of the Western Platform 48

6.2 A carbon isotope reference curve for the sub-equatorial eastern 49 Pacific

6.2.1 Bulk-micrite carbon isotope ratios 49

6.2.2 Bulk organic carbon isotope ratios 50

6.2.3 Comparison with other reference curves for Albian OAEs

viii

6.3 Impact of environmental changes on Peruvian neritic carbonate factory 54

7 Conclusions 56

Acknowledgements 57

References 57

CHAPTER 4

ONGOING CENOMANIAN – TURONIAN HETEROZOAN CARBONATE PRODUCTION IN THE NERITIC SETTINGS OF PERU 64

1 Introduction 65

2 Regional tectonic and stratigraphic setting 67

3 Methods and materials 68

3.1 Field work and thin-section microscopy 68

3.2 Carbon isotope stratigraphy 69

13 3.2.1 Bulk micrite data (δ CCarb) 69

13 3.2.2 Bulk organic matter data (δ Corg) 69

4 Data description and interpretation 69

4.1 Facies associations and depositional environments 69

4.1.1 Shallow subtidal inner ramp setting 69

4.1.2 Open marine middle ramp setting 71

4.1.3 Outer ramp setting 73

4.2 Sequence stratigraphic interpretation 74

4.3 Carbon-isotope stratigraphy 78

5 Discussion 79

5.1 Cenomanian–Turonian chemostratigraphy of Peru: Significance and comparison to other reference curves 79

5.1.1 Chemostratigraphic features of the Cenomanian isotope excursion and their relation to climate and sea-level change 82

5.1.2 The Cenomanian–Turonian boundary carbon isotope excursion 82

5.1.3 The Turonian post-excursion stage 86

5.2 Western Platform environmental patterns during OAE2 87

6 Conclusions 89

Acknowledgements 90

References 91

CHAPTER 5

RESPONSE OF WESTERN SOUTH AMERICAN EPEIRIC-NERITIC HETEROZOAN ECOSYSTEM TO OAE 1D AND 2

1 Introduction 99

2 Regional tectonic and stratigraphic setting 100

3 Methods and materials 102

3.1 Field work and thin-section microscopy 102

3.2 Carbon isotope stratigraphy 104

13 3.2.1 Bulk micrite data (δ CCarb) 104

13 3.2.2 Bulk organic matter data (δ Corg) 105

4 Data description and interpretation 105

4.1 Facies associations and depositional environments of the late Albian to early Cenomanian 105

4.2 Facies associations and depositional environments of the late Cenomanian to early Turonian 106

4.2.1 Shallow subtidal inner ramp setting 106

4.2.2 Open marine middle ramp setting 110

4.3 Sequence stratigraphic interpretation 110

4.3.1 The late Albian to early Cenomanian of the Jumasha Formation 111

4.3.2 The late Cenomanian to early Turonian of the Jumasha Formation 112

4.4 Carbon-isotope stratigraphy 116

4.4.1 Late Albian isotope excursion 116

4.4.2 Late Cenomanian isotope excursion 116

5 Discussion 117

5.1 Environmental implications of oyster and P. peruviana mass occurrences 117

5.2 Carbon cycle disturbances 119

5.2.1 Impact of environmental changes (OAEs 1d and 2) on the Peruvian neritic carbonate factory 121 x

5.2.2 OAEs 1d and 2: Significance and comparison to other reference curves 122

6 Conclusions 124

Acknowledgements 125

References 125

CHAPTER 6

SYNTHESIS 133

1 What was the impact of global mid-Cretaceous palaeoenvironmental perturbation on the neritic carbonate platforms in the sub-equatorial eastern Pacific setting of Peru? 133

2 Are the mid-Cretaceous palaeoceanographic patterns found in Peru comparable to those found in the proto-Atlantic and the Tethyan realm? 135

3 Did mid-Cretaceous perturbations of the carbon cycle leave a significant record in Peru or not, and if yes, how would this be expressed within the epeiric-neritic domain? 136

4 Specifically, is the OAE2 in Peru as scarcely represented and its organic content stratigraphically expanded (as opposed to being confined to a narrow interval) as suggested from e.g., Site 465, Hess Rise? 136

5 Outlook 137

References 139

ANNEX

GEOCHEMICAL DATA 142

CURRICULUM VITAE 149

xii

CHAPTER 1

INTRODUTION

1. Rationale

According to the Intergovernmental Panel on Climate Change report for policymakers released in 2013 (IPCC, 2013), climate change is a change in average weather conditions of the atmosphere at a certain place and time with reference to temperature, pressure, humidity, wind, and other key parameters (meteorological elements) lasting for an extended period of time (i.e., >30 years to Myrs). The causes for climate change respond to a complex set of factors such as solar radiation, biotic processes, volcanic eruptions, and plate tectonics. According to the National Academy of Sciences (NAS, 2010), since the Industrial Revolution, a change in the average weather conditions is taking place. These observations are based on multiple lines of research, documenting that the current climate undergoes "global warming" and that much of this warming is largely caused by human activity (NAS, 2010; IPCC, 2013).

In recent years, increasing attention has been paid to the impact of global warming, which is already having significant and harmful effects on ecosystems and society (Tan et al., 2015; Myers et al., 2015 and references therein). Indeed, the way the Earth system is reacting through a wide variety of climatic and oceanographic conditions points out a profound impact of human activity on the environment as a result of increased concentrations of carbon dioxide, methane, and nitrous oxide in the atmosphere and ocean over the last few centuries (IPCC, 2013). Over decades, observations based on the atmosphere, land, oceans, and cryosphere have been performed in order to decipher the causes and linkages of global warming, as well as how mitigate its impact (IPCC, 2013). This has led to reveal indicators of a sever global change as evidenced by, for instance, global average air and ocean temperatures, widespread melting of snow and ice, rising global average sea level, more common dangerous heat waves, increasing of extreme storm events and droughts in many areas (e.g., IPCC, 2013; Tan et al., 2015).

Fig. 1. Human activities in response to climatic forcing and natural processes in Watts per square metre (energy per unit area). Warming effect is represented by positive forcing. Source: NAS (2010) and references therein.

Fig. 2. Atmospheric CO2 curve (in ppm) indicating the monthly averaged data (red curve) with its respective averaged data (black curve), showing a clear upward trend. Source: NAS (2010) after Tans (2010). 2

Since the late 19th century, carbon dioxide (CO2) is derived from a wide scale burning of fossil fuels induced by human activities (IPCC, 2013). These CO2 emissions represent the largest single climate forcing agent, and together with other human activities, are playing a critical role in both the physical climate system and the Earth’s biosphere (NAS, 2010; Fig. 1).

Although roughly one-quarter of the CO2 released by human activities is absorbed in the sea, continuously increasing level of CO2 concentrations in the atmosphere at present rates could intensify the acidification of the world’s oceans (NAS, 2010). This may lead a reduction of the ocean ability to take up CO2, a positive feedback on global warming and, thus, disrupting many biological processes such as coral reefs (NAS, 2010).

Fig. 3. CO2 curve obtained through air bubbles trapped in an ice core from Law Dome in Antarctica during the last 1,000 years (in ppm), displaying a sharp rise in atmospheric CO2 at the beginning of the late 19th century. Source: NAS (2010) after Etheridge et al. (1996).

Through in situ CO2 measurements at many sites around the world, a 50-year well- calibrated, precise atmospheric CO2 curve (“the Keeling curve”) was elaborated (Tans, 2010;

Fig. 2). This curve evidences that there is a continuous increase of atmospheric CO2 by more than 20 percent since 1958 (Blasing, 2008; Tans, 2010). Studies performed in Greenland and

Antarctic ice sheets revealed that CO2 values were relatively constant for thousands of years before the Industrial Revolution, indicating that the current CO2 values are higher than they have been for at least 800,000 years (Etheridge et al., 1996; ASCC, 2010; Fig. 3).

Furthermore, measurements obtained of the isotopic abundances of the CO2 molecules in the

atmosphere, revealed a human origin of elevated CO2 product of large amounts of fossil fuel burning such as coal, oil, and natural gas (Keeling et al., 2005). Improving our understanding and estimates of current and projected future fluxes of CO2 is a key research need (NAS, 2010). Applying this argument geologist and palaeo-climate researchers exploit ancient climate archives. This is because these provide an alternative/complementary approach by using observations and theoretical models.

In this thesis, the mid-Cretaceous interval and more precisely, the Albian–Turonian interval (ca. 113–90 Ma; Ogg and Hinnov, 2012) has been studied in order to shed light on the complexity of climate change and global warming under greenhouse conditions in the southern hemisphere (Fig. 4). Mid-Cretaceous temperature gradients from the equator to the poles were reduced and polar ice caps seem absent at least over longer time intervals (e.g.,

Larson et al., 1993). Moreover, levels of atmospheric pCO2 were 2 to 6 times higher during the Cretaceous compared to pre-industrial values (e.g., Wallmann, 2001; Heimhofer et al., 2004; Haworth et al., 2005). All of this has been used to argue that the mid-Cretaceous represents a natural laboratory to study climatic extremes (Fig. 4). Having said this, any worker aiming at a contrast-comparison of mid-Cretaceous to future end 21st century future climate must acknowledge significant differences including rates, processes, and products.

2. The Albian–Turonian interval

The Albian (~113 to 100.5 Ma; Ogg and Hinnov, 2012) corresponds to the last stage of the Early Cretaceous (Fig. 4). It is preceded by the Aptian (126.3 to 113 Ma, Early Cretaceous) and followed by the Cenomanian (100.5 to 93.9 Ma) and Turonian (93.9 to 89.8 Ma) of the Late Cretaceous. The Albian Stage has been defined in northeast France (e.g., Ogg and Hinnov, 2012). Despite a significant bulk of published data, the exact position of the Aptian/Albian boundary is still undecided (Ogg and Hinnov, 2012). Agreement seems to exist, however, that the base of the Albian Stage should be assigned to the lowest occurrence of the calcareous nannofossil Praediscosphaera columnata, near the “Kilian” carbon isotope excursion at the base of the Microhedbergella rischi foraminifer zone (e.g., Selby et al., 2009; Huber et al., 2011; Ogg and Hinnov, 2012; Fig. 4). The Cenomanian stage has been defined in northern France as well and its base is defined by the lowest occurrence of the planktonic foraminifer Rotalipora globotruncanoides, being slightly lower than the lowest occurrence of the Cenomanian ammonoid marker of Mantelliceras mantelli (Ogg and Hinnov, 2012 and references therein; Fig. 4). The official global stratotype section and point (GSSP) for the Cenomanian was chosen in the Mont Risou section in southeast France (Kennedy et al.,

2004). The Turonian stage has been constantly redefined and extensively studied in France (e.g., Bengtson et al., 1996). Its base is suggested by the lowest occurrence of the ammonite Watinoceras devonense, near the global oceanic anoxic event 2 (Bengtson et al., 1996). Its GSSP is situated at Rock Canyon Anticline, east of Pueblo (Colorado, west-central USA) (Bengtson et al., 1996; Kennedy and Cobban, 1991; Ogg and Hinnov, 2012).

Fig. 4. Mid-Cretaceous numerical time scale with biostratigraphic zonations, carbon isotopes and position of sea level. Source: Ogg and Hinnov (2012).

2.1 Palaeogeography

During the Early Cretaceous, a new palaeogeographic configuration is established (Moulin et al., 2010; Fig. 5). This was marked by the disintegration of Gondwana and Laurentia, themselves further subdivided into smaller plate tectonic units, resulting in the development of the North Atlantic Ocean (Labrador Sea), between North America and Greenland (e.g., Austin and Howie, 1973). In the Aptian, a breakup occurred separating the Indian and Australian continents that formed the North and SW Indian Ocean (e.g., Sahabi, 1993) and initiated the separation between Africa and South America (e.g., Aslanian et al., 2009; Moulin et al., 2010). Thereafter, in the Aptian–Albian transition, the breakup of Africa and South America ended (e.g., Moulin et al., 2010). From this time towards the Turonian, the opening of the Equatorial Atlantic Ocean took place, leading to a connection between North and South Atlantic Ocean water masses (Eagles, 2006; Fig. 5).

Fig.5. Palaeogeographic map of South America during the mid-Cretaceous (modified after Blakey, 2011) indicating the position of what today are the sedimentary rocks formerly deposited on the Peruvian platform (yellow arrow).

2.2 Palaeoclimate

During the mid-Cretaceous (Albian to Turonian), the temperatures reached a maximum, declining slightly to a more accentuated cooling until the top of this period (Friedrich et al.,

2012). The warmest temperatures in the mid-Cretaceous are likely to be associated to a critical time that encompasses the final disintegration of Pangea, resulting in high volcanic activity and sea-floor spreading (Chumakov et al., 1995; Hay, 2008; Hay and Floegel, 2012; Bodin et al., 2015; Figs. 6A–B). This triggered a first-order sea-level highstand, as well as anomalously large inputs of CO2 into the atmosphere during the formation of Large Igneous Provinces (LIPs). This caused hothouse episodes as indicated by the formation oceanic anoxic events such as the early Aptian OAE1a (120 Ma) and the Cenomanian–Turonian OAE 2 (93 Ma) (Kidder and Worsley, 2012; Fig. 4).

At sites localized in high latitudes (60°–90°; Figs. 6A–B), the climate was temperate and humid, and sensitive to climatic variations (Hay and Flögel, 2012), indicated by terrestrial fauna and flora (Chumakov et al, 1995). Transient episodes of short cooling events were likely to have occurred here (Chumakov et al, 1995; Jarvis et al., 2011; Bodin et al., 2015). Relatively warm climate in this belt is supported by the occurrence of thermophilic insects and terrestrial fauna, which temperatures roughly fluctuating in 2–5°C for the coldest months and 5–15°C for the warmest months (Chumakov et al, 1995).

Across mid-latitudes of both hemispheres (30°–60°; Figs. 6A–B), warm and humid belts established during the mid-Cretaceous as evidenced by flora, terrestrial fauna and widespread occurrence of laterite, kaolin, bauxite, and coal deposits (Chumakov et al., 1995; Hay and Flögel, 2012). Moreover, an annual temperature approaching 20°C in southern and 13–20°C in northern mid-latitudes and equable climate in all epicontinental seas and land regions is commonly suggested (Chumakov et al., 1995; Hay, 2008; Hay and Flögel, 2012). This was perhaps regulated by warm trade-wind coming from the Pacific and eastern Tethys (Chumakov et al., 1995).

In the tropical-equatorial regions (Figs. 6A–B), the presence of a tropical hot arid belt has been suggested by several previous workers. Arid conditions were typical for the Aptian– Albian transition, represented by widespread evaporite deposits associated to red beds (Chumakov et al., 1995; Hay and Flögel, 2012). Temperatures in the atmosphere may have been as warm as 42°C (Bice and Norris, 2002; Bice et al., 2006), resulting in a dead zone effect (Hay and Flögel, 2012). These extreme temperatures may have formed an over-regional life-limiting parameter (Hay and Flögel, 2012). The Intertropical Convergence Zone (ITCZ) took place in a unified tropical-equatorial belt, separating the northern and southern hemispheres (Chumakov et al., 1995; Hay and Flögel, 2012).

2.3 Palaeoceanography

Fig. 6A. Albian and; (B) Cenomanian climate indicators and zones. Source: Hay and Floegel (2012) after Chumakov et al. (1995).

As pointed out by Hay (2008), the Cretaceous ocean circulation may have been very different from that of today, because it was characterized by episodes of local anoxia occurring during the earlier Cretaceous, and regional to global ones during the mid- Cretaceous. It was also suggested an ocean circulation that could carry warmth to the polar regions and invasion of calcareous oceanic plankton into epicontinental seas, responding to sea-level rise during the Early to the Late Cretaceous (Hay, 2008; Figs. 6A–B). This resulted in widespread oligotrophic stratified oceanic water onto the shelves (Gale et al., 2000). Investigations performed at several sites in the Tethys and North Atlantic points at two different ocean circulation modes for the Albian – Cenomanian interval (Giorgioni et al., 2015). One is indicated by less connection and unstable circulation for the early up to early late Albian, whereas more stable circulation with better connected basins and better stratified water masses in the late Albian to early Cenomanian (Giorgioni et al., 2015). For the Cenomanian – Turonian interval (Trabucho-Alexandre et al., 2010), this was characterized by the connection between the Atlantic and Pacific oceans by currents at the surface, reversed circulation at intermediate depths and disconnected circulation at deeper depths (Fig. 7). Water masses from the Pacific were likely moving eastward via the equatorial Central American Seaway into the Atlantic Basin. This modelled circulation pattern represents an estuarine circulation and accounts for the upwelling zones and the OAE black shales depositions described from the middle Cretaceous world (Trabucho-Alexandre et al., 2010).

2.4 Transgressive-regressive cycles and sea-level fluctuations

The revised Cretaceous curves for both long-term and short-term sea-level variations published in Haq (2014), suggests higher average sea levels than the present day mean sea level (PDMSL; Fig. 8). Haq (2014) suggested that the Cretaceous sea level reached a trough in mid-Valanginian (~75 m above PDMSL), followed by a first maximum point in the early Barremian (~160–170 m above PDMSL) and by a highest point in earliest Turonian (~240– 250 m above PDMSL; Fig. 8). Evidence for this comes from western European basins, the Russian and the Arabian platforms and further supporting data from the US and the Pacific and is constrained with oxygen-isotopic trends. Specifically for the mid-Cretaceous, the Albian stage comprises 8 “eustatic” sea-level events labelled as KAl1–8, classifying almost all as medium in magnitude (between 25 and 75 m). Whereas the Cenomanian is characterized by five “eustatic” sea-level events labelled as KCe1–5 (Haq, 2014). Finally, the Turonian is characterized by five “eustatic” sea-level events labelled as KTu1–5 (Haq, 2014).

Fig. 7. Model circulations favouring black shale deposition during the mid-Cretaceous OAEs in the North Atlantic: (a) At the surface, (b) at 1 km depth and (c) at 4 km depth. Source: Trabucho-Alexandre et al. (2010).

∼ ∼ 10

Fig. 8. The Cretaceous eustatic cycle chart, showing the maximum long-term sea-level position during the mid- Cretaceous (Albian, Cenomanian and Turonian stages). Source: Haq (2014). 11

Fig. 9. Carbonate carbon isotope composite age-calibrated curve (Gradstein et al., 2012) and major mid- Cretaceous paleoceanographic events (after Herrle et al., 2015).

3. Oceanic anoxic events (OAEs) and periods of disturbed Carbon cycle

The Cretaceous comprises several oceanic anoxic events (OAEs; Jenkyns, 2010; Fig. 9). The term “oceanic anoxic event” was first introduced by Schlanger and Jenkyns (1976) as a geological time represented by an expanded and intensified oxygen minimum zone in the world oceans. OAEs coincide well with carbon isotope ratio (δ13C) high amplitude fluctuations; therefore, it was suspected that these OAEs were related to carbon-cycle disturbances (Schlanger and Jenkyns, 1976; Jenkyns, 2010; Bodin et al., 2015). OAEs are associated to enhanced accumulation of black shale deposition in basinal and deeper settings, anomalously high burial rates of marine organic carbon on a global scale and significant ecological changes in many faunal and floral groups (Schlanger and Jenkyns, 1976; Jenkyns, 2010; Lamolda et al., 1997; Föllmi et al., 2006; Kaiho et al., 2014; Bodin et al., 2015). In the NW Tethyan Platform, OAEs have been associated to complex alternations between growth and demise of carbonate systems, linked themselves by trophic level changes (e.g., Föllmi et al., 2006; Immenhauser et al., 2005; Bodin et al., 2006; Andrieu et al., 2015). Changes in trophic levels is evidenced with the switch from oligo- to meso- to eutrophic conditions, causing changes in the geometry, depositional, and carbonate-secreting organisms shaping the evolution of the northern Tethyan platform (e.g., Föllmi et al., 2006; Immenhauser et al., 2005; Bodin et al., 2006; Andrieu et al., 2015). Despite a significant bulk of published data, the underlying controls of OAEs are still debated, but during the two decades a linkage between massive volcanism and OAEs has been suggested (e.g., Larson, 1991; Mort et al., 2008; Du Vivier et al., 2014; Bodin et al., 2015).

The well-known OAEs of global significance for the Cretaceous (e.g., Jenkyns, 2010) are represented by the OAE1a of the early Aptian (Selli event, 120 Ma), the OAE 1b set of the early Albian ( 111 Ma) and the OAE2 of the Cenomanian–∼Turonian (Bonarelli event, 93 Ma). Amongst ∼these, other OAEs have been recognized in the Tethyan domain such as∼ the Albian OAE1c and the OAE1d, as well as OAE3 of the Coniacian–Santonian in the Atlantic, but their distribution is somewhat problematic due to patchy records (see Jenkyns, 2010). In the following account, Albian OAEs and OAE2 are here discussed with more detail as they are of great significance in the present study (Fig. 9).

OAE1b was a long-lasting event with duration of about 6.3 Myr, characterized by elevated surface water fertility and represented by negative excursions of carbon isotope records (δ13C) with amplitudes of 1.5 to 2‰ (Premoli Silva et al., 1989; Herrle et al., 2003, 2004; Reichelt, 2005; Browning and Watkins, 2008). OAE1b was linked to a global cooling and in 13

consequence to a sea-level fall (Weissert and Lini, 1991). In the Vocontian Basin (Western Tethys), OAE1b is represented by four black shale levels, which are the uppermost Aptian Jacob Level and the lower Albian Kilian, Paquier and Leenhardt levels (Reichelt, 2005). The most expanded and best studied of these sub-events are the Kilian and Paquier levels (Fig. 4; Trabucho-Alexandre et al., 2011).

Fig. 10. The mid‐Cretaceous estuarine circulation for the North Atlantic, suggesting accumulation of organic-rich black shale accentuated in deeper-open marine environments near to the low latitude North Atlantic and Tethys domain. Source: Trabucho-Alexandre et al. (2010).

OAE1d is stratigraphically localized at the Albian–Cenomanian boundary (Jarvis et al., 2006). It is ascribed as the Niveau Breistroffer in the Vocontian Basin (France, Bornemann et al., 2005) and the Pialli Level in the Umbria-Marche Basin (Italy, Coccioni, 2001; Fig. 9). In the Tethys Ocean, OAE1d is characterized by the deposition of rhythmic organic-rich black

13 shales exhibiting enhanced contents of bulk organic matter (δ Corg) and carbon positive 14

excursion of ~ 1‰ (Gale et al., 2011). This event is of widespread distribution being reported in other settings such as the Atlantic and Pacific Ocean (see Gambacorta et al., 2015 and references therein).

OAE2 represents one of the most extreme carbon cycle perturbations of the Phanerozoic and is characterized by widespread organic-rich black shale deposition and a positive

13 13 excursion of 2–3‰ in δ Ccarb and 3–4‰ in δ Corg spanning ~600kyr. This pattern is ubiquitously recorded in marine carbonate, and both marine and terrestrial organic matter (e.g., Schlanger and Jenkyns, 1976; Sageman et al., 2006; Jarvis et al., 2006; Elrick et al., 2009; Gertsch et al., 2010; Meyers et al., 2012; Ma et al., 2014). OAE2 was preceded by the Mid-Cenomanian Event I (MCEI; Andrieu et al., 2015). The MCEI is defined by two positive carbon-isotope excursions (MCE Ia, MCE Ib; Mitchell et al., 1996). Contrary to OAE2, MCEI is not characterised by organic-rich black shale deposition, but is recorded in Tethyan and North Atlantic settings as a major perturbation in carbonate platform productivity (Giraud et al., 2013).

4. Tectonic of Peru

The western margin of South America was largely influenced by the late stages of Gondwana breakup that culminated in the separation of Africa and South America during the Aptian–Albian transition and the opening of the Equatorial Atlantic Ocean (Moulin et al., 2010). These events caused continental arc volcanism (Lower Cretaceous), plutonism (Late Cretaceous), and subduction along the western margin of South America (Soler and Bonhomme, 1990). Along the western margin, an epeiric-neritic ramp was installed, characterized by subsequent mixed carbonate-siliciclastic deposition and controlled by NNW–SSE-trending structures (Atherton and Webb, 1989; Jaillard, 1987; Soler and Bonhomme, 1990; Pindell and Tabbutt, 1995; Robert and Bulot, 2004). The western structures are represented by the Paracas structural high and by the potential barrier effects of the Albian volcanic arc, which perhaps locally protected the ramp from impinging Cretaceous Pacific swell (Jaillard, 1987). Towards the southeast, deposition was attached to the Marañon Massif that controlled transient intervals of important source area of continental runoff of the exposed continental basement in present-day Brazil (Jaillard, 1986; Figs. 11 and 12).

Fig. 11. Mid-Cretaceous palaeogeographic map of Peru (based on Pindell and Tabbutt, 1995; Robert and Bulot, 2004). Black squares indicates study area of Cajamarca region (Northern Peru) and Oyon region (Central Peru). Note size of study area relative to the vast dimensions of the Western Platform. A–A´ denotes position of transect in Fig. 12.

Fig. 12. Hypothetical transect across the Andean Basin during Neocomian–Coniacian times (based on Benavides- Caceres, 1956; Atherton and Webb, 1989; Jaillard, 1987) with inferred Western Platform ramp morphology. Refer to approximate position of transect in Fig. 11 (A–A´).

The northern and central regions of Peru have previously been studied by a limited number of authors with rather traditional biostratigraphic and sedimentologic concepts in mind (e.g., Benavides-Caceres, 1956; Jaillard, 1986, 1987; Jaillard and Arnaud-Vanneau, 1993; Robert et al., 2009 and references therein). Combining evidence from the rather scarce previously published data set on the Albian to Turonian pictures a mainly north-south trending, exceptionally large outcrop belt with evidence for deepening to west, i.e. the Palaeo- Pacific. During the mid-Cretaceous in Peru, two transgressive pulses led the deposition of two large-scale sequences, which comprise eight carbonate formations. These carbonate formations in the Cajamarca region (northern Peru), are represented by the Inca, Chulec, Pariatambo, and Yumagual formations of the early Albian–early Cenomanian (Fig. 12), overlain by the Mujarrún, Romirón, Coñor, and Cajamarca formations of the early Cenomanian–middle Turonian (Fig. 12). In the Oyon region (Central Peru; Fig. 11), the mid- Cretaceous is represented by the Pariahuanca, Chulec, and Pariatambo formations (Benavides-Caceres, 1956). These are overlain by the Jumasha Formation of the late middle Albian– early Turonian.

5. Open research questions and motivation of this study

The number of modern, process-oriented studies dealing with mid-Cretaceous shallow- marine carbonate deposits in South America is limited and even more so with reference to

Peru. Indeed, to date, much of the present knowledge of oceanic anoxia in the mid- Cretaceous world is biased towards data derived from hemipelagic and pelagic sections in Europe (Tethys) and North America (Western Interior Basin) as well as data from various ocean drilling projects (e.g., Paul et al., 1999; Price, 2003; Robinson et al., 2004; Dumitrescu et al., 2006; Sageman et al., 2006; Ando et al., 2008; Voigt et al., 2008; Jarvis et al., 2011). Less attention has been paid to the impact of oceanic anoxic events on shallow water carbonate systems. These arguably bear crucial information to understand OAEs causes and consequences, as indicated in the NW Tethyan Platform, where OAEs have been linked to trophic level changes (e.g., Föllmi et al., 2006; Immenhauser et al., 2005; Bodin et al., 2006). Evidence for the sedimentological and geochemical signatures of various anoxic events in Peru was first explored by Jaillard and Arnaud-Vanneau (1993) for the OAE2 interval. These authors performed their work mainly from the viewpoint of biostratigraphy. The absence of bituminous shales in Peru led these authors to conclude that anoxia was not an important environmental feature of these epeiric water masses. To present, no modern geochemical and palaeoceanographic study have been undertaken to verify this claim.

As a consequence, the timing of pre-OAE ecosystems, the syn-OAE ecosystems and those during the recovery as well as the potential collapse modalities in the vast epeiric shelf of Peru represent a hitherto underrepresented archive of Albian–Turonian OAEs. The absence of black, organic-rich deposits in these sections implies that the corresponding biotic and environmental effects differ significantly from those of, for example, hemipelagic and pelagic sections. Moreover, the significant continentality of this vast epeiric platform, nutrient runoff, local carbon cycles superimposed on provincial and global ones and differences in the aquafacies render the direct comparison of these settings with open oceanic ones challenging (Thomas et al., 2004; Immenhauser et al., 2008). Nevertheless, based on the enormous dimensions of the western South American carbonate shelf, its complex subdivision in ramps, basins, and structural highs and the setting with a large oceanic basin to the west and a very large uplifted continental basement area to the east makes this region a fascinating natural laboratory. All of this forms a strong motivation for the present research project that is guided by the following main research questions:

• What is the impact of global mid-Cretaceous palaeoenvironmental perturbation on the neritic carbonate platforms in the equatorial eastern Pacific setting of Peru? • Are the mid-Cretaceous palaeoceanographic patterns found in Peru comparable to those found in the Atlantic and the Tethyan ones? • Did mid-Cretaceous perturbations of the carbon cycle leave a significant record in Peru or not, and if yes, how is this event expressed within the epeiric-neritic domain?

• Specifically, is the OAE2 scarcely represented and organic content stratigraphically expanded (as opposed to being confined to a narrow) interval as suggested so from e.g., Site 465, Hess Rise?

In essence, the research proposed here was driven by the following working hypotheses:

• The mid-Cretaceous OAEs (OAE 1b, 1d and 2) as recorded in the Tethyan and Atlantic realm are part of a series of global events affecting the Cretaceous ocean-atmosphere system and biogeochemical cycling and are thus recorded in Peru, representing the eastern Pacific domain. The scarcity of previous records is merely the result of incomplete sections. • Peru lies at the cross-roads between the Tethys and the Pacific Oceans and well- constrained data sets from this part of the world form a “data bridge” between those two realms.

6. Aims, scientific approach, and outline of this thesis

In order to answer the above stated questions and validate the hypotheses, we made use of the excellent outcrop conditions of the hitherto poorly studied mid-Cretaceous (Albian– Turonian) sections located in the northern and central Andes of Peru (Fig. 11). This study is based on a multidisciplinary approach making used of a broad, multi-proxy approach (C, and Sr isotopes, sedimentology, sequence stratigraphy etc.). Geochemical analyses are coupled with a detailed sedimentological and palaeo-ecological assessment of carbonate platform successions. Carbonate platform ecology is examined in detail to trace changes in the carbonate depositional mode and biotic and abiotic changes in marine environment (e.g., Föllmi et al., 2006; Immenhauser et al., 2005; Bodin et al., 2006). The results are compared and discussed to coeval findings from the Atlantic and Tethyan realm in order to understand causal linkages among global geological patterns during the mid-Cretaceous greenhouse word. The thesis is subdivided into six chapters:

Chapter 1 deals with basic concepts and terminologies used in this manuscript. This chapter discusses the state-of-the-art of the Albian–Turonian (mid-Cretaceous) interval and provides the fundament to understand the scientific discussion (e.g., oceanic anoxic events, palaeoclimate changes, carbon cycle disturbances etc.). This also documents the regional tectonic setting of the present thesis.

Chapter 2 documents the methods and materials used to perform this study.

Chapter 3 presents and discusses a new early Albian–early Cenomanian carbon isotope

13 13 (δ Ccarb and δ Corg) curve, based on heterozoan epeiric-neritic successions from Peru. Chemostratigraphic evidence for the expression of OAEs 1b, 1c, and 1d is presented, revealing an overall good agreement with published sections from the Pacific, proto-Atlantic, and the Tethys domains. In order to achieve this objective, the δ13C records were supported by biostratigraphic evidence and 87Sr/86Sr isotope stratigraphy using well-preserved oyster shells. A special focus was the complex interplay of climatic changes, nutrient supply and platform drowning associated to the impact of the OAE1b set (Kilian, Paquier and Leenhardt Levels), OAE1c and OAE1d.

Chapter 4 documents the ongoing heterozoan carbonate production in the neritic settings of Peru during the early Cenomanian–middle Turonian. Moreover, this chapter provides a new

13 13 early Cenomanian–middle Turonian carbon isotope (δ Ccarb and δ Corg) curve that matches well with global published high-resolution data from the English Chalk and the Portland # 1 core. This led to the following conclusions regarding the ecological and environmental structure of the western platform in Peru: (i) An Albian to early late Cenomanian heterozoan ramp recording the middle Cenomanian event I (MCEI); (ii) a late Cenomanian OAE2 interval, characterized at the base by a progressive deepening leading to the short-lived establishment of middle ramp type sedimentation. Followed by (iii) benthonic inner ramp sedimentation during a sea-level highstand in the δ13C plateau, and by a recovery of δ13C values at the end of OAE2 associated to reduced carbonate production. It is shown that the ramp system did not suffer any carbonate crisis during OAE2.

Chapter 5 characterizes the main contribution of environmental interaction between tectonic events, sea-level change, and climatic patterns during the late Albian–early Cenomanian and late Cenomanian–Turonian. In order to deal with these three different parameters, the main changes in term of carbonate factory and/or clastic inputs recorded were compared. Special focus was on the environmental implication of OAEs 1d and 2 on the neritic carbonates of the western South American domain in Peru.

Chapter 6 provides an overview of the main scientific outcome of this thesis. Suggestions for future work are given.

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CHAPTER 2

METHODS AND MATERIALS

1. Field work and thin-section microscopy

Well-exposed sections from the northern and central Andes in Peru (Cajamarca and Oyon regions) were chosen for their well-exposed strata of the early Albian–Turonian age. In total, ca. 1600 m of section have been logged. These field data, obtained rock samples and thin-sections provide the fundament for the petrographic and facies interpretations presented here. In analogy with the nomenclature and approach presented in Immenhauser et al. (2004), non-skeletal and skeletal components were analysed semi-quantitatively with numbers indicating their relative abundance: 0 = absent, 1 = present, 2 = frequent, 3 = abundant, 4 = dominant.). Oyster shells were collected as potential archive material for radiogenic strontium-isotope stratigraphy. Limestone classification is according to Dunham (1962). Discontinuity surfaces, mainly marine hardgrounds, were recorded stratigraphically and traced laterally between the two sections. Combined with under- and overlying facies, discontinuities acted as pining points for a first attempt to understand the sequence stratigraphic evolution of the study area. Emphasis must be laid on the fact that sequence stratigraphy as based on a limited number of sections is at best regional in nature.

2. Assessing the diagenetic alteration of shell material

2.1 Cathodoluminescence microscopy

Cathodoluminescence (CL) microscopy examination of selected outer shell layers of five oyster specimens were investigated for their CL pattern in order to avoid diagenetically altered shell material. This was carried out with the ‘hot cathode’ CL microscope (type HC1-LM) at the Ruhr-University Bochum, Germany. The acceleration voltage of the electron beam is 14 kV and the beam current is set to a level gaining a current density of ~9 μA mm-2 on the sample surface. Refer to Christ et al. (2012) for details on the analytical procedure.

2.2 Trace and major element analysis

Oyster shells that passed initial cathodoluminescence inspection were subsequently analysed for major and trace elemental composition at the Ruhr-University Bochum, Germany. Aliquots of the powder samples (1.35–1.65 mg) were analysed by inductively coupled plasma atomic emission spectroscopy (ICP-AES) for strontium, magnesium, iron and manganese. Refer to Huck et al. (2011) for details on the analytical procedure.

3. Carbon isotope stratigraphy

13 3.1 Bulk micrite (δ CCarb)

Carbonate bulk rock specimens were sampled by use of tungsten drill bits. While drilling the carbonate powder, areas rich in sparry cement, large bioclasts and diagenetic calcite vein material were avoided. In order to reveal the variability of carbon-isotope composition within a single rock sample, several sub-samples were drilled from 20% of all the hand specimens collected. Refer to Immenhauser et al. (2005) for more details of the analytical procedure. Carbon-isotope analysis of 123 micrite samples was performed using a Thermo Finnigan MAT delta-S mass spectrometer at the isotope laboratory of the Ruhr-University Bochum, Germany (see Appendix data). Repeated analyses of certified carbonate standards (NBS 19, IAEA CO–1 and CO–8) and internal standards show an 13 13 external reproducibility of ≤0.06‰ for δ Ccarb. The δ Ccarb values are expressed on a per mil (‰) basis relative to the Vienna-Pee Dee Belemnite standard (V-PDB).

13 3.2 Bulk organic matter (δ Corg)

Carbon-isotope analyses were performed on bulk organic matter samples (see Appendix data). Two grams of sample powder were placed in 50mL centrifuge tubes and 6N HCl was added, this procedure was repeated until no carbonate reaction was visible. The samples are then centrifuged and the supernatant removed, the residues were rinsed with distilled water via centrifuge and dried at 40 °C. The δ13Corg measurements were performed with an elemental analyser (CE 1110) connected online to ThermoFinnigan Delta V Plus mass spectrometer. All carbon isotope values are reported in the conventional δ‐notation in permil relative to V-PDB (Vienna‐PDB). Accuracy and reproducibility of the analyses was checked by replicate analyses of international or laboratory standards. Organic

carbon analyses were conducted at the stable isotope laboratory of the Friedrich-Alexander University of Erlangen-Nuremberg and at the stable isotope laboratory of the Leibniz University Hannover, Germany.

4 Strontium-isotope analysis

Sample powders of low–Mg calcite shells of oysters were taken by use of tungsten drill bits. The required sample quantity of 500 ng was determined as based on the strontium content of each sample. The Sr fraction was separated twice using quartz glass columns filled with BioRad ion exchange resin AG50W-X8 according to column calibration pattern. The 87Sr/86Sr ratio was determined with a 7 collector Thermal Ionisation Mass spectrometer (TIMS) MAT 262 at the Ruhr-University Bochum, Germany, in 3 collector dynamic mode. As standard reference materials NIST NBS 987 and USGS EN–1 were chosen. The total blank for Sr isotope analysis, including chemical separation and loading blank, is 1.5 ng. The repeatability was tested with the USGS EN–1 modern bivalve carbonate, which passed through the same procedures as all samples. The average 87Sr/86Sr value is 0.709160 ± 0.000027 2σ (n=209). The reproducibility of Sr measurements represented by NBS987, which was directly loaded onto a Re filament, is 0.710240 ± 0.00034 2σ (n=233). Refer to Huck et al. (2011) for more details of the analytical procedure.

References

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Huck, S., Heimhofer, U., Rameil, N., Bodin, S., Immenhauser, A., 2011. Strontium and carbon-isotope chronostratigraphy of Barremian–Aptian shoal-water carbonates: Northern Tethyan platform drowning predates OAE 1a. Earth and Planetary Science Letters 304, 547–558.

Immenhauser, A., Hillgärtner, H., Sattler, U., Bertotti, G., Schoepfer, P., Homewood, P., Vahrenkamp, V., Steuber, T., Masse, J.P., Droste, H., Taal-van Koppen, J., Van der Kooij, B., Van Bentum, E., Verwer, K., Hoogerduijn-Strating, E., Swinkels, W., Peters, J., Immenhauser-Potthast, I., Al Maskery, S., 2004. Barremian–lower Aptian Qishn Formation, Haushi-Huqf area, Oman: a new outcrop analogue for the Kharaib/ Shu'aiba reservoirs. GeoArabia 9, 153–194.

CHAPTER 3

RECORD OF ALBIAN TO EARLY CENOMANIAN ENVIRONMENTAL PERTURBATION IN THE EASTERN SUB- EQUATORIAL PACIFIC

Navarro-Ramirez J.P., Bodin S., Heimhofer U., Immenhauser A.

ABSTRACT

The present paper documents and discusses a new Albian–early Cenomanian carbon isotope

13 13 (δ Ccarb and δ Corg) curve from the subequatorial Eastern Pacific in Peru. Chemostratigraphic evidences for the expression of the OAE1b set and for OAE1c and OAE1d are presented. This dataset is relevant inasmuch as previous work is strongly biased towards study sites in North America (Western Interior Basin), in Europe (Tethys) and the Pacific realm. A comparison of the carbon isotope stratigraphy obtained in Peru with published sections from the Central and Western Pacific, the Western Atlantic and Northern and Western Tethys reveals an overall good agreement supporting the global nature of the isotope patterns described here. The δ13C from Peru record is constrained by biostratigraphic evidence and 87Sr/86Sr isotope stratigraphy using well-preserved oyster shells. Furthermore, we document the development of a heterozoan epeiric-neritic mixed carbonate-siliciclastic ramp in the Western Platform of Peru and its corresponding sedimentary facies associations. This dataset was used to elucidate the complex interplay of climatic changes, nutrient supply, and platform drowning, leading to the following conclusions: (i) An upper Aptian–lower Albian major change from siliciclastic-dominated to carbonate sedimentation coincided with the impact of the Kilian Level. (ii) A lower Albian incipient platform drowning linked to the impact of the Paquier Level. (iii) A lower middle Albian major demise of neritic carbonate production that coincides with the Leenhardt Level, followed by middle Albian condensed sedimentation that reports

13 prominent negative values in δ Ccarb prior to the onset of OAE1c. (iv) Finally, renewed carbonate ramp production during the upper Albian–lower Cenomanian. The data shown here represent the foundation for future work documenting the mid–Cretaceous of Peru and its implications for the palaeoceanography of the SE subequatorial Pacific.

1. Introduction

The Albian–early Cenomanian (mid-Cretaceous, 113–96.5 Ma; Ogg and Hinnov, 2012) witnessed a series of oceanic anoxic events (OAEs, ∼Leckie et al., 2002; Heimhofer et al., 2004; Jarvis et al., 2006; Sprovieri et al., 2013; Lorenzen et al., 2013). One of the most prominent features related to these events is represented by the accumulation of organic matter in marine sediments (Schlanger and Jenkyns, 1976). These organic-rich sediments (black shales) can be traced in ocean basins worldwide as individual, distinct horizons or as bundles with several layers enriched in organic carbon (Heimhofer et al., 2006; Emeis and Weissert, 2009; Trabucho-Alexandre et al., 2012). This specific facies and related carbon isotope anomalies represent short-lived perturbations of the global carbon cycle (Jenkyns, 2010; Melinte-Dobrinescu and Bojar, 2010; Trabucho-Alexandre et al., 2010; Bodin et al., 2013). OAEs are also associated with minor extinction and rapid turnover phases in marine life, changes in carbonate platform ecology and phases of platform drowning (Hallock and Schlager, 1986; Wilmsen, 2000; Föllmi, 2012; Krencker et al., 2014). Specifically, during the mid–Cretaceous (Gale et al., 2011), OAEs took place (i) during the late Aptian–early Albian (OAE1b, 114–109 Ma), (ii) during the early late Albian (OAE1c, 102 Ma) and (iii) at the

Albian–Cenomanian∼ boundary (OAE1d, 99.5 Ma). ∼

Much of the present understanding∼ of mid-Cretaceous climate and its perturbations

13 comes from the carbon-isotope proxy making use of bulk micrite (δ Ccarb) and bulk organic

13 matter (δ Corg) sample sets. Obtained chemostratigraphic sections from numerous basins worldwide display excursions in the carbon isotope record that have significance as time or event markers and shed light on processes involved (e.g. Erbacher et al., 1996; Bralower et al., 1999; Herrle et al., 2004, Jarvis et al., 2011; Krencker et al., 2014). Amongst these, OAE1b was a long-lived event lasting about 6.3 Myr, characterized by a bundle of up to four black shale levels and associated perturbations in the carbon cycle recorded as excursions in carbon isotopes (Herrle et al., 2004, 2003; Reichelt, 2005; Madhavaraju et al., 2013). In the Vocontian Basin (Western Tethys), these levels are represented by the uppermost Aptian Jacob Level and the lower Albian Kilian, Paquier and Leenhardt levels. Here, the Kilian and Paquier levels are the most widespread, deposited during eustatic highstands (Bréhéret, 1994) and composed by marine organic matter with a low terrigenous input (Trabucho- Alexandre et al., 2011).

On the other hand, in the Tethys realm, the early late Albian OAE1c is a black shale level characterized by abundant terrigenous organic matter (Amadeus Level; Coccioni et al., 1993; Galeotti, 1998; Luciani et al., 2007). The chemostratigraphic pattern of this event is not uniform, a feature most likely related to differences in sedimentation rates, the interaction of 31

local and global processes and other, so far not well constrained factors. In this sense, only a

13 weakly developed negative δ Ccarb shift has been reported in Central Europe for OAE1c

13 (Erbacher et al., 1996), whereas a prominent negative excursion in δ Corg has been identified in Mexico (Bralower et al., 1999) or in sections on Japan (Takashima et al., 2010). At the Albian–Cenomanian transition, OAE1d has been ascribed to a black shale level known as the

13 Breistroffer Level in France (Gale et al., 1996). A long-lasting positive excursion of δ Ccarb within this interval has led several authors to suggest globally significant organic-carbon burial (Nederbragt et al., 2001; Schröder-Adams et al., 2012; Scott et al., 2013). It is likely, however, that different OAEs have different driving mechanisms and differential organic matter-rich sedimentation in different localities during an event as reflected in different types of organic matter found in specific black shale intervals (Kuypers et al., 2001).

Much of the present knowledge of the Albian–early Cenomanian OAE records is biased towards data derived from sections in Europe (Tethys) and North America (Western Interior Basin). In contrast, very limited information on mid–Cretaceous OAEs is available from the eastern sub-equatorial Pacific and generally, the western South American realm. A limited series of relatively well dated, albeit usually incomplete records of Albian to Cenomanian marine strata from the Pacific stems from ocean drilling projects (e.g., Hess Rise, Shatsky Rise, Resolution Guyot: Price, 2003; Robinson et al., 2004; Dumitrescu et al., 2006; Ando et al., 2008). Other outcrop-based studies from the Pacific realm have been reported from Japan (Nemoto and Hasegawa, 2011 and references therein). Judging from available data, OAE1a and OAE2 seem to be more represented in the Pacific region whilst other OAEs show a less pronounced record. Albian–early Cenomanian occurrences of organic-rich black shales seem to be less thick and rather patchy in distribution (Robinson et al., 2004). In conclusion, the eastern sub-equatorial Pacific is remarkably underrepresented and poorly constrained with respect to continuous and well-dated chemostratigraphic reference sections for the mid- Cretaceous interval (Fig. 1).

In order to close this gap, a field-based project in northern Peru has been undertaken and extended Albian–early Cenomanian sections are here reported. The results shown indicate that OAE1b, 1c and 1d are recorded in the Andean Basin of Peru. This paper has the following aims: To (i) document und interpret the sedimentology of well-exposed Albian–early Cenomanian sections of the Andean Basin of Peru; to (ii) provide a carbon-isotope reference curve from these sections for the marginal SE Pacific and to (iii) discuss and correlate the Peruvian findings with coeval records documented from the Tethyan, the proto-Atlantic and the Pacific domains.

Fig. 1. Palaeogeographic map of Gondwana during the mid–Cretaceous (modified after Larson and Pitman, 1972; Torsvik et al., 2009; Moulin et al., 2010) indicating position of what today is the Andean Basin. Black square denotes approximate position of study area.

2. Regional tectonic and stratigraphic setting

The information concerning the regional tectonic can be found in chapter 1.6.

Given the scarcity of stratigraphic data on the Albian–early Cenomanian of Peru, a summary of the existing knowledge is given for reference (Figs. 2, 3 and 4; Benavides- Caceres, 1956; Jaillard, 1987; Robert et al., 2009). In the northern Andes (Cajamarca region), the Lower Cretaceous is represented by the Goyllarisquizga Group that encompasses the Chimu, Santa, Carhuaz and Farrat formations (Benavides-Caceres, 1956). The Goyllarisquizga Group is assigned to the Valanginites broggii Zone at the base and an Aptian age was assumed for the top (Benavides-Caceres, 1956). This group is overlain by Albian transgressive deposits that resulted in shelf deposition of the Inca, Chulec, Pariatambo and Yumagual formations reaching the onset of the early Cenomanian.

The Inca Formation consists of iron-rich, sandy and marl-limestones beds, assigned an early Albian age based on the ammonite Neodeshayesites nicholsoni (Robert and Bulot, 2004; Robert et al., 2009). The Inca is unconformably overlain by the Chulec Formation with a discontinuity surface separating the two units. The Chulec Formation is characterized by marl-limestone alternations with a very abundant and diversified outer shelf fauna (Jaillard, 1987). Numerous ammonites have been reported and were assigned to the Knemiceras raimondii Zone (Robert and Bulot, 2004; Robert et al., 2009), indicating a middle early Albian–early middle Albian age. The Chulec Formation is conformably overlain by the Pariatambo Formation. 33

Fig. 2. Mid-Cretaceous palaeogeographic map of Peru (based on Pindell and Tabbutt, 1995; Robert and Bulot, 2004). Black square indicates study area. Note size of study area relative to the vast dimensions of the Western Platform. A–A´ denotes position of transect in Fig. 3.

Fig. 3. Hypothetical transect across the Andean Basin during Neocomian–Coniacian times (based on Benavides- Caceres, 1956; Atherton and Webb, 1989; Jaillard, 1987) with inferred Western Platform ramp morphology. Refer to approximate position of transect in Fig. 2 (A–A´).

The Pariatambo Formation is characterized by fossiliferous, black, bituminous, fetid marly limestones facies and includes fine lamination and siliceous intervals. Ammonites and planktonic foraminifera are common, evidencing the Oxytropidoceras carbonacrium and Prolyelliceras ulrichi Zones assigned to a middle Albian age (Robert et al., 2009). The Pariatambo Formation grades upsection into nodular bioclastic grey marl-limestones of the Yumagual Formation containing few ammonites. Bivalve biostratigraphy places the Yumagual Formation within the Ostrea scyphax and Exogyra mermeli Zones, indicating a late middle Albian–early Cenomanian age (Benavides-Caceres, 1956).

Fig. 4. Cretaceous chronostratigraphy of the Northern Andes. Note Goyllarisquizga Group, Inca, Chulec, Pariatambo and Yumagual formations and corresponding biostratigraphic framework (Benavides-Caceres, 1956; Robert and Bulot, 2005; Robert et al., 2009). Variations in seawater 87Sr/86Sr values during early Albian–early Cenomanian times are shown (after Howarth and McArthur, 1997; McArthur et al., 2001). Timescale is from Ogg and Hinnov (2012). Horizontal lines indicate strontium isotope values of best-preserved oysters used in this study. Numerical ages are derived following the procedures of Howarth and McArthur (1997).

3. Methods and materials

3.1. Field work and thin-section microscopy

Two well-exposed sections (Pulluicana and Huameripashga) were chosen for their well- exposed strata of the early Albian–early Cenomanian age. Sections were located along the 35

road from Cajamarca to Celendin (Fig. 5). In total, ca. 450 m of section have been logged and studied following the method describes in chapter 2.1.

Fig. 5. Geological map of the Cajamarca region (northern Andes) showing main structural elements and location of Pulluicana and Huameripashga study areas (modified after INGEMMET, 2006).

3.2. Assessing the diagenetic alteration of shell material

3.2.1. Cathodoluminescence microscopy

Cathodoluminescence (CL) microscopy examination of selected outer shell layers of five oyster specimens were investigated for their CL pattern using the analytical methods described in chapter 2.2.1

3.2.2 Trace and major element analysis

Oyster shells that passed initial cathodoluminescence inspection were subsequently analysed for major and trace elemental composition (Table 1) using the analytical methods described in chapter 2.2.2.

87 86 Age Locality Sample no. Sr Mg Fe Mn Sr/ Sr 2σ mean [Ma] [ppm] (×10−6) Min Mean Max

Huameripashga Hua 199 380 2598 720 76.6 0.707435 7 97.93 101 104.12

Pulluicana Puy 91.3 825 873 147 6.3 0.707380 6 108.87 109.36 109.74

Table 1. Selected elemental and 87Sr/86Sr data of low–Mg calcite oysters from the Cajamarca region (Western Platform, Peru) and derived numerical ages.

3.3. Carbon isotope stratigraphy

13 3.3.1. Bulk micrite (δ CCarb)

Carbon-isotope analysis were performed on 123 micrite (see Appendix A) from the Albian and early Cenomanian interval using the analytical methods described in chapter 2.3.1.

13 3.3.2. Bulk organic matter (δ Corg)

Carbon-isotope analyses were performed on 84 bulk organic matter samples (see Appendix B) from the Albian and early Cenomanian interval using the analytical methods described in chapter 2.3.2.

3.4. Strontium-isotope analysis

Sample powders of low-Mg calcite shells of 5 oysters were taken by strontium isotope analyses as described in chapter 2.4.

4. Data description and interpretation

4.1. Facies associations and depositional environments

The mid-Cretaceous facies belts along the Andean Basin of South America crop out over a north-south extension of about 2000 km (Figs. 1 and 2). Hence, despite the wide regional 37

extent of sections measured in the context of this study, the observational window represented by our data is dwarfed by the size of the Cretaceous Western Platform. Having said this, the direct match of coeval portions of measured chemostratigraphic sections argues for an over-regional significance of the data shown here. More fieldwork in neighbouring areas will support this working hypothesis or claim revision.

Fig. 6. Epeiric ramp model for the Western Platform in Peru exposed to occasional storm events and depositional environments recorded in the Inca, Chulec, Pariatambo and Yumagual formations of the Cajamarca region. Facies types (1a through 3a) are indicated. Refer to Table 2 for more details.

Table 2. Overview of facies classification and interpretation. Numbers indicate the relative abundance of non- skeletal and skeletal components: 0 = absent, 1 = present, 2 = frequent, 3 = abundant, 4 = dominant 38

The data obtained so far and previous work point to a mixed, carbonate-siliciclastic ramp with decreasing argillaceous influence towards the west (i.e. Pacific-wards) pinching out against the Albian volcanic arc (Figs. 1 and 2). The most proximal settings represented by coastal, deltaic and tidal flat facies in the eastern part of the platform (present-day eastern Peru and western Brazil, Fig. 2) were not visited in the context of this project. Integrating previous work (e.g., Benavides-Caceres, 1956; Robert et al., 2009), three main depositional environments, each with a number of standard facies types, are here established for the subtidal to outer ramp (Fig. 6 and Table 2). The documentation of these fundamental sedimentological data is important due to the lack of accessible previous work. Estimates of the approximate bathymetry of corresponding depositional environments are based on previous works dealing with comparable depositional environments (Yanin, 1983; Lukasik et al., 2000; Immenhauser, 2005, 2009).

4.1.1. Shallow subtidal inner ramp setting

The shallow subtidal inner ramp depositional environment comprises three facies types: 1a, 1b and 1c (Fig. 6 and Table 2). The most proximal deposits are composed of grey argillaceous mudstone and occasionally iron-rich sandstone, representing intervals of two to five metres thick (upper portions of Inca Formation). Scarce fauna occurs as undifferentiated shell debris and flora as plant remains (Facies 1a). In more distal sub-environments, the beds show thickening-upward packages and the facies is characterized by packstones and argillaceous wackestones, exhibiting a nodular fabric due to an increased argillaceous content, intraclasts and oncoids are also observed in variable amounts (Facies 1b; Fig. 8A). Evidence for bioturbation is limited and primary sedimentary structures such as plane bedding are well preserved (Fig. 7A).

Towards more open marine settings, Facies type 1b grades into thicker beds, characterized by grainstone and occasionally floatstone facies and reaching 1 to 3 m in thickness (Facies 1c, lower portions of Chulec Formation). Facies 1c is typically rich in mixed and fragmented shell debris, mainly of oysters and gastropods and echinoderms (Fig. 8B). Bioturbation is locally more intense and the bedding blurred. Abundant minor discontinuity surfaces are observed. These are characterized by FeO mineralization and increased levels of bioturbation of underlying rocks.

Fig. 7A): Outcrop image of thickly-bedded Yumagual Formation. B) Bioclastic rudstone of the Yumagual Formation, characterized by oysters in life position in shell debris host facies (Facies 2a). C) Oyster specimen used for strontium analyses in this study (Facies 2b). D) Discontinuity surface (SB1) that marks the contact between medium-scale sequences (MsS) 1 and 2. Note olistolithic block. MsS 2 displays thin dark-grey to thick-bedded limestones of the Pariatambo Formation.

The lithological and palaeontological characteristics of the innermost subtidal ramp facies exposed in the study area supports a restricted, low-energy environment, where punctuated tempestite intervals, abundant terrigenous influx including plan remains and the scarcity of a marine ichnofacies indicate a significant level of continentality. Estimates of the palaeo-bathymetry as based on comparable settings elsewhere point to water depths of a few metres only (Facies 1a).

Limestones deposited in the more distal inner ramp setting, i.e. towards the west, display a more diverse fauna and much decreased terrigenous influx suggesting a protected, shallow subtidal environment under low hydrodynamic energy near the expectedly very shallow fair- weather wave base (Immenhauser, 2009; Facies 1b). The presence of transported, fragmented and worn skeletal grains points to periodical storm events (Facies 1c). In the absence of corals, calcareous green algae or ooidal facies, an estimate of the palaeo- bathymetry of the distal inner ramp is difficult. In comparison to data shown in Yanin (1983) water depths from near-sea level to about 5 m are tentatively estimated for deposits of facies type 1c (Fig 8B). In this sense, it seems likely that facies types 1a through c are not differentiated by bathymetry but rather through their decreasing level of clastic influx and increasingly marine water masses.

Fig. 8A). Packstone of Chulec Formation composed of oncoids and variegated shell debris (Facies 1b). B) Bioclastic floatstone of the Chulec Formation, showing mixed and fragmented skeletal debris (Facies 1c). C) Wacke- to packstone of the Chulec Formation, showing some gastropods, echinoderms and unspecified shell debris (Facie 2c). D) Mudstones with Chondrites burrows of the Pariatambo Formation (Facies 3a). 41

4.1.2. Open marine middle ramp setting

The more open marine, middle ramp depositional environment comprises facies types 2a, 2b and 2c (Fig. 6 and Table 2). Deposits of the mid-ramp setting are characterized by packages of oyster bioherms, interbedded by rare grainstone units yielding various skeletal elements and intraclasts (Facies 2a; Figs. 7B). Oysters are a dominant faunal element and associated with echinoderms and occasional ammonites. Skeletal elements are locally imbricated, elsewhere bioclasts are arranged in lenses and display a chaotic fabric of non- sorted and fragmented bioclasts. Oyster packstones represent bioherms with dimensions of between 5 and some tens of metres and thicknesses of up to one metre (Facies 2b; Fig. 7C). Basin-wards, beds display thinning and are represented by bioturbated mud- to wackestones, less commonly by packstones. Faunal elements consist of gastropods, echinoderms and planktonic foraminifera and of unspecified shell debris (Facies 2c; Fig. 8C).

Generally, oyster-debris facies witness storm hydrodynamics that eroded, entrained and re-deposited oyster bioherms that were otherwise situated beneath the effective fair-weather wave base (Facies 2a), i.e. the depth at which wave-orbitals are still capable of moving sediment particles (Immenhauser, 2009). According to Yanin (1983), oyster bioherms in mid-ramp settings are often situated at palaeo-water depths of between 10 and 20 metres (Facies 2b). In modern, wave-exposed oceanic margins the average effective fair-weather wave base is in the order of 30 ± 15 metres. Given the broad epeiric ramp setting studied here, fair-weather and swell waves from the Cretaceous Pacific Ocean probably lost much of their energy due to wave base-seafloor interaction and the shallower end-member of this bathymetric range is expected (Keulegan and Krumbein, 1949). Moreover, the significance of the Albian volcanic arc to the west is difficult to quantify in terms of wave climates. The combination of oyster bioherms, combined with gastropods and echinoids, the presence of planktonic foraminifera and ammonites may indicate an open marine setting with a water depth between 10–30 m.

∼

4.1.3. Outer ramp setting

The outer ramp setting is characterized by facies type 3a (Fig. 6 and Table 2) and is located below the reach of the effective storm wave base. The facies obtained primarily consists of mudstones in dm-thick beds and yields traces of sulphide deposits (pyrite 42

content). The carbonate facies is dark grey in weathering colour (Fig. 7D). Chondrites-type bioturbation is common (Fig. 8D). The faunal composition includes planktonic foraminifera and echinoderms (Facies 3a). Fine-grained deposits characterized by dark grey colouring, fine lamination, high pyrite content, chondrites and scarce fauna may indicate an environment below the reach of storm waves in relatively dys- to anoxic water conditions at the sea floor.

In modern, open storm exposed oceanic margins, the effective storm wave base is in the order of 150 metres or less (Immenhauser, 2009). Considering the epeiric shelf and the potential barrier effects of the Albian volcanic arc to the west (Atherton and Webb, 1989), however, this depth range seems unlikely due to wave base frictional energy loss and a depth range of 30–50 metres is suggested in a tentative manner. This is the typical depth range of smaller, epeiric-neritic basins that might represent some sort of an analogue of the broad Peruvian shelf in the mid–Cretaceous but the modern world provides no suitable analogues to the Cretaceous Western Platform of Peru.

4.2. Sequence stratigraphic interpretation

Acknowledging the fact that the present data set covers a comparable small portion of the vast Western Platform of the Cretaceous of Peru, a first-order assessment of the sequence stratigraphy of these sections is here presented. The significance of these data lies fundamentally in the fact that they represent the foundation for future work from this poorly documented part of the world and allow for a comparison with sequence stratigraphic interpretations from neighbouring basins. The facies analyses of the Pulluicana- Huameripashaga composite section indicates changes in accommodation space of a least two different orders of magnitude. Essentially, the Inca, Chulec, Pariatambo and Yumagual formations represent a large-scale sequence (LsS1) with a minimum stratigraphic thickness of 450 m (Figs. 9 and 10). As the stratigraphic top of the Yumagual Formation is not reached in the outcrops visited, LsS1 is perhaps stratigraphically thicker than documented here. Overall, large-scale sequence 1 is made up by three medium-scale sequences with thicknesses ranging between 125–150 m, commonly delimited by well-marked discontinuity surfaces indicating changes in the depositional system including abrupt shifts of facies belts (Fig. 9).

Medium-scale sequence 1 (MsS1) overlies the argillaceous mudstone and iron-rich sandstone alternations of the Inca Formation. At the base, the transgressive system tract is represented by an alternation of wacke- and mudstone beds intercalated with some floatstone-packstone units (Chulec Formation). It is topped by a discontinuity surface 43

bearing field evidence of a marine hardground, which marks the maximum flooding surface and, thus, the contact between the Chulec and the Pariatambo formations. The hardground is overlain by more calcareous and thicker bedded highstand deposits comprising wacke- to floatstone units that are often bioturbated. A regionally significant discontinuity (SB1) characterized by the impregnation with secondary iron-oxides and intense bioturbation is recognized in both, the Huameripashga and Pulluicana sections respectively. On this discontinuity surface an olistolithic block is observed (Fig. 7D) that may indicate local tectonic faulting along the northern Andes (Jaillard, 1987). In terms of its sequence stratigraphic significance, SB1 marks the transgressive surface of the MsS2 (Fig. 9). Due to the lack of evidence for prolonged subaerial exposure of SB1, such as karstification, bleaching or saw-tooth-shaped carbon and oxygen-isotope excursions, the medium-scale sequences shown in Figure 9 might represent parasequences topped by marine transgressive surfaces.

Medium-scale sequence 2 (MsS2), i.e. the transgressive systems tract comprises wackestones, locally with pack- float- and rudstone alternations and some rare calciturbidites. The maximum flooding interval is probably represented by thinly-bedded, black argillaceous facies and, thus, the contact between the Pariatambo and the Yumagual formations. Upsection, this facies grades into highstand deposits characterized by packstone with rud- to floatstone alternations (lower to middle portions of the Yumagual Formation). At the top, discontinuity surface SB2 is characterized by impregnation and intense bioturbation and marks the top of MsS2 and represents the transgressive surface at the base of MsS3 (Fig. 9).

Medium-scale sequence 3 (MsS3) comprises the upper exposed portions of the Yumagual Formation. Here, transgressive deposits are represented by alternations of packstones and argillaceous nodular wackestones, overlain in turn by a minor marine discontinuity that may well define the maximum flooding surface. Highstand deposits are represented by thickly- bedded, grey wackestones that grade upsection into even thicker bedded, bioturbated grain- and floatstones facies. The top of the Yumagual Formation is not exposed in the study area.

5. Chemostratigraphy

5.1. Radiogenic strontium isotope stratigraphy

Fig. 9. Regional sequence stratigraphic interpretation, carbon-isotope stratigraphy and relative sea-level record of the Pulluicana-Huameripashga composite section with indication of Albian OAEs and corresponding sub- levels. Pe1–Pe9 refers to chemostratigraphic segments defined for the Peruvian reference curve. Data points shaded grey correspond to OAEs based on bio- and chemostratigraphic data. Note position of 87Sr/86Sr data. Biostratigraphic data by Benavides-Caceres (1956) and Robert et al. (2009).

Fig. 10. Legend for Fig. 9, denoting colour codes for different facies types and corresponding carbonate factory.

In light of limitations related to available biostratigraphic data (e.g., Benavides-Caceres, 1956; Robert et al., 2009 and references therein) from the Inca, Chulec, Pariatambo and Yumagual formations, it is critically important to apply independent stratigraphic tools in order to better constrain the chronological framework (Fig. 4 and Table 1). Strontium-isotope (87Sr/86Sr) analyses of well-preserved low–Mg calcite biominerals (rudists, belemnites, brachiopods and oysters) represent a powerful tool in chemostratigraphy (Huck et al., 2011; Mutterlose et al., 2014) that is well adapted for the mid–Cretaceous. The first and most important step, however, is the evaluation of well-preserved biogenic carbonate material for further analysis.

The outer shell layer of a series of macro- and mesoscopically well-preserved oysters has been investigated for their diagenetic overprint (Fig. 11). Intact valves and/or fragments of oysters display various levels of preservation. Cathodoluminescence microscopy revealed well-preserved valves with formerly organic-rich layers now replaced by secondary luminescent phases (Figs. 11A and 11B), whilst others display growth increments replaced by secondary luminescent carbonate phases (Figs. 11C and 11D). The latter specimens were not further considered here. Well-preserved valves were selected for further geochemical analyses (Sr, Mg, Fe and Mn elemental abundances). High manganese and iron concentrations (>100ppm) in calcitic oyster shell are used to determine recrystallization under reducing conditions (Huck et al., 2011 and references therein). Two out of five selected oyster specimens fulfilled these boundary conditions and were used for chemostratigraphy. As shown in Table 1, these shells yield Sr elemental abundances ranging from 380–825 ppm and Mn elemental abundances ranging from 6–76 ppm. Iron abundances exhibits more than 100 ppm, an observation that is in agreement with the mild degree of alteration observed 46

under the cathodoluminescence microscope being spatially limited to formerly organic-rich shell portions.

Numerical ages are derived from mean values of theses best-preserved samples using the `look up´ table of McArthur et al. (2001). Upper and lower age limits were obtained by adding two standard errors of the mean values of isotopic results to the statistical uncertainty of the seawater curve (Howarth and McArthur, 1997). The results indicate numerical ages of 109±0.4 Ma (early middle Albian) for the middle part of the Chulec Formation and 101±3 Ma (late Albian–early Cenomanian) for the Yumagual Formation (Huameripashga section; Table 1). Combined with the results of the detailed carbon-isotope curve documented below and previously published biostratigraphic data (e.g., Benavides-Caceres, 1956; Robert et al., 2009 and references therein) these chemostratigraphic pinning points provide a solid fundament for the stratigraphic framework applied here.

Fig. 11A): Cathodoluminescence photomicrograph of oyster shell, dark blue luminescence indicates well- preserved shell material. B) Same image under crossed polarizers. C) Cathodoluminescence photomicrograph of altered shell. D) Same image under crossed polarizers.

5.2. Carbon-isotope stratigraphy

13 13 The δ Ccarb and δ Corg data sets as described here were subdivided in chemostratigraphic

13 segments (Pe1–Pe9; Fig. 9). In the case of the Inca Formation, data are limited to δ Corg due to the very low carbonate content.

13 The δ Corg section commences with an abrupt negative shift reaching -26.3‰ (Pe1), continued by a recovery phase (Pe2) and followed by Pe3 that exhibits a gradual negative shift reaching -27‰. Further upsection, Pe4 reaches a recovery phase and Pe5 represents a

13 plateau. The overlying Pe6 shows a sharp δ Corg decrease down to values as low as -28‰.

13 This feature is followed by a recovery phase (Pe7). Upsection, δ Corg values exhibit a short negative shift (Pe8) and finally a plateau phase (Pe9).

13 The δ Ccarb section starts with an abrupt negative shift reaching values as low as +0.1‰ (Pe3), continued by a recovery phase (Pe4). Thereafter, Pe5 shows a plateau that oscillates

13 between 0 and +2.3‰, Pe6 presents a sharp δ Ccarb decrease reaching -0.1‰, followed by

13 relatively invariant values (Pe7). Pe8 displays a pronounced negative shift in δ Ccarb ratios, reaching a minimum of -3‰ (the lowest values measured in these sections). Upsection,

13 δ Ccarb values display a gradual return to pre-excursion values (Pe9a), followed by a plateau with values in the order of between +1.3–+2.1‰ (Pe9b).

6. Discussion

6.1. Large-scale depositional setting of the Western Platform

Controls that play a main role in the type of mid–Cretaceous neritic carbonate deposition (Lukasik et al., 2000; Schlager, 2005) include seawater temperature, degree of restriction, nutrient levels or latitudinal diversity gradients. Data shown here document the predominance of heterozoan skeletal elements in the Western Platform of Peru. In tropical settings, heterozoan carbonate secreting organisms are perhaps most common in open marine and relatively cool seawater settings (Philip and Gari, 2005), or where enhanced nutrient levels prevail (Halfar et al., 2004). Hence, a potential cause for the dominance of a heterozoan association in the palaeo-tropical setting of Peru is best found in the lower to middle Albian transgression that likely favoured an ingression of nutrient-rich oceanic water masses from the Pacific or the Atlantic. Nevertheless, evidence for the establishment of an open marine circulation in the Andean basin is indicated by the occurrence of ammonite taxa of Tethyan affinities (Robert and Bulot, 2005). According to Trabucho-Alexandre et al.

(2010), nutrient-rich surface currents were flowing along the southern margin of the proto- Atlantic into the Pacific via the northern margin of South America. It is at least conceivable that some of these water masses were flooded onto the Andean epeiric ramp. Moreover, the Palaeo-Pacific Ocean to the east represents a potential upwelling zone for cooler, nutrient- rich bottom waters similar to the present-day situation. However, the Cretaceous ocean water structure and current patterns were probably considerably different from those of the modern Pacific (D'Hondt and Arthur, 1996; Hay and Floegel, 2012) and differences in the plate tectonic setting were of importance, too. Finally, both the Marañon Massif to the east, as well as exposed continental basement in present-day Brazil in the east (Figs. 2 and 3) formed an important and spatially large source area for continental runoff delivering freshwater, nutrients and terrigenous influx. Judging from the argillaceous facies in the innermost exposed ramp settings (facies type 1a; Fig. 6 and Table 2), Albian transgressive pulses might have rework clay and silt facies in the flooded coastal settings.

Taking these potential nutrient-sources into account, the overall scenario of a mixed carbonate-siliciclastic ramp - protected perhaps locally by the Paracas structural high from impinging Cretaceous Pacific swell - is proposed (Fig. 2). A potential analogue, albeit not in size, is perhaps found in the Tertiary Murray Basin in SE Australia (Lukasik et al., 2000), whose environmental deposits pass gradually from a low energy shallow subtidal environment to more open settings and, then into outer ramp environment in sub-storm- wave environments. There, water masses were at least temporally sub-oxic in nature (compare with Fig. 6).

6.2. A carbon isotope reference curve for the sub-equatorial eastern Pacific

6.2.1. Bulk-micrite carbon isotope ratios

Secular changes in the isotopic composition of the dissolved inorganic carbon (DIC) pool of the world’s oceans are, under favourable conditions, recorded by marine carbonates

13 (δ Ccarb) resulting in characteristic chemostratigraphic patterns. This approach is widely accepted and despite numerous problems and shortcomings, mid–Cretaceous OAEs are

13 recognized and documented from bulk micrite δ Ccarb curves worldwide (Jarvis et al., 2002; Weissert and Erba, 2004; Erbacher et al., 2005; Jenkyns, 2010; Wendler, 2013; Horikx et al. 2014).

In order to provide a chemostratigraphic data set that has the potential for supra-

13 regional correlation, patterns in δ Ccarb must at first be tested for diagenetic alteration

masking palaeoceanographic patterns (Marshall, 1992; Immenhauser et al., 2002). In shallow-neritic bulk micrite samples, meteoric waters related to transient subaerial exposure

18 13 stages may introduce O-depleted oxygen, whereas C-depleted soil-zone CO2 can alter carbonate samples (Allan and Matthews, 1977; Christ et al., 2012). Generally, carbon isotope values are less prone to diagenesis as the carbonates form the dominant C reservoir compared to oxygen ratios, sourced mainly from fluids (e.g. Marshall, 1992; Jarvis et al.,

13 2002). Moreover, fractionation of δ Ccarb during burial is almost insensitive to temperature- controlled processes (about 0.035‰ per °C; Lynch-Stieglitz, 2003). In this context, it is of importance that we found no clear field evidence such as karsting, bleaching or traces of the roots of land plants in any of our sections. Similarly, evidence for the typical saw-tooth- shaped isotope patterns described in Allan and Matthews (1977) and many subsequent papers is lacking. Similar to Jurassic ramps settings in Morocco (Christ et al., 2012), it is tentatively concluded that low-amplitude sea-level fall was insufficient to expose portions of the carbonate ramp studied here. Reasons for this might include rapid basement subsidence, a feature that is probably supported by the generally stratigraphically thick successions studied.

Given that the chemostratigraphy presented here is mainly based on bulk micrite data, the δ18O record of the analysed carbonates is not interpreted in terms of its palaeoenvironmental significance but serves to understand the degree of diagenetic alteration of sections studied. Specifically, when plotting the carbon versus oxygen isotope data from the Huameripashga and Pulluicana sections no correlation is observed (R² = 0.03) between the two. Often, but not always, the lack in covariance between carbon and oxygen isotopes serves as an indication of less than pervasive diagenetic resetting. Generally, the

13 δ Ccarb ratios found are typical for mid-Cretaceous oceanic signatures reported from many basins worldwide (Veizer et al., 1999; Prokoph et al., 2008; Schulte et al., 2011). An exception is found in the significantly more negative ratios in the lower Yumagual Formation. This feature is not related to changes in facies or subaerial exposure-related diagenesis and might represent a local palaeoceanographic pattern that awaits further study. Summing up, whilst we do not exclude minor degrees of diagenetic alteration, the data shown here support the concept of a carbon isotope curve that predominantly reflects patterns in oceanic DIC.

6.2.2. Bulk organic carbon isotope ratios

13 Patterns and amplitudes of δ Corg ratios may result from variations in the source of organic carbon as well as from selective preservation and degradation of organic compounds 50

during diagenesis (Bralower et al., 1999; Wendler, 2013; Trefry et al., 2014). In order to test

13 the significance of the δ Corg curve obtained here we compare the pattern obtained with that from inorganic carbon isotopes (Fig. 9). The observed similarities are considered encouraging whilst the two records also show distinct differences in the amplitude of carbon isotope excursions. These may be related to variations in photosynthetic fractionation, differential organic carbon preservation (e.g. Jenkyns, 2010; Wendler, 2013), differential sources of carbonate mud, differential diagenesis, facies change or hidden hiatal surfaces (Immenhauser et al., 2008 and references therein; Turpin et al., 2014). A prominent example

13 is the isotope segment Pe8 (Fig. 9) that has a high-amplitude in δ Ccarb, while we observe a

13 low-amplitude negative inflection in the δ Corg curve (~1‰). In conclusion, the comparison between organic and carbonate carbon isotopes supports the overall concept of a mainly palaeoceanographic record found.

6.2.3 Comparison with other reference curves for Albian OAEs

13 13 The Pulluicana-Huameripashga composite δ Ccarb and δ Corg chemostratigraphy results from the stratigraphically most complete intervals of each section. The perhaps best test for the over-regional significance of the isotope sections from Peru is their contrast-comparison with well-established reference curves from other basins. Here, we compare the Peru section with coeval successions in European Tethys (Vocontian Basin, France; Piobbico section, Italy), the Pacific realm (Western Platform, Peru; Pacific Guyots and Hokkaido, Japan) and the Western Atlantic domain (Sierra Madre, Mexico). Specific focus is on mid–Cretaceous OAE signatures (OAEs1b, 1c and 1d). In terms of the temporal framework, this correlation is primarily based on the available biostratigraphic and chemostratigraphic analysis as documented here. Following previous well-established approaches, we assign chemostratigraphic labels to characteristic segments of the isotope curves (Pe1 though Pe9; Fig. 12).

13 The isotope excursion Pe1 in the δ Corg curve from Peru is dated by the ammonite Neodeshayesites nicholsoni as early Albian in age (Benavides-Caceres, 1956; Robert et al., 2009). The Pe1 event is correlated with the Kilian Level in the European Tethys (i.e. the second organic-rich level of the OAE1b bundle; Herrle et al., 2004; Reichelt, 2005; Fig. 12). In the Vocontian Basin, this interval consists of black, organic-rich sediment located close to the Aptian–Albian boundary (Bréhéret, 1986). This level is tentatively correlated with the Monte Nerone black shale interval in the Umbria Marche Basin, Italy (Erbacher, 1994).

Fig. 12. Correlation of Peruvian δ13C reference curve for the subequatorial western Pacific with reference curves from the Central Pacific (Jenkyns and Wilson, 1999), the Western Pacific (Takashima et al., 2010), the Western Atlantic (Bralower et al., 1999), the Northern Tethys (Reichelt, 2005) and the Western Tethys (Erbacher et al., 1996). Black shale intervals are shaded grey.

13 13 The second negative excursion Pe3, observed both in the δ Ccarb and δ Corg records, is correlated with the lower Albian Paquier Level from the European Tethys (Fig. 12; Herrle et al., 2004, 2003; Reichelt, 2005). This is in line with the occurrence of Glottoceras raumundi within the Chulec Formation (Robert, 2002). The Paquier Level is associated with one of the most widespread carbon-isotope events of the mid–Cretaceous and typically characterized by marine black shale located in the middle of the Hedbergella planispira foraminifera Zone of the Vocontian Basin (Herrle et al., 2004). Furthermore, in the Western Atlantic, the Paquier Level is recorded by organic-rich sediments belonging to the Hedbergella planispira foraminifera Zone (Bralower et al., 1999). In terms of its δ13C amplitude (>1‰; Fig. 12), Pe3 in Peru is comparable with the mid-Pacific record (Jenkyns and Wilson, 1999).

In segment Pe5, a 87Sr/86Sr ratio of 0.707380 ± 0.000006 was obtained from a well- preserved oyster shell (Fig. 9). This value is typical for an early to middle Albian, post Paquier Level, marine strontium isotope signature (Kennedy et al. 2000) and is in good agreement with our interpretation of the carbon isotope stratigraphy. A good correlation of segment Pe5 in Peru is found with the Mexican record of Bralower et al. (1999). The subsequent negative excursion Pe6 recorded in the Pariatambo Formation is assigned to the middle Albian Prolyelliceras ulrichi ammonite Zone (Robert, 2002). Segment Pe6 occurs on top of a marine transgressive surface (SB1; Fig. 9) and spans the transgressive interval of the MsS2. Segment Pe6 is tentatively correlated with the Leenhardt Level recorded in the European Tethys (top of the OAE1b set; Herrle et al., 2004). The Leenhardt Level in the Vocontian Basin is characterized by a prominent, basin wide black shale interval (Bréhéret, 1986) in the planktonic foraminifer Ticinella primula Zone (Reichelt, 2005). Interestingly, in the Vocontian Basin, both the Paquier and the Leenhardt levels have been linked to transgressive phases (Bréhéret, 1994). A similar situation is observed in the Peruvian record. There, the two negative shifts (Pe3 and Pe6) fall in the initial transgressive phase of MsS1 and 2 (Fig. 9). This may evidence widespread flooding and associated carbon cycle perturbations both for the Paquier and Leenhardt levels.

13 13 The Pe8 negative shift in the Peruvian δ Ccarb and δ Corg section (Fig. 9) is also observed in limestones deposited during the Biticinella breggiensis Zone in Mexico, on Pacific Guyots and in Japan (Bralower et al., 1999; Jenkyns and Wilson, 1999; Takashima et al., 2004). In the Piobbico area in Italy, Pe8 occurs in the planktonic foraminifer Ticinella praeticinensis Subzone (Erbacher et al., 1996), but displays subdued amplitude only. In the northern

13 Tethyan domain, Pe8 is best correlated to a δ Ccarb negative shift observed in the Ticinella primula Zone, making this feature somewhat older than the other records. This discrepancy might be explained by the error bars in age models applied. Above this negative trend, the Amadeus Level (OAE1c, Erbacher et al., 1996) is coeval to a short-lived, small-amplitude δ13C excursion to lower values. This level, recorded in the Umbria Marche Basin was assigned to the Biticinella breggiensis Zone and the Ticinella praeticinensis Subzone (Central, Italy; Galeotti, 1998). In summary, the contrast-comparison of the Peruvian composite isotope curve with coeval sections elsewhere is indicative of a good comparability of Pacific and Western Atlantic records and indicates that the sections measured recorded, and preserved, an above-regional chemostratigraphy.

In the Yumagual Formation, a positive excursion labelled Pe9a (Fig. 12) is present in the

13 δ Ccarb curve whereas it is probably absent in organic carbon. In Europe, OAE1d is defined by a positive excursion (the Breistoffer Level; Bréhéret, 1997), characterized by black shale

13 horizons dated as Rotalipora appenninica Zone (Reichelt, 2005). In Peru, the δ Ccarb Pe9a segment corresponds to that in the Japan curve (Pe9a). The observed partial discrepancy in Peruvian curves is interesting but not uncommon in chemostratigraphy. Interpretations commonly brought forward include local patterns in organic carbon cycling or secondary features such as differential preservation of organic matter or differential diagenesis of

13 δ Ccarb.

6.3. Impact of environmental changes on Peruvian neritic carbonate factory

The carbonate-siliciclastic ramp deposits in Peru document temporal changes in sedimentary environments and relative sea-level fluctuations. Moreover, observed patterns of enhanced carbonate production or demise probably relate to the impact of transient carbon cycle perturbations (OAEs 1b, 1c and 1d).

Within the OAE1b set, the negative excursion of the Kilian Level is recorded in the North Atlantic and in the Western Tethys. Between sections, considerable similarities include the coeval occurrence of organic-rich facies (Bralower et al., 1999; Herrle et al., 2004; Trabucho-

13 Alexandre et al., 2011). In the Peruvian sections, the Pe1 δ Corg negative excursion is probably the equivalent of the Kilian Level (Fig. 9). In Peru, this event is recorded by marls and sandstones alternating with iron-rich facies of a shallow subtidal environment. These are clearly different lithological patterns when compared to the Tethyan and North Atlantic realms. Along the northern Tethyan margin (Helvetic Platform), the onset of the OAE1b set is

associated with the final demise of the neritic carbonate factory (Föllmi et al., 2006) giving way to siliciclastic sedimentation and phosphogenesis. Due to a gap in exposure in the

13 Peruvian sections, the recovery of the negative δ Corg excursion is not recorded. The Pe2 excursion might, however, represent an increase in terrestrial organic matter influx during the turnover phase from a mainly argillaceous sedimentation to a neritic carbonate factory. This patter is in good agreement with the negative δ13C excursion of the Kilian Level, showing the characteristics of a global carbon cycle perturbation (Jenkyns and Wilson, 1999).

The expression of the Paquier Level in Peru (Pe3 segment) is, in terms of its facies, represented by an alternation of wackestones and dark-grey mudstones. Here, the base of the Pe3 segment corresponds to a change from a heterozoan-dominated carbonate factory to a decreasing benthic carbonate production in the lower portion of the MsS1 transgressive phase (Fig. 9). In our view, this shift in sedimentation represents an incipient phase of platform demise indicative of a global perturbation of the carbon cycle during the Paquier Level. The absence of organic-rich deposits in the Pacific realm (Peru and mid-Pacific Guyot) indicates that the accumulation of organic matter was either due to local signatures in carbon cycling superimposed on global patterns (Robinson et al., 2004) or due to the deposition (or preservation) of organic matter preferentially in deeper settings.

In Peru, the isotopic signature of the Leenhardt level (Pe6; Fig. 12) is found in dark-grey mudstones. These probably correspond to a major demise of neritic carbonate production marked in the field by a major discontinuity surface associated to the MsS2 transgressive surface (SB1, Fig. 9). Conversely, in the Tethyan and Atlantic deep-water record, the most prominent expression of the OAE1b event is assigned to the Jacob sub-event associated with the final demise of the neritic carbonate factory (Föllmi et al., 2006), whereas in the Western Platform of Peru the OAE1b bundle is best recorded by its final sub-event (Leenhardt Level).

Following the drowning related to the Leenhardt Level, the Peruvian sections reflect condensed sedimentation typified by a heterozoan biofacies in dark-grey mudstones. These condensed beds represent maximum flooding stages during LsS1 and MsS2 (Fig. 9). Deposition was probably influenced by ingression of more basinal, nutrient-rich waters from the Pacific during rising sea level (Trabucho-Alexandre et al., 2010). This oceanographic reorganization must have resulted in reduced carbonate production and episodic shutdown

13 13 of ramp carbonate deposition. This concept is in line with a shift in δ Ccarb and δ Corg towards more negative values (Pe8). Differences in amplitude may reflect local environmental patterns in the partitioning of carbon between organic carbon and carbonate or variations in the values of total dissolved inorganic carbon. Interestingly, the Pe8 event seems to be most pronounced in the Pacific sections (including Japan) and in Mexico,

whereas Tethys sections do not reveal this event. This may point to a circum-Pacific pattern not affecting the Tethyan realm (Fig. 12).

During the Albian–Cenomanian transition, an oyster-rich carbonate ramp is established.

13 13 Carbonate production and sedimentation rates are high. Here, both the δ Ccarb and δ Corg chemostratigraphy exhibit a plateau phase, suggesting stable sedimentation, oceanographic conditions and a phase of regional tectonic quiescence.

7. Conclusions

An Albian–Cenomanian composite section from the vast Western Platform in Northern Peru comprises a total of seven facies associations. These are assigned to a shallow to distal ramp environment with the most proximal portions not being exposed in the study area.

In the shallow subtidal inner ramp setting, argillaceous deposits are intercalated with restricted muddy carbonates. The open mid-ramp is mainly characterized by oyster biostromes established beneath the effective fair-weather wave base. Storms eroded, entrained and re-deposited oyster debris resulting in widespread oyster packstone to rudstone facies. The outer ramp setting is located below the reach of the effective storm wave base. Owing to the potential barrier effects of the Albian volcanic arc to the west (Marañon Massif) and wave degradation across the shallow ramp, wave-base depths were probably shallow.

Both, the Marañon Massif and the Brazilian Shield to the east of the Western Platform represent important source areas for continental runoff whilst currents transported proto- Atlantic and Pacific nutrient-rich water masses onto to the Western Platform as expressed in a tropical heterozoan carbonate factory. The Western Platform in Peru recorded

13 13 characteristic global patterns in Albian and early Cenomanian δ Ccarb and δ Corg ratios. The carbon isotope stratigraphy is constrained by biostratigraphic data and by two oyster-based 87Sr/86Sr marker points. The correlation with other reference sections in the Central and Western Pacific, the Western Atlantic and the Tethyan realm evidences that the Peru chemostratigraphy is of over-regional significance and has potential as a new reference curve for the subequatorial eastern Pacific realm. Important chemostratigraphic features observed include the OAE1b set (Kilian, Paquier and Leenhardt levels), OAE1c and OAE1d.

The negative excursion of the Kilian Level in the Peruvian data, representing part of the OAE1b set, is probably representative of a global carbon cycle perturbation and linked to the transition from a siliciclastic-dominated sedimentation to a neritic carbonate factory. The 56

expression of the Paquier Level corresponds to a change from a heterozoan-dominated carbonate factory to decreasing benthic carbonate production during incipient carbonate factory deterioration. The impact of the Leenhardt level in Peru is characterised by major demise of neritic carbonate production, marked in outcrops by a regional discontinuity. This surface is followed by condensed sedimentation, triggered by the influx of more basinal, nutrient-rich water masses from the Palaeo-Pacific. Given that a similar pattern is not known from the Tethyan realm, this feature seems circum-Pacific in nature. Further upsection, an Albian- Cenomanian transition oyster-rich carbonate ramp is established, suggesting stable environmental conditions.

The reference curve shown here is significant as (i) the vast area of the subequatorial eastern Pacific has previously not been covered in terms of its mid–Cretaceous carbon- isotope record and (ii) pronounced similarities with well-established chemostratigraphic patterns elsewhere evidence the over-regional significance of the Peruvian data. Nevertheless, the enormous dimensions of the Western Platform require the incorporation of regionally distributed coeval sections in order to shed light on spatial variations in carbon isotope stratigraphy.

Acknowledgements

This project was supported by the Deutsche Forschungsgemeinschaft (DFG, project n° BO-365/2-1) and by the Deutscher Akademischer Austauschdienst (DAAD) through a scholarship to a J.P.N. (PKZ: A/11/97387). We thank the Geological survey of Peru (INGEMMET) for important logistical support. W. Gomez is thanked for his help during field expedition. Analytical work was performed in the isotope laboratories at Bochum and Hannover by A. Niedermayr, D. Buhl and C. Wenske. The present manuscript benefitted from the comments of two anonymous Palaeo3 reviewers and editorial guidance by F. Surlyk.

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CHAPTER 4

ONGOING CENOMANIAN – TURONIAN HETEROZOAN CARBONATE PRODUCTION IN THE NERITIC SETTINGS OF PERU

Navarro-Ramirez J.P., Bodin S., Immenhauser A.

ABSTRACT

The present paper reports on the sedimentological and geochemical record of Albian– Turonian neritic carbonates from the eastern subequatorial Pacific domain in Peru. The focus is on one of the most extreme carbon cycle perturbations of the Phanerozoic, the Oceanic Anoxic Event 2 (late Cenomanian–early Turonian). Thanks to the very expanded and well- exposed sections in Peru, the OAE2 interval was sampled at high temporal resolution for both bulk micrite and bulk organic matter carbon isotopes. Despite the scarcity of significant amounts of organic matter or evidence for oxygen deficiency, the δ13C curve matches well with global published high-resolution data for coeval successions such as those reported from the English Chalk and the Portland # 1 core. Biostratigraphic data and the detailed sequence stratigraphic interpretation of these sections are combined with the carbon-isotope chemostratigraphy documented here. Applying the characteristic peak and trough chemostratigraphic terminology for OAE2 (A – C), the following main environmental and carbon isotope stratigraphic features are observed from the late Albian to the early middle Turonian in Peru: (i) An Albian to early late Cenomanian heterozoan ramp recording the pre- OAE2 δ13C excursions, specifically the Mid-Cenomanian Event. (ii) A late Cenomanian trough of δ13C values (B) showing a progressive deepening leading to the short-lived establishment of middle ramp type sedimentation. (iii) A late Cenomanian to early Turonian δ13C plateau (C) characterised by benthonic inner ramp sedimentation during a sea-level highstand phase. (iv) A recovery of δ13C values at the end of OAE2 associated to a trophic change, increased influx of argillaceous facies and reduced carbonate production. (v) A early to middle Turonian fluctuating δ13C curve, linked to a maximum flooding phase in the Mammites nodosoides Zone and carbonate production during the Collignoniceras woollgari Zone. The data shown here are particularly relevant as they come from very expanded neritic sections in the sub-equatorial eastern Pacific. Many of the features recognized share important similarities with Tethyan and Atlantic sections whilst the ramp system as such did not suffer from a carbonate crisis during OAE2 as recorded, for instance, in Mexico and Tibet.

1. Introduction

The mid-Cretaceous strata recorded one of the most extreme carbon cycle perturbations of the Phanerozoic (Oceanic Anoxic Event 2= OAE2), characterised by widespread organic- rich black shale deposition and a positive δ13C excursion (Schlanger and Jenkyns, 1976; Schlanger et al., 1987; Jenkyns, 1980). Since four decades, a considerable amount of research has focused on the causes and consequences of OAE2. Previous studies were mainly focused on Tethyan and proto-Atlantic hemi-pelagic and pelagic settings (e.g., Pratt, 1985; Pratt et al., 1985; Arthur et al., 1985, 1988; Elder, 1985, 1989; Kennedy and Cobban, 1991; Cobban, 1993; Kauffman, 1995; Dean and Arthur, 1998; Bowman and Bralower, 2005; Gale et al., 2005; Sageman et al., 2006; Friedrich et al., 2006; Mort et al., 2007a, 2007b; Voigt et al., 2006, 2008; Jarvis et al., 2006; 2011; Meyers et al., 2012; Ma et al., 2014; Du Vivier et al., 2014, 2015; Gambacorta et al., 2015). Despite a significant bulk of published data, the underlying controls of OAE2 are still debated, but during the two past decades a linkage between massive volcanism and OAE has been suggested (e.g., Larson, 1991; Larson and Erba, 1999; Mort et al., 2008; Du Vivier et al., 2014; Bodin et al., 2015).

The Cenomanian–Turonian OAE2 was preceded by the Mid-Cenomanian Event I (MCEI; Paul et al., 1994; Coccioni and Galeotti, 2003; Keller et al., 2004; Gertsch et al., 2010b; Giraud et al., 2013; Andrieu et al., 2015). The MCEI is defined by two positive δ13C peaks separated by a trough (MCEIa, MCEIb; Mitchell et al., 1996; Gertsch et al., 2010b; Giraud et al., 2013; Andrieu et al., 2015). Contrary to OAE2, MCEI is not characterised by organic-rich black shale deposition, but is recorded in Tethyan and North Atlantic settings as a major perturbation in carbonate platform productivity (Giraud et al., 2013). The Mid-Cenomanian Event I is also linked to a major sea-level fall recorded in the early–middle Cenomanian transition (Gale et al., 2002, 2008; Wilmsen, 2003, 2007; Wilmsen et al., 2005; Gertsch et al., 2010b; Giraud et al., 2013; Andrieu et al., 2015). Following the MCEI, long-term eustatic sea-level rise took place during the middle Cenomanian to late Cenomanian, resulting in the establishment of large shallow epicontinental seas (Haq et al., 1987; Hardenbol et al., 1998; Haq, 2014). This has led to the deposition of the basal Bridge Creek Limestone in the Western Interior Seaway (Cobban and Scott, 1972; Hattin, 1975, 1986; Kauffman, 1977; Elder, 1985; Dean and Arthur, 1998; Sageman et al., 2006; Ma et al., 2014), the extension of the carbonate platforms towards the North and South from the Tethys realm (Philip and Airaud-Crumiere 1991; Mort et al., 2008; Gertsch et al., 2010a, 2010b; Lézin et al, 2012; Bomou et al., 2013), accumulation of organic-rich black shale in deeper-open marine environments and formation of upwelling areas near to the low latitude North Atlantic and Tethys Ocean (Trabucho-Alexandre et al, 2010).

In organic-rich and in carbonate rocks, OAE2 is frequently recorded by a carbon isotope

13 excursion with an amplitude on the order of 2–4% in both bulk micrite (δ Ccarb) and organic

13 matter (δ Corg), exhibiting a typical pattern of subordinated peaks and troughs (named A – C), used as globally significant chemostratigraphic markers (Tsikos et al., 2004; Jarvis et al., 2006; Voigt et al., 2008; Du Vivier et al., 2014). Moreover, the positive δ13C excursions associated with OAE2 encompassed not only open-deep marine settings, but shallow water environments as well (Gertsch et al., 2010a; 2010b; Elrick et al., 2009; Lézin et al., 2012; Lebedel et al., 2013). The isotope shift represents a short-lived perturbation of ~600 Kyr duration according to astronomical time scales (Meyers et al., 2012; Ma et al., 2014). Published work on northern and southern Tethyan shallow-water settings report on the late Cenomanian–early Turonian demise of carbonate platforms, including rudists and large benthic foraminifera (Philip and Airaud-Crumiere 1991; Drzewiecki and Simo, 1997). This demise was marked by widespread extinction of marine biota (Sepkoski, 1996; Lamolda et al., 1997; Kaiho et al., 2014).

Fig.1. Palaeogeographic map of South America during the mid-Cretaceous (modified after Blakey, 2011) indicating the position of what today is the Western Platform Andean Basin (yellow arrow).

The vast Western Platform of Peru presents a hitherto underexplored archive of the OAE2 interval and generally, only few well-dated records from the sub-equatorial eastern Pacific have been reported (Fig. 1; Elrick et al., 2009; Du Vivier et al., 2015). The scarcity of black organic-rich shales in Cenomanian – Turonian Western Platform sections implies that the corresponding biotic and environmental effects differ significantly from those of, for example, the open Tethyan realm. Moreover, the significant continentality of this vast epeiric platform, nutrient runoff, local carbon cycles superimposed on provincial and global ones

and differences in the aquafacies render the direct comparison of these settings with open oceanic ones challenging (Thomas et al., 2004; Immenhauser et al., 2008). In order to close this gap, a field-based project in northern Peru with focus on the Cenomanian–Turonian epeiric sedimentary record has been undertaken and the outcome is reported here. Whilst the focus is on the OAE2, underlying sedimentological evidence is reported in some detail. This is important as no previous author has documented these sections in sufficient detail. Specifically, the aims of this paper are: to (i) document and discuss the sedimentological and palaeoecological features of Cenomanian–Turonian section in Peru; to (ii) provide a carbon- isotope chemostratigraphic reference curve for the eastern sub-equatorial proto-Pacific and to (iii) discuss and correlate the Peruvian sections chemostratigraphically with coeval records from the Tethyan and proto-Atlantic domains.

2. Regional tectonic and stratigraphic setting

The information concerning the regional tectonic can be found in chapter 1.6.

In the Cajamarca region (Northern Peru; Benavides-Caceres, 1956; Fig. 2), the Cretaceous epeiric-neritic system is represented by the Inca, Chulec, Pariatambo and Yumagual Formations (early Albian–early Cenomanian; Fig. 2; Navarro-Ramirez et al., 2015), being overlain by the Mujarrún, Romirón, Coñor and Cajamarca Formations studied in the context of this paper (lower Cenomanian– middle Turonian; Fig. 2 and 3A–E).

The Mujarrún Formation is characterised by marls and nodular limestone alternations. Common occurrences of Exogyra africana Lamarck, Exogyra olisiponensis Sharpe, Exogyra polygona von Buch, Neithea tenousklensis Coquand and Orthopsis titicacana are found, indicating a middle Cenomanian age (Benavides-Caceres, 1956). The Mujarrún Formation is overlain by the Rómiron Formation. The lower part of the Rómiron Formation is built by yellowish and brownish sandy marls, whereas its upper part is made of thickly-bedded limestones (Fig. 3C). Benavides-Caceres (1956) identified Lissoniceras mermeti, Forbesiceras sp., several Acanthoceras species, and Neolobites kummeli Benavides and abundant Exogyras, defining the Acanthoceras chasca Zone, attributed to a late Cenomanian age. This unit is overlain by the Coñor Formation (Fig. 3C) characterised by bluish to greyish marls that are in places nodular and intercalated with dark grey argillaceous limestone beds (Fig. 3D). The fauna includes Mammites nodosoides, Pseudoaspidoceras reesidei, Thomasites fischeri and Hoplitoides inca Benavides, suggesting an early Turonian age (Benavides-Caceres, 1956; Jaillard and Arnaud-Vanneau, 1993). The overlying Cajamarca Formation is characterised by light bluish to greyish massive, thickly bedded limestones (Fig. 67

3E). This unit yields a facies containing Coilopoceras newelli, Cardium lissoni Brüggen, Inoceramus peruanus Brüggen, Hemiaster fourneli Deshayes, and Cyphosoma peruanum, suggesting a middle to late Turonian age (Benavides-Caceres, 1956; Jaillard, 1987; Jaillard and Arnaud-Vanneau, 1993).

Fig. 2. Geological map of the Cajamarca region (northern Peru, San Marcos Map code: 15-g; modified after INGEMMET, 1980) showing location of Quebrada Chinchin and Piedra Parada areas of this study, as well as Pulluicana and Huameripashga locations of the previous study (Navarro-Ramirez et al., 2015).

3 Methods and materials

3.1 Field work and thin-section microscopy

Two well-exposed sections (Quebrada Chinchin and Piedra Parada localities) were chosen to provide maximum coverage of middle Albian to middle Turonian deposits (Fig. 2). In total, ca. 1000 m of section have been logged and studied following the method describes in chapter 2.1.

3.2. Carbon isotope stratigraphy

13 3.2.1. Bulk micrite data (δ CCarb)

Carbon-isotope analysis were performed on 298 micrite samples (see Appendix C) from the Cenomanian and Turonian interval using the analytical methods described in chapter 2.3.1.

13 3.2.2. Bulk organic matter data (δ Corg)

Carbon-isotope analyses were performed on 184 bulk organic matter samples (see Appendix D) from the Cenomanian and Turonian interval using the analytical methods described in chapter 2.3.2.

4 Data description and interpretation

4.1 Facies associations and depositional environments

Previously, Navarro-Ramirez et al. (2015) documented the development of a heterozoan mixed carbonate-siliciclastic ramp for the Albian–Cenomanian time interval of the epeiric- neritic Western Platform of Peru. Following this work, the fieldwork dataset, as well as biostratigraphic and thin section data from the Cenomanian–middle Turonian succession, led to the definition of seven facies representing three main depositional environments.

4.1.1 Shallow subtidal inner ramp setting

Facies 1a — Sandy marls and argillaceous mudstones (Fig. 4A). This facies is composed of yellowish to brownish sandy marls and grey argillaceous mudstone, showing mm- to cm- thick lamination in units of one to three metres thick (eg., lower part of the Coñor Formation). Scarce fauna occurs as undifferentiated shell debris and flora as plant remains. 69

Quartz grains are abundant, commonly angular, less often rounded (Fig. 4A). Given the scarcity of fauna and bioturbation and clay-rich facies, facies 1a is interpreted to evidence shallow, restricted marine, nearshore environments.

Fig. 3A): Limestones of the facies 2a, showing some ripples (Yumagual Formation). B) Oyster bioherms of facies 2a. C) Field image showing the thickly-bedded limestones of the medium-scale sequence 9 (MsS9; Romirón Formation) and thinly-bedded sandy marls and marly limestones of the medium-scale sequence 10 (MsS10; Coñor Formation). OAE2 is present within MsS9 where it exhibits the characteristic A–D segments. CTB= Cenomanian–Turonian boundary; SB9=sequence boundary 9. D) Thinly-bedded marls of the maximum flooding interval of MsS10. E) Thickly-bedded limestones of the highstand deposits of MsS10 (Cajamarca Formation).

Facies 1b — Oncoid- and miliolid-bearing wacke-packestones (Fig. 4B): This facies is made of brownish wacke- to light grey peloidal packstones, displaying dm- to cm-thick lamination in units of two to five metres thick. Bioturbation is rare, but primary sedimentary 70

structures such as plane bedding are preserved. Oncoids display variable sizes, but reach diameters of 4 mm in average (Fig. 4B). Nuclei commonly consist of a bivalve fragments coated by irregular micritic laminae. Next to oncoids, skeletal clasts with micrite envelopes, peloids and birdseyes are present. The fauna is dominated by miliolids, gastropods, ostracods and benthic foraminifera. Echinoids are sparse, oysters and planktonic foraminifera are rare. Associations of coated grains, birdseyes and skeletal elements such as miliolids, ostracods, gastropods and bivalves indicate deposition from nearshore shallow to deeper parts of the littoral zone (Flügel, 2004), with low-energy hydrodynamic conditions (Lézin et al., 2012) and slightly restricted environments (Elrick et al., 2009). Associations of large miliolids, oncoids and peloids reflect moderate trophic resources (James, 1997).

Facies 1c — Gastropod-bearing float-grainstones (Figs. 4C–D): This facies is built by dm to m-thickly-bedded grey packstones and occasionally by floatstones. At their base, beds present discontinuous horizontal laminae with scours and bioturbation. The fauna is abundant in gastropods, oysters and ostracods, commonly echinoids associated with coated grains of different sizes (oncoids and peloids). Facies 1c includes highly to moderately mixed and fragmented shell debris mainly consisting of gastropods, oysters, ostracods and echinoids. Following previous workers, this type of deposit indicates high-energy, shallow subtidal shoals and storm event beds (Elrick et al., 2009), just above the fair-weather wave base (FWWB; Lézin et al., 2012) and are linked to shallowing upward sequences during regressive intervals (Elrick et al., 2009). The predominance of gastropods suggests warm waters and normal marine conditions (Nield and Tucker 1985).

4.1.2 Open marine middle ramp setting

Facies 2a — Oyster-bearing floatstones, rudstones to biostromes (Figs. 3B and 4E–F): This facies is characterised by oyster bioherms and oysters- floatstones/rudstones. Individual oyster bioherms are 1 to 5 m thick and built by bluish grey massive oyster-biostromes (Fig. 3B), extending laterally over several hundreds of metres. The predominance of oyster- biostromes indicate restricted shallow sub-tidal environments (<20 m), high energy, low salinity aquafacies and mesotrophic nutrient level (Gertsch et al., 2010a, 2010b). In what might be an analogous setting, in shallow environments in Morocco, the oyster facies was associated with high kaolinite contents indicating predominantly humid climate during the Cenomanian (Gertsch et al., 2010a, 2010b).

Oyster floatstones represent thickening upwards sequences each one being 20 to 80 m thick, and they are situated in highstand deposits. Rocks are nodular in appearance a feature 71

typical for this facies. Oyster floatstones occur as bivalve shells with occasional gastropods and rather frequent ostracods, embedded in micrite matrix that displays crinoids, echinoids and planktonic foraminifera (Fig. 4F). Generally, oyster-debris facies witness storm hydrodynamics that eroded, entrained and re-deposited oyster bioherms that were otherwise situated beneath the effective fair-weather wave base, i.e. the depth at which wave-orbitals are still capable of moving sediment particles (Fig. 4F; Tucker and Wright, 1990; Immenhauser, 2009). According to previous workers, oyster floatstones are interpreted to indicate deep subtidal settings and low hydrodynamic levels (Lézin et al., 2012). On the other hand, the nodular features of the Facies 2a may indicate storm deposits, interpreted to be formed after early lithification in environments above the storm wave base (SWB; Burchette and Wright, 1992). All of this probably evidences a water-depth of 30–50 m in an open marine middle ramp environment (Herkat, 2007).

Facies 2b — Echinoid-bearing packstones to wackestones (Fig. 4G): This facies is bluish to brownish in weathering colour, showing dm to cm-thick laminations. The fauna is rich in echinoids and crinoids (Fig. 4G) skeletal material is associated with bryozoans, oysters, and planktonic foraminifera all floating in a micrite matrix. Conversely, gastropods and ostracods are rare. Skeletal elements are hydrodynamically sorted in places, probably representing distal storm beds, whilst they are chaotic elsewhere. Echinoderms are commonly found in proximal to middle ramp environments with normal salinity and perhaps elevated trophic resources (Nield and Tucker 1985; Lukasik et al., 2000) and are not especially tolerant of low oxygen conditions (Gale et al., 2000). However, their association with bryozoans, planktonic foraminifera and the scarcity of gastropods, oysters and ostracods may indicate a middle ramp environment (Flügel, 2004).

Facies 2c — Bryozoan-bearing wackestones to mudstones: This facies is bluish to dark grey in weathering colour, displays 10 to 50 cm-thick lamination, and consists of bioturbated mud- to wackestones, less commonly by packstones. Faunal elements consist of variably sized fragments of bryozoans, echinoids and planktonic foraminifera. Oyster shells are rare. In what might be analogous settings in Tasmania, the diverse bryozoan biota associated with echinoids and planktonic foraminifera represents deposition in open marine, low mesotrophic, shallow marine waters (Lukasik et al., 2000) between 50 and 80 water depth (Amini et al., 2004). Normally, bryozoan-rich beds are evidence of low terrigenous input and low sedimentation rate in the basin (Lézin et al., 2012) and suggest cool-temperate shelves (Flügel, 2004). The lack of fragmented and abraded shell debris, the frequent presence of bioturbation and the excellent faunal preservation suggests low hydrodynamic influence (Lukasik and James, 2003).

4.1.3 Outer ramp setting

Fig.4. Scanned thin-section images showing: A) Sandy marls displaying undifferentiated shell debris (Sd) in facies 1a. B) Packstone-bearing facies 1b showing oncoids (On), echinoids (Ech), gastropods (Ga) and bivalves (Bi). C) Grainstone showing fragments of oysters (Oy) and gastropods (Ga) in facies 1c. D) Packstone-bearing 73

facies 1c showing oysters (Oy), gastropods (Ga) and abundance argillaceous content (Si). E) Oyster (Oy) floatstone-bearing facies 2a. F) Oyster floatstone-bearing facies 2a, fabric is chaotic and characterised by unsorted and fragmented oyster bioclasts (Oy). G) Packstone- bearings facies 2b showing echinoid (Ech) and crinoids (Cri). H) Mudstone-bearing facies 3a.

Facies 3a — Planktonic foraminifera-bearing mudstones (Fig. 4H): This facies shows 5 to 10 cm-thick horizontal bedding and consists of dark grey mudstones with brown-ochre chert nodules. The chert nodules are oriented parallel to bedding. The fauna consist of planktonic foraminifera and rare benthic bioclastics. Storms beds are rare. The overall muddy facies, as well as planktonic foraminifera, suggest low-hydrodynamic levels in an open marine setting below the reach of the effective storm wave base (SWB; Lézin et al., 2012).

4.2 Sequence stratigraphic interpretation

A basic interpretation of the sequence stratigraphic scheme of the Western Platform was presented in Jaillard (1987) and Jaillard and Arnaud-Vanneau (1993). Here we document and discuss the first detailed scheme based on facies patterns, the stratigraphic framework and regionally significant discontinuities. This is considered important as the concept of sea- level fluctuations is particularly useful in understanding trophic resources and epeiric ramp carbonate production, two issues that are of major relevance here and that merit a certain degree of detail (Andrieu et al., 2015). Furthermore, the sequence stratigraphy model used follows the definition of Embry (2009), as well as that used by Andrieu et al. (2015) for shallow marine environments. This implies that the depositional sequences recognized here are bounded by sequences boundaries (SB), which represent either maximum regressive surfaces or transgressive surfaces (Van Wagoner et al., 1988). These surfaces indicate a shift from a shallowing-upward to a deepening-upward pattern (Embry, 2009; Andrieu et al., 2015). Moreover, a depositional sequence is a depositional cycle that comprises: transgressive deposits (TD) that mark a change to upwards deepening facies (retrograding structures); maximum flooding interval (MFI), a portion of the section that marks the transition from deepening-upward to shallowing-upward patterns; and a highstand deposit (HD), representing shallowing facies grading upwards into more proximal facies (prograding structures).

The stratigraphic framework of the sections studied is subdivided into two large-scale sequences. In this context, medium-scale shallowing-upward sequences, their origin and their association to possible trophic resources (Figs. 5–6) is discussed. Large-scale sequence 1 includes the lower Albian–lower Cenomanian (Inca, Chulec, Pariatambo and Yumagual

formations; Navarro-Ramirez et al., 2015) rock record and large-scale sequence 2 the middle Cenomanian–middle Turonian interval (Mujarrún, Romirón, Coñor and Cajamarca formations; this study; Figs. 5–6), respectively. One of the most prominent of the sequences boundaries is present at the base of the Mujarrún Formation (sequence boundary 7 = SB7; Fig. 5). There it marks a major change in depositional style from a regressive to transgressive trend that probably represents a hiatal interval. Below sequence boundary 7, six significant sequences boundaries (SB1–6; Figs. 5–6) mark the top of six medium-scale sequences (MsS1–6) and above sequence boundary 7, two others (SB8 and 9; Figs. 5–6) the top of two medium-scale sequences (MsS8–9). Medium-scale sequence 7 is topped by sequence boundary 7. The top of medium-scale sequence 10 was not reached stratigraphically in this study.

Medium-scale sequences 1–2 represent the Chulec and Pariatambo formations (lower Albian to middle Albian; Fig. 5), and comprise facies types 3a to 2a with thicknesses of about 100 m and represent the transgressive and maximum flooding interval of the large-scale sequence 1 (Navarro-Ramirez et al., 2015). The transition from medium-scale sequence 2 into 3 is gradual, passing upwards to shallower facies from mid-ramp to nearshore restricted environments until the top of medium-scale sequence 7. Medium-scale sequences 3–7 are all in all about 100 m thick, characterised by thickening-upwards nodular limestone beds, displaying a stacking pattern including parasequences topped by marine flooding surfaces (sequences boundaries 1–6; Fig. 5). These sequences represent the regressive trend of the large-scale sequence 1 and yield the massive nodular limestones of the Yumagual Formation (middle Albian to lower Cenomanian; Fig. 5). This parasequence set includes well-defined carbonate highstand-system-tracts built by facies types 2b to 1c and maximum flooding intervals that are, however, not well-defined here due to the overall shallow water depths of this epeiric-ramp. These facies patterns indicated that trophic resource levels increased consistently towards highstand deposits (Fig. 5) as reflected by successively more restricted marine environments, less fossiliferous and increasingly argillaceous-rich facies until the top of medium-scale sequence 7.

Medium-scale sequence 8 (MsS8, Fig. 5) is 160 m thick, corresponds to the Mujarrún Formation (middle Cenomanian to upper Cenomanian) and represents the beginning of the large-scale sequence 2. Medium-scale sequence 8 comprises a thickening-upward transition through facies types 1a to oncoid facies 1b, defining a nearshore restricted to inner ramp environments and perhaps moderate trophic levels. This unit is followed by an abrupt deepening trend to outer ramp sedimentation typified by facies 3a with abundant planktonic foraminifera, defining the maximum flooding interval. Further up-section, a thickening- upward succession from facies 2c to 2a, increasingly enriched in quartz grains and oysters (Fig. 4F) evidences high trophic levels of the highstand. The top of the medium-scale 75

sequence 8 is not exposed in the study area. Based on regional correlations across northern Peru this interval comprises thinly-bedded glauconitic and sandy marls interbedded with marly limestones (Jaillard and Arnaud-Vanneau, 1993). It is conceivable, that these units may represent a transgressive surface or transgressive deposits of the medium-scale sequence 9.

Medium-scale sequence 9 (MsS9, Fig. 5) is ca 60 m thick and comprises the Romirón Formation (upper Cenomanian). The transgressive system tract is represented by a proximal storm bed (facies 1c and 2a) at the base, rich in oysters and gastropods. As sea level continued to rise, these relatively shallow deposits passed gradually upwards into progressively deeper depositional environment of brownish marly echinoids-bearing limestones (facies 2b). With upward reduced creation of accommodation space, thickly- bedded gastropods limestones (Figs. 4C–D; facies 1c), indicate the highstand systems tract. As sea level subsequently dropped, shallower facies 1c passed gradually upwards via sequence boundary 9 into oncoid- and peloid-rich facies 1b and facies 1a (Fig. 4A). These facies patterns of medium-scale sequence 9 indicated that trophic resource levels decreased from high to moderate conditions (Fig. 9) as reflected by a transition from oyster-gastropod to oncoid-peloid facies until the top of medium-scale sequence 9. Above sequence boundary 9, thin sandstone and siltstone sheets (facies 1a) probably typify a regressive surface.

Medium-scale sequence 10 (MsS10, Fig. 5) is at least 160 m thick (but not fully exposed) and is more marly than the previous units. This sequence includes the argillaceous Coñor Formation and the lower part of the Cajamarca Formation (lower Turonian to middle Turonian). Facies 1a through 1b rich in oncoids and miliolids are well-defined in the lower part of the medium-scale sequence 10 (Fig. 4B), showing a nearshore restricted environment grading upwards into inner ramp type sedimentation. This interval is perhaps best assigned to the transgressive system tract. Upsection, facies 1b grades rapidly into bryozoan-facies 2c and then into outer ramp marly facies 3a rich in planktonic foraminifera (Fig.4H) i.e. the maximum flooding. The transition from facies 3a into the overlaying thickly-bedded limestones of facies 2c is gradual, marking the highstand system tract. Highstand deposits yield common echinoids, but scarce benthonic organisms (oysters, gastropods, etc.) and very low argillaceous content. These facies patterns in medium-scale sequence 10 indicate that trophic resource levels were moderate towards the top of medium-scale sequence 10 (Fig. 5), as inferred by successively oncoid-miliolid to bryozoan-echinoid facies, scarce macrofauna and low terrigenous input.

Fig. 5. Regional sequence stratigraphic interpretation and carbon-isotope stratigraphy of the Western Platform composite section, showing five-point moving average δ13Ccarb curve, carbonate production and relative sea-level change. Mid-Cretaceous OAEs and corresponding chemostratigraphic Pe1–Pe13 segments are indicated. Data points shaded grey correspond to OAEs based on bio- and chemostratigraphic data. Data in green of Pulluicana section is taken from Navarro Ramirez et al. (2015). Biostratigraphic data is taken from Benavides-Caceres (1956), Robert et al. (2009) and Jaillard and Arnaud-Vanneau (1993). Cenomanian zones have been obtained from oyster fauna associations and Albian as well as Turonian zones are based on ammonite fauna.

Fig. 6. Legend for figs. 5 and 8, denoting colour codes for different facies types and corresponding carbonate factory.

4.3 Carbon-isotope stratigraphy

Previous work (Navarro-Ramirez et al., 2015) subdivided the Albian of the Peru Western Platform into chemostratigraphic segments referred to Pe1–Pe9 segments, which represent

13 either positive/negative δ C shifts or features such as onset, trough, and plateau. Building on this approach, we subdivide the Cenomanian to Turonian δ13C chemostratigraphic record as described here, in terms of segments Pe10–Pe13 (Figs. 5, 7 and 8). Moreover, in order to characterize specific sub-intervals of the Cenomanian–Turonian isotope excursion, we apply a chemostratigraphic terminology including terms such as onset, trough, plateau and recovery (Paul et al., 1999; Tsikos et al., 2004) and refer to datum levels A, B and C as used in Tethyan and North Atlantic sections (Du Vivier et al., 2014, 2015; Fig. 8).

At the Quebrada Chinchin section, segments Pe5–Pe9 match those previously described in the Pulluicana-Huameripashga composite section situated 20 km away (Fig. 7; Navarro- Ramirez et al., 2015), a feature that documents the regional significance of these patterns. 78

Stratigraphically upsection in the lower to middle Cenomanian interval, segment Pe10 of the

13 δ Ccarb curve is characterized by low values ranging from –1 to +1.5‰, showing short-term

13 inflections from decline to rise, whereas the δ Corg curve also displays inflections from decline to rise ranging from –27 to –23‰. In segment Pe11, both δ13C curves rise from 1‰ to a maximum of +3.2‰ and –26 to –22‰ followed by a mild decrease in isotope values

13 13 towards the top reaching +0.5‰ in δ Ccarb and –26‰ in δ Corg values.

At Piedra Parada section in the late Cenomanian to middle Turonian interval, the lower part of the prominent carbon isotope excursion related to segment Pe12, was not covered due to poor exposures. However, the Pe12 segment displays a significant rise from mean values of

13 13 +1‰ to a first peak of +4.4‰ in δ Ccarb and a high ratio 19.5‰ in δ Corg, labelled as Pe12A sub-segment (Fig. 8). This sub-segment may represent the top of the onset of the Cenomanian–Turonian prominent positive excursion. Upsection, both bulk micrite and

13 13 organic carbon ratios are characterised by decreasing C-values of 3.4‰ in δ Ccarb and –

13 26.5‰ in δ Corg, followed by rapidly enriched values reaching +5.7‰ and –20.1‰, respectively forming a `trough´ according to the Tethyan scheme (Du Vivier et al., 2014; Fig.8), here labelled Pe12B. Thereafter, in the Pe12C sub-segment, there is a plateau of

13 13 δ Ccarb values ranging between +3.8 and +5.1‰, whilst δ Corg values display a shift to less positive values of –26‰ followed by 13C-enriched values, representing a third episode of peak

13 isotope values at –18‰ δ Corg. The top of the prominent Cenomanian–Turonian positive isotope shift ends with Pe12D that corresponds to the recovery segment, characterised by a

13 13 rapid decline to a base level of +1.4‰ in δ Ccarb and –26.3‰ in δ Corg.

The Turonian excursion is comprised in the Pe13 segment and represented in its lower part by five short-term, repetitive patterns in chemostratigraphy all of similar duration and

13 amplitude. They are characterised by short-term inflections from rise to decline in δ Ccarb

13 and δ Corg ratios. The upper part of the Pe13 segment is characterised by a positive excursion.

5 Discussion

5.1 Cenomanian–Turonian chemostratigraphy of Peru: Significance and comparison to other reference curves

In order to provide a robust chemostratigraphic data set that has the potential for supra-

13 regional correlation, patterns in δ Ccarb must be tested for diagenetic features masking or

overprinting palaeoceanographic patterns (Marshall, 1992; Immenhauser et al., 2002; Swart, 2015). In shallow-neritic bulk micrite samples, meteoric waters related to transient subaerial

18 13 exposure stages may introduce O-depleted oxygen, whereas C-depleted soil-zone CO2 can alter carbonate samples (Allan and Matthews, 1977; Christ et al., 2012). We found no clear field evidence for karsting, bleaching of carbonate rocks beneath discontinuities or traces of roots of land plants in the sections studied here. This may imply that Cenomanian to Turonian low-amplitude sea-level fall was insufficient to expose portions of the carbonate ramp studied here. Reasons for this might include rapid basement subsidence, a feature that is supported by the stratigraphically thick successions studied (Fig. 6). The expanded nature of the Peru sections is well exemplified when comparing for example the OAE2 interval from its trough to the end of plateau phase in the Piedra Parada composite section (26 m) with previously described sections from the English Chalk (7.5 m; Jarvis et al., 2011) or Portland #1 Core, USA (2.4 m; Sageman et al., 2006). Similarly expanded shallow marine C–T transition intervals are known for example in Central Mexico (Elrick et al., 2009) and in Tibet (Bomou et al., 2012).

The lack of prolonged subaerial exposure as based on field evidences is confirmed by thin-section observation. Essentially, we did not find any evidence for micro-karst, vadose silt, or meteoric cements in any of the thin sections studied. From this we tentatively conclude that Cretaceous meteoric diagenesis did not significantly reset the

13 chemostratigraphic patterns observed. Furthermore, the regional comparison of δ Ccarb and

13 δ Corg patterns between different sections is considered encouraging. This is because different sections, situated 20 km apart (Figs. 2 and 7), display directly comparable chemostratigraphic patterns, despite different facies distributions in these sections. This supports the concept that the carbon isotope record documented here is principally recording palaeoceanographic patterns and that burial diagenesis did not alter these rocks to a degree that would make the correlation of their geochemical archive impossible. The interpretation based on a regional correlation in Peru can be tested by comparing chemostratigraphic patterns found with such from other basins.

We have correlated the Albian to Turonian chemostratigraphy of Peru with the Tethyan carbon isotope composite age-calibrated curve for the Barremian–Turonian as published by Herrle et al. (2015; Fig. 7). This composite curve is based on sections in the English Chalk (Jarvis et al., 2006), southeast France (Herrle et al., 2004; Gale et al., 2011) and Italy (Erba et al., 1999). The outcome shows convincing similarities as documented in figure 7, except for the onset of OAE2 with is not clear due to a poor exposure. Summing up, the δ13C chemostratigraphy of the Western Platform of Peru is typical for mid-Cretaceous oceanic signatures reported from many basins worldwide (Veizer et al., 1999; Prokoph et al., 2008;

Schulte et al., 2011; Wendler, 2013; Joo and Sageman, 2014) and can be used as a reference curve for further chemostratigraphic work in the subequatorial western Pacific.

Fig. 7. Correlation of the Western Platform δ13C composite curve with the Tethyan carbon isotope composite curve for mid-Cretaceous by Herrle et al. (2015), based on sections of the English Chalk (Jarvis et al., 2006), southeast France (Herrle et al., 2004; Gale et al., 2011) and Italy (Erba et al., 1999). Section denotes the Pe1–Pe13 segments for chemostratigraphic correlation as well as the main OAEs. Age calibration is from Ogg and Hinnov (2012). Biostratigraphic data is taken from Benavides-Caceres (1956), Robert et al. (2009) and Jaillard and Arnaud-Vanneau (1993. Albian chemostratigraphic data in green is taken from Navarro-Ramirez et al. (2015).

5.1.1 Chemostratigraphic features of the Cenomanian isotope excursion and their relation to climate and sea-level change

The early to middle Cenomanian transition in Peru is reflected by the segment Pe10 comprised in the regressive trend of the large-scale sequence 1 characterised by high trophic levels and Pe11 in the transgressive trend of large-scale sequence 2 displaying moderate to high trophic levels upsection (Fig. 5). This pattern of sea-level change is also observed in the Western Interior Basin (USA), the Anglo-Paris Basin and in Morocco (Gale et al., 2002, 2008; Wilmsen, 2003, 2007; Wilmsen et al., 2005; Gertsch et al., 2010b; Giraud et al., 2013; Andrieu et al, 2015). Here, a major sea-level fall, followed by a rapid sea-level rise is recorded near the early/middle Cenomanian boundary. Furthermore in the Anglo-Paris Basin, a marly interval in the lower Cenomanian displays (based on nannofossils) elevated trophic levels due to enhanced runoff related to humid conditions. Grading upsection to oligotrophic conditions at the early/middle Cenomanian boundary, the facies is associated to a positive shift in

13 δ Ccarb labelled MCE Ia (mid-Cenomanian Ia; Giraud et al., 2013) that is also coeval with a

13 major sea-level fall. Further up low mesotrophic conditions in a positive δ Ccarb trend labelled MCE Ib (mid-Cenomanian Ib; Giraud et al., 2013) are recorded. These coincide with the termination of a transgressive system tract (Giraud et al., 2013). A change in trophic conditions has been suggested by Andrieu et al, (2015) as well, reporting that MCEI was synchronous with a turnover in carbonate secreting organisms. Commonly, MCEI is

13 characterised by two positive δ C peaks (MCEIa and MCEIb) separated by a significant through (Cocioni and Galeotti, 2002, Gertsch et al, 2010b). All of this may possibly follow the same controls as those affecting Pe10 and Pe11 sub-segments in Peru. In this context, MCEIa event may correspond to the upper part of the Pe10 segment and the lower part of Pe11 to the MCEIb (Fig. 5), matching quite well the findings of Gertsch et al. (2010b) and Andrieu et al. (2015).

5.1.2 The Cenomanian–Turonian boundary carbon isotope excursion

Fig. 8. Lithological, chemostratigraphy and sequence stratigraphy correlation between the Western Platform, Eastbourne (Jarvis et al., 2011) and the Portland # 1 core (Sageman et al., 2006) sections for the OAE2 interval with indication of the chemostratigraphy markers described by Du Vivier el al. (2014) and Tsikos et al. (2004).

The Cenomanian–Turonian transition in Peru is marked by a pronounced negative trend

13 13 in both δ Ccarb and δ Corg isotope ratios (upper part of Pe11), followed by a positive excursion labelled Pe12. Four sub-segments (Pe12A–D, Figs. 5, 7 and 8) can be defined within the medium-scale sequence 9 (Romirón Formation; Benavides-Caceres, 1956). The pronounced positive carbon isotope excursion of the Pe12 segment is assigned to OAE2 and the sub- segments Pe12A–D are considered equivalents of the datum levels known from the English Chalk (Jarvis et al., 2011) and from the Portland #1 Core, USA (Sageman et al., 2006) for global correlation (Fig. 8). For the OAE2 interval (Fig. 8), the Gun Gardens section at Eastbourne represents an international NW European standard section (e.g., Paul et al., 1999; Jarvis et al., 2006; 2011; Pearce et al., 2009). The Portland # 1 core, drilled close to the Global Stratotype Section and Point (GSSP) of the Cenomanian – Turonian boundary near Pueblo, Colorado (Western Interior Seaway; Sageman et al., 2006) provides a critical framework for global correlation (Meyers et al., 2012; Ma et al., 2014). For OAE2 correlation purposes (Fig. 8), we use the following chemostratigraphy markers as described by Du Vivier el al. (2014) and references therein: `A´ is the onset of OAE2 and its base represents the beginning of the δ13C positive excursion dated as 94.38 Ma (Meyers et al., 2012). Subsequently, there is a trough labelled `B´ dated as 94.23 Ma, followed by a chemostratigraphic plateau where the last relatively enriched δ13C values correspond to the peak `C´, before the values shift back to pre-excursion levels (recovery trend `D´), dated as 93.95 Ma.

Due to poor exposure of some portions in the Piedra Parada section (Figs. 5 and 8), the lowermost portion of the OAE2-related (Pe12A segment) carbon isotope excursion is not accessible. The regional distribution of sandy marls and hardgrounds across the Andes at this stratigraphic level (lower medium-scale sequence 9, Fig. 8) is most likely indicative of a sea- level fall and the shedding of continent-derived clastic material. Across Europe, this interval coincides with a sea-level fall of at least 5–7 m at the base of the Metoicoceras geslinianum Zone (Wilmsen, 2003, 2007; Wilmsen et al., 2005; Voigt et al., 2006, Pearce et al., 2009). Concluding, despite the poor exposures, the sandy marls at the base of the OAE2 are probably indicative of a regional sea-level event.

Assuming that the chemostratigraphy and the biostratigraphy applied here is correct, the top of the onset `A´ in the OAE2 interval may be founded in the chemostratigraphic sub- segment Pe12A in Peru (Fig. 8), and is present in the transgressive system tract of medium- scale sequence 9. At this level, oyster facies 2a documents the establishment of a heterozoan carbonate factory and elevated trophic conditions. A similar oyster facies is also reported from the onset of OAE2 of Wadi El Ghaib (Egypt, Gertsch et al., 2010b) and near Azazoul (Morocco, Gertsch et al., 2010a). The basinal equivalent of this chemostratigraphic segment,

for example bed 3 of the Plenus Marls in the Anglo-Paris Basin and bed 63 at Pueblo Colorado (USA, Elder et al., 1994; Gale et al, 2005, Fig. 8), were interpreted to reflect an initial rise in relative sea level with an amplitude in the order of 10–15 m during the mid M. geslinianum Zone (Voigt et al., 2006). This event may coincide well with the transgressive deposits of medium-scale sequence 9 in Peru.

The trough `B´ of the OAE2 interval could be present in the sub-segment Pe12B (Fig. 8). The arguably equivalent feature in the Portland # 1 core is dated at 94.23 Ma (Meyers et al., 2012). The trough in the sub-segment Pe12B lies in the maximum flooding interval of medium-scale sequence 9, characterised by the echinoid-rich facies 2b indicating normal marine salinity, somewhat elevated nutrient levels and well oxygenated conditions (Gale et al., 2000). In the Anglo-Paris Basin, the sub-segment interval `B´ coincides well with the incursion of cooler water boreal fauna, defining the Plenus Cold Event (Jarvis et al., 2011). Throughout the Plenus Marl Member, a continuous sea-level rise in the order of 10 m has been proposed (Voigt et al., 2006). However, the positioning of the TS in Eastbourne is still a matter of debate (Voigt et al, 2006). In fact, this issue is more complicate, since the sea-level rise took place in Bed 3, below the trough in the Cushmani zone and the MFS is located above in bed 8 at the base of C sub-segment. Hence, if our correlation is correct, the MFS from Eastbourne is not coeval with the one observed in Peru within the limitations of the time framework shown here.

Based on the section shown here (Fig. 8), the flooding pulse across the Western Platform diminished towards the base of the Pe12C plateau sub-segment (Fig. 8). The plateau in the OAE2 interval represents one of the most remarkable and widespread OAE2 features in

13 13 δ Ccarb and δ Corg curves from many basins worldwide (Tsikos et al., 2004). The stratigraphically highest, relatively 13C-enriched value in the plateau is dated at 93.95 Ma in the Portland # 1 core (Meyers et al., 2012). The Pe12C plateau is recorded in the highstand system tract of medium-scale sequence 9 in Peru, characterised by a return of gastropod facies 1c, but lacking oysters. This pattern is in agreement with the interpretation of Voigt et al. (2006) for the Anglo-Paris Basin, suggesting that the plateau of OAE2 forms during a sea- level highstand (Gale et al., 2005; Voigt et al., 2006).

Towards the top of medium-scale sequence 9, i.e., in the recovery sub-segment Pe12D of OAE2 (Fig. 8), facies types recorded display a general dominance of peloids, oncoids and scarcity of macrofaunal elements. This is accompanied by the influx of argillaceous material and reduced carbonate production. According to Gale et al. (2000), towards the top of the OAE interval, there is a significant shift in surface water conditions from high-mesotrophic to distinctly oligotrophic conditions. Following this line of evidence, this may explain the faunal change in Peru from oyster-gastropod dominated facies to oncoid-peloid-dominated facies.

This may indicate a change in trophic conditions in Peru. In Egypt and Morocco (Gertsch et al., 2010a, 2010b), this sub-segment Pe12D corresponds to anoxic conditions and significant biotic stress.

Despite the fact that the Pe12A–D sub-segments in Peru represent well-described features of the OAE2, differences exist as well. In Peru, the range in 13C values is from +1.2 to

13 13 +5.9% in bulk micrite (δ Ccarb) and –26.8 to –17.9‰ in bulk organic matter (δ Corg).

13 Conversely, values range from +2.8 to +5.37% (δ Ccarb) in Eastbourne (e.g., Jarvis et al.,

13 2006) and between –26.3 and –22.1‰ for δ Corg in the Portland # 1 core (Sageman et al., 2006). Possible interpretations of this difference in absolute values include diagenetic differences, or the concept of “aged” neritic platform-top water masses (Lloyd, 1964; Holmden et al., 1998; Saltzman, 2003; Immenhauser et al., 2008; Bomou et al., 2012) of the vast Western Platform in Peru undergoing restriction from vigorous exchange of (and mixing with) the open oceanic water masses. In these settings, the water and the carbon balance are likely decoupled from that of the open oceans potentially leading to either more positive or more negative carbon isotope DIC. Case examples include the recent Florida Bay (Patterson and Walter, 1994), the Pennsylvanian of Asturias, or the Aptian of the Oman platform (see discussion in Immenhauser et al., 2008). Given that the modern world offers no suitable analogue of a similarly extensive epeiric-neritic ramp setting, the interpretation of these data remains difficult at present.

5.1.3 The Turonian post-excursion stage

The Pe13 segment falls into the medium-scale sequence 10 (Figs 5–7). Here, the transgressive deposits and the maximum flooding interval are comprised in the Coñor Formation assigned to the Coilopoceras jenksi Zone (early Turonian; Benavides-Caceres, 1956) and M. nodosoides Zone (Jaillard and Arnaud-Vanneau, 1993). The lower part of the Pe13 segment shows short-term fluctuations that, in the absence of more precise biostratigraphy, are difficult to firmly correlate to other Tethyan sections. The upper part of the Pe13 segment yields a positive excursion that is comprised in the highstand deposits of medium-scale sequence 10 in the Cajamarca Formation (middle Turonian; Benavides- Caceres, 1956). This small positive excursion in Pe13 (Figs 5–7) perhaps matches that of the Round Down Event defined in the early middle Turonian of England (lower C. woollgari Zone, Jarvis et al., 2006).

The deepening trend of medium-scale sequence 10 marks the reduction of neritic carbonate production in the M. nodosoides Zone in the epeiric ramp of Peru and coincides 86

well with a transgressive pulse in the Northern and Southern Tethyan platforms during the M. nodosoides Zone (Philip and Airaud-Crumiere 1991; Drzewiecki and Simo, 1997; Gertsch et al., 2010a, 2010b; Lézin et al., 2012). There, transgression resulted in condensed carbonate facies. This relative sea-level rise is most likely related to the opening of the Equatorial Atlantic Seaway (Trabucho-Alexandre et al., 2010), and/or climatic changes impeding the development of corresponding carbonate platforms (Philip and Airaud-Crumiere 1991; Drzewiecki and Simo, 1997; Gertsch et al., 2010a, 2010b; Lézin et al., 2012; Lebedel et al., 2013).

Shallow marine carbonate production is re-established during the highstand interval in medium-scale sequence 10, i.e. the C. woollgari Zone equivalent of Peru and is characterized by carbonate- and echinoid-rich wacke-mudstones. The oyster-facies is not re-established before the Coniacian-early Santonian (Benavides-Caceres, 1956).

5.2 Western Platform environmental patterns during OAE2

During the pre-isotope excursion interval of the Cenomanian, a heterozoan-dominated carbonate ramp is established and the Mid-Cenomanian Event I is recognized in the Western Platform section of Peru (Figs. 5 and 7). Humid-and high mesotrophic conditions caused enhanced continental influx leading to the deposition of nodular, massive oyster and gastropod facies. At the onset of the Mid-Cenomanian Event I, a maximum sea-level fall resulted in increased argillaceous influx and an overall reduced carbonate production followed by a rapid rise in the early–middle Cenomanian transition. This later rise in sea level coincides with transgressive intervals in the Western Interior Basin (Gale et al., 2002, 2008), in France (Giraud et al., 2013; Andrieu et al., 2015), Morocco (Gertsch et al., 2010b) and in Germany (Wilmsen, 2003, 2007; Wilmsen et al., 2005).

From a chemostratigraphic point of view, the Cenomanian–Turonian boundary carbon isotope excursion recorded in Peru displays important similarities with those of the English Chalk (Jarvis et al., 2011) and the Portland # 1 core (Sageman et al., 2006; Fig. 8). This confirms the wide extent of OAE2 in the very expanded neritic sections in the Western Platform of Peru.

The averaged sedimentation rates (not corrected for compaction) across the OAE2- equivalent units can be computed for the Piedra Parada section (Fig. 8). The stratigraphic thickness from the trough (Pe12B) of the OAE2 to the peak interval `C´ is 28 m and if one adds the strata until the termination of the recovery sub-segment Pe12D, a stratigraphic

thickness of 38 m results. According to the geochronology and astrochronology of the GSSP section (Meyers et al., 2012), the estimates for the duration of the OAE2 ranges between 280 kyr when plotted from the trough `B´ to peak `C´ and 455 kyr when plotted from the trough `B´ until the termination of the OAE2 event. Thus, the approximate mean sedimentation rates of 10 to 8.4 cm/kyr result for the OAE2 interval in Peru. This is in good agreement with previous estimates of epeiric-neritic carbonate depositional settings (e.g., Elrick et al., 2009; Bomou et al., 2013), but also suggests that the OAE2 interval in Peru represents one of the most expanded OAE2 records worldwide directly comparable to those in Mexico and Tibet (Elrick et al., 2009; Bomou et al., 2013).

Jaillard and Arnaud-Vanneau (1993) placed the Cenomanian–Turonian boundary in the lower portion of the Coñor Formation. In the view of these authors, dark-coloured, argillaceous units of the lower Romirón and between the lower and middle Coñor Formation beds are considered evidence of anoxic conditions related to OAE2. This stands in contrast with the chemostratigraphic scheme presented here suggesting that the OAE2 is confined within the Romirón Formation. In contrast, the corresponding interval of the Romirón Formation is rich in benthic fossils indicating well-oxygenated water masses. In other neritic environments across the southern Tethyan realm (Gertsch et al., 2010a, 2010b; Lezin et al., 2012), anoxic/dysoxic conditions are associated to a maximum flooding at the end of the OAE2 interval. Contrary to that, and pending that our chemostratigraphic scheme is correct, the implication of this is that the late Cenomanian to early Turonian Western Platform water masses remained oxic throughout OAE2. Note, similar concepts have been brought forward for sections in Mexico and in Tibet (Elrick et al., 2009; Bomou et al., 2013).

The rather scarce evidence of anoxia/dysoxia in the Western Platform water masses may be associated with to a dominance of shallow water productivity (>50m) in the epeiric-neritic zone rather than upwelling expanding the oxygen minimum zone. We here have presented sedimentological and palaeoecological evidence that humid conditions with high run-off prevailed in Peru during the onset of the OAE2. Conversely, according to Trabucho- Alexandre et al. (2010), strong upwelling occurred along portions of the Pacific and the southern margins of the North Atlantic, leading to the deposition of black shales in deeper water environments. This may have led to the export of nutrient-rich waters into the southern margins of the North Atlantic, enhancing the deposition of organic‐rich sediments in the North Atlantic localities (e.g., Site 1260, Demerara Rise; Site 367, Cape Verde Basin; Tarfaya Basin; Trabucho-Alexandre et al., 2010).

During the geochemical trough interval of OAE2, transgression resulted in the deposition of marly limestones of middle-ramp type sedimentation with normal salinity in Peru. The plateau interval of OAE2 coincides with a sea-level highstand characterised by warm and

normal marine aquafacies. The recovery trend of the OAE2 is marked by a change in trophic conditions, leading the deposition of oncoids and peloids. The early Turonian maximum flooding event during the M. nodosoides Zone is associated with the reduction of neritic carbonate production in the Western Platform (Fig. 5). For the sea-level highstand in the C. woollgari Zone, echinoids facies 2b and micritic carbonate content is dominant.

Concluding, the Cenomanian in Peru was characterised by transient intervals of more humid climate, resulting in an increase in continental run-off from the exposed basement onto the Western Platform. The Cenomanian–Turonian Boundary Event (OAE2) may represent a shift to more seasonally dry conditions, establishing high to moderate trophic resource levels and warm waters towards the top of the OAE2 interval. The Turonian of the Western Platform was characterised by more variable environmental conditions resulting from unstable palaeoceanographic conditions with periods of relatively cool-temperate waters, significantly reduced run-off and low nutrient aquafacies across the Western Platform of Peru.

6 Conclusions

Within Albian–Turonian composite sections from the vast Western Platform in Northern Peru, a total of seven facies associations is found. These are assigned to a shallow to distal ramp environments. In the shallow subtidal inner ramp setting, sandy marls and argillaceous wackestones are the most common facies types whereas oncoid-bearing packstones, grading into gastropod-bearing grainstones, typify higher hydrodynamic level and in some cases, proximal tempestites. The open mid-ramp comprises oyster- and gastropod-bearing float- packstones. These were arguably deposited beneath the effective fair-weather wave base whereas echinoid-bearing wacke- to packstones interpreted as representing transient intervals of more elevated hydrodynamic levels. The outer ramp setting is characterised by planktonic foraminifera-bearing wackestones and mudstones defining deep marine settings below the reach of the effective storm wave base.

By applying detailed carbon isotope chemostratigraphy, we demonstrate that the major mid-Cretaceous OAE’s, as reported for example from the Tethyan realm, are recorded in the epeiric-neritic carbonate successions of Peru. This holds particularly true for the well- expressed patterns assigned to OAE2. The OAE2 carbon isotope excursion is present within the Romirón Formation, but neither organic-rich facies nor anoxic environmental conditions are established. The average sedimentation rate for the OAE2 interval was on the order of ~10 to 8.4 cm/kyr. The implication of this is that these successions represent some of the 89

most expanded OAE2 records worldwide. The δ13C curve in Peru matches well (with the exception of the A_B sub-segments) with globally recognized high-resolution data from coeval successions such as the Tethyan composite curve, or the chemostratigraphy of the English Chalk, or the Portland # 1 core.

During the Albian to early late Cenomanian, a heterozoan-dominated carbonate ramp linked to transient intervals of more humid climate resulted in an increase in continental run-off from the exposed basement onto the Western Platform, recording pre-OAE2 patterns in δ13C such as the Breistoffer and Mid-Cenomanian Event I. At the onset of the Mid- Cenomanian Event I, a maximum sea-level fall is followed by a rapid rise in the early to middle Cenomanian transition.

During the onset of OAE2 (Pe12A), heterozoan carbonate production associated with humid conditions and high run-off is established in the context of a transgressive system tract. In the Pe12B trough of OAE2, a progressive deepening is recorded causing a faunal change to more echinoid-dominated facies. At the plateau of OAE2 (Pe12C), the return to heterozoan carbonate production as part of a highstand system tract is recorded. Finally, the recovery stage of OAE2 (Pe12D) is marked by a change in trophic conditions leading to increased influx of argillaceous facies and reduced carbonate production. Towards the top of the OAE2 interval in Peru a shift to drier and warmer water conditions, and high to moderate trophic resources levels is indicated.

The Turonian of the Western Platform was characterised by more variable environmental conditions resulting from unstable palaeoceanographic conditions with periods of relatively cool-temperate waters and low nutrient aquafacies. This also includes a large-scale maximum flooding event during the M. nodosoides Zone associated with the reduction of neritic carbonate production and a highstand interval in the C. woollgari Zone characterised by micrite-rich carbonate production.

The data shown here are significant as they (i) represent one of the most expanded sections worldwide for this time interval and (ii) because they document that the eastern sub- equatorial Pacific realm recorded global palaeoceanographic patterns as previously described from the Tethyan and Atlantic domain.

Acknowledgements

This project was supported by the Deutsche Forschungsgemeinschaft (DFG, project n° BO-365/2-1) and by the Deutscher Akademischer Austauschdienst (DAAD) through a

scholarship to J.P.N. (PKZ: 91540654). We thank the Geological survey of Peru (INGEMMET) for important logistical support. Analytical work was performed in the isotope laboratories at Bochum and Erlangen-Nuremberg. We acknowledge critical comments of Sedimentary Geology reviewers T. Adatte and B. Sageman, as well as the editorial guidance of Prof. Jones.

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CHAPTER 5

RESPONSE OF WESTERN SOUTH AMERICAN EPEIRIC- NERITIC HETEROZOAN ECOSYSTEM TO OAE 1D AND 2

Navarro-Ramirez J.P., Bodin S., Immenhauser A.

ABSTRACT

The present paper reports on the environmental implications of late Albian– Cenomanian and Cenomanian–Turonian neritic carbonates from the western South American domain in Peru. The focus is on the Oceanic Anoxic Event 1d (late Albian– Cenomanian) and one of the most extreme carbon cycle perturbations of the Phanerozoic, the Oceanic Anoxic Event 2 (late Cenomanian–early Turonian). Thanks to the very expanded and well-exposed sections in central Peru, the OAE 1d and 2 intervals were sampled at high temporal resolution for both bulk micrite and bulk organic matter carbon isotopes. Data shown here document ongoing Albian–Turonian heterozoan carbonate production throughout OAEs 1d and 2 under enhanced deposition of mesotrophic heterozoan fauna. Carbonate producers such as oyster and the endemic large foraminifera Perouvianella peruviana seem better adapted to environmental perturbations associated to OAEs 1d and 2. The carbonate production in Peru was able to keep pace with high rates of basement subsidence, related relative sea-level changes and transient intervals of increased continental run-off under humid climate conditions. Both, the inorganic and the organic δ13C records were probably influenced by sea-level changes. Transgressive deposit towards the maximum flooding interval reveal an initial increase in δ13C values, followed by a decrease in δ13C values, whereas highstand deposit shows positive δ13C excursions.

1 Introduction

The mid-Cretaceous (i.e., the Barremian–Turonian interval, ca. 130–90 Ma; Ogg and Hinnov, 2012) is considered as a time governed by greenhouse conditions due to emission of elevated atmospheric pCO2 levels (e.g., Barron and Washington 1985; Du Vivier et al., 2014; Bodin et al., 2015) produced by a massive pulse of ocean crust production (Arthur et al., 1985; Larson, 1991). These warm conditions were accompanied by reduced temperature gradients from the equator to the poles (e.g., Huber et al. 1995), an accelerated hydrological cycle (e.g., Weissert et al. 1998; Wortmann et al. 2004), and are coeval with a long-term eustatic sea-level rise (e.g., Haq, 2014) resulting in the establishment of vast shallow epicontinental seas (e.g., Cobban and Scott, 1972; Hattin, 1975, 1986; Kauffman, 1977; Elder, 1985; Philip and Airaud-Crumiere 1991; Elder et al., 1994; Mort et al., 2008; Gale et al., 2008; Gertsch et al., 2010a, 2010b; Lézin et al, 2012; Bomou et al., 2013; Navarro-Ramirez et al., 2015a, 2015b). During the mid-Cretaceous, Earth’s oceans were more susceptible to the development of oxygen deficits, leading to the formation of several Oceanic Anoxic Events (OAEs, Schlanger and Jenkyns, 1976; Jenkyns, 2010). The well-known OAEs of global significance for the Cretaceous (e.g., Jenkyns, 2010) include the OAE1a of the early Aptian (Selli event, 120 Ma), the OAE 1b set of the Aptian–Albian ( 111 Ma), the Albian OAE 1c and OAE 1d, ∼and the OAE 2 of the Cenomanian–Turonian (Bonarelli∼ event, 93 Ma).

The OAE1d is known as the Niveau Breistroffer in the Vocontian∼ Basin (France, Bornemann et al., 2005) and the Pialli Level in the Umbria-Marche Basin (Italy, Coccioni, 2001). This event has also been recorded in the Atlantic, the western Tethys, and the Pacific Ocean (Nederbragt et al. 2001; Wilson and Norris, 2001; Erbacher et al., 2001; Strasser et al., 2001; Giraud et al. 2003; Bornemann et al., 2005; Gale et al., 2011; Schröder-Adams et al., 2012; Scott et al., 2013; Melinte-Dobrinescu et al., 2015; Gambacorta et al., 2015). At all these localities, the OAE1d interval is recognized by the deposition of rhythmic organic-rich black shales and a positive carbon isotope excursion with amplitude varying between 0.5 and 1.5‰, suggesting a global perturbation of the carbon cycle. Despite a significant bulk of published data, the palaeogeographic interpretations differ and forcing mechanisms behind such phenomena are still not clearly understood (e.g., Giraud et al., 2003; Bornemann et al., 2005; Gambacorta et al., 2015). Perhaps remarkably, so far no studies have focussed on the impact of OAE1d on epeiric-neritic carbonate settings.

Contrary to OAE1d, OAE2 has been recorded in open, deep marine settings and in shallow water environments (Elrick et al., 2009; Gertsch et al., 2010a, 2010b; Lézin et al., 2012; Bomou et al., 2013; Navarro-Ramirez et al., 2015b). OAE2 represents a short-lived perturbation of ~600 kyr (Sageman et al., 2006; Meyers et al., 2012; Ma et al., 2014) and is 99

frequently recorded by a carbon isotope excursion with an amplitude in order of 2–4% in

13 13 both bulk micrite (δ Ccarb) and organic matter (δ Corg; Pratt, 1985; Pratt et al., 1985; Arthur et al., 1985, 1988; Elder, 1985, 1989; Schlanger et al., 1987; Kennedy and Cobban, 1991; Cobban, 1993; Kauffman, 1995; Dean and Arthur, 1998; Bowman and Bralower, 2005; Gale et al., 2005; Sageman et al., 2006; Tsikos et al., 2004; Friedrich et al., 2006; Jarvis et al., 2006, 2011; Voigt et al., 2008; Du Vivier et al., 2014, 2015; Gambacorta et al., 2015). During OAE2, the highest sea level of the Phanerozoic (e.g., Haq 2014) combined with a climate change towards aridity conditions (e.g., Bomou et al., 2013) caused a long-term decrease in detrital influx (Gertsch et al., 2010a, 2010b; Bomou et al., 2013), a transition from mesotrophic coastal waters to oligotrophic oceanic waters (Gale et al., 2000; Hay, 2008), significant ecological changes in many faunal and floral groups and dysoxic conditions in carbonate platforms along the northern and southern Tethyan (Philip and Airaud-Crumiere 1991; Drzewiecki and Simo, 1997; Sepkoski, 1996; Lamolda et al., 1997; Gertsch et al., 2010a, 2010b; Lezin et al., 2012; Bomou et al., 2013; Kaiho et al., 2014).

In order to interpret the timing of the recovery, the type of ecosystem, and the collapse modalities, of the mid-Cretaceous carbonate interval, detailed stratigraphy, geochemical and sedimentological studies are crucial to identify the impact of the transient carbon cycle perturbations (e.g., OAEs 1d and 2). In terms of carbonate producer, outcrops in Central Peru are characterized by large benthic foraminifera Perouvianella peruviana, an endemic form only known from this setting throughout the late Cenomanian–early Turonian transition (Jaillard and Arnaud-Vanneau, 1993).

In this study (Fig. 1A), we explore carbon isotope data obtained from sections in Central Peru. Specifically, we constrain potential spatial and bathymetric variability in epeiric-neritic DIC patterns with respect to sections in northern Peru (500 km distance, Navarro-Ramirez et al., 2015b). The aims of this paper are: to (i) document and discuss the sedimentological and palaeoecological features of the OAEs 1d and 2 intervals in Central Peru; to (ii) provide a precise chronostratigraphy of the central Peruvian carbonate ramp evolution and to (iii) attribute major steps in ramp evolution to palaeoenvironmental and palaeoceanographic patterns assigned to OAE 1d and 2.

2 Regional tectonic and stratigraphic setting

The information concerning the regional tectonic can be found in chapter 1.6.

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Fig. 1A) Palaeogeographic map of South America during the mid-Cretaceous (modified after Blakey, 2011) indicating the position of what today is the Western Platform of South America (yellow arrow). B) Map of Peru (modified after INGEMMET, 1980) showing location of the Oyon region where the Uchucchacua and Lauricocha sections documented in this study were measured. Moreover, the location of the Piedra Parada and Quebrada Chinchin sections in the Cajamarca region (Navarro-Ramirez et al., 2015b) is shown. C) Geological map of the Oyon region (modified after INGEMMET, 1980) showing the location of the Uchucchacua and Lauricocha sections. 101

In Central Peru (Oyon region; Fig. 1B), the Lower Cretaceous is represented by the Goyllarisquizga Group that encompasses the Chimu, Santa, Carhuaz, and Farrat formations, assigned to the Valanginites broggii Zone at the base and an Aptian age was assumed for the top (Benavides-Caceres, 1956). This group is overlain by Albian carbonate deposits that represent shelf deposition of the Pariahuanca, Chulec, and Pariatambo formations (Benavides-Caceres, 1956). These are overlain by the Jumasha Formation studied in the context of this paper (late middle Albian– early Turonian; Fig. 1C).

The Pariatambo Formation is attributed to the mid-Albian (Prolyelliceras ulrichi and Oxytropidoceras carbonarium zones, Benavides-Caceres, 1956; Robert et al., 2009). It is overlain by the Jumasha Formation, followed by the Celendin Formation of Conician– Santonian age based on ammonite findings (Wilson, 1963; Jaillard, 1986). The Jumasha Formation is ca. 1270 m thick at Uchucchacua, Oyon region, Central Peru (Jailard, 1986). It is made of very massive, thickly-bedded, light yellowish brown to brownish grey limestones, including three marly ledges (Jaillard, 1986; Fig. 2A). The base of the Jumasha Formation was attributed to the late middle Albian due to ammonite associations including Lyelliceras ulrichi Knechtel, Oxytropidoceras douglasi (Benavides-Caceres, 1956; Wilson, 1963; von Hillebrandt, 1970; Jaillad, 1986; Fig. 2B). At the first marly ledge (at 115m of the base of the Jumasha Formation at Uchucchacua), Romani (1982) reported Exogyra costagyra cf olisoponensis and Merlingina cretacea of Cenomanian age, whereas in the third marly ledge, the presence of Sellialveolina sp. indicates a middle to late Cenomanian age (Jaillad, 1986; Fig. 2C). Further south, the third marly ledge of the Jumasha Formation (Jaillard, 1986) contains age assigning microfauna, especially Hedbergela washitensis, H. delrioensis, Globigerinelloides bentonensis and Heterohelix seewashitensis, indicating a middle to late Cenomanian age (Hillebrandt, 1970). Finally at Uchucchacua, at the top of the Jumasha Formation, the presence of Coilopoceras sp. has been reported (Romani, 1982), suggesting a Turonian age for the uppermost part of this formation (Fig. 2C).

3. Methods and materials

3.1 Field work and thin-section microscopy

Two well-exposed sections (Uchucchacua and Lauricocha lake localities; Fig. 1C) were chosen to provide maximum coverage of the Albian–Cenomanian and Cenomanian– Turonian transitions (Figs. 2A–D). In total, ca. 620 m of section have been logged and studied following the method describes in chapter 2.1.

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Fig. 2A): The Jumasha Formation at Uchucchacua, Oyon region, Central Peru, built by massive, thickly- bedded, light-grey limestones. B) Lower part of the Jumasha Formation at Uchucchacua, including the top of the middle Albian and the the lower Cenomanian and the medium-scale sequences 3–6. C) The upper part of the Jumasha Formation at Uchucchacua, including the late Cenomanian to lower Turonian and the medium-scale sequences 9–10. D) The upper part of the Jumasha Formation at Lauricocha, including the late Cenomanian to lower Turonian and the medium-scale sequences 9–10 and the small-scale sequences 1–5.

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Fig. 3. Field images taken at Uchucchacua. A) Nodular limestones of the facies 2a. B) Thickly-bedded limestones of the highstand deposits of medium-scale sequence 9 (MsS9). C) Floatstones of the facies 1d. D) Bioturbation at the level of the sequence boundary 9 (SB9) representing a transgressive surface. E) Thickly-bedded limestones of medium-scale sequence 10. F) Image displaying thin lamination of facies 2c.

3.2 Carbon isotope stratigraphy

13 3.2.1 Bulk micrite data (δ CCarb)

Carbon-isotope analysis were performed on 53 micrite samples (see Appendix E) from the late Albian and 118 from the Cenomanian and Turonian intervals using the analytical methods described in chapter 2.3.1. 104

13 3.2.2 Bulk organic matter data (δ Corg)

Carbon-isotope analyses were performed on 45 bulk organic matter samples (see Appendix F) from the late Albian and 138 from the Cenomanian and Turonian intervals using the analytical methods described in chapter 2.3.2.

4 Data description and interpretation

4.1 Facies associations and depositional environments of the late Albian to early Cenomanian

The data obtained in the context of this study and from previous work point to a mixed, carbonate-siliciclastic ramp with decreasing argillaceous influence towards the west (i.e., Pacific-wards) pinching out against the Albian volcanic arc (Navarro-Ramirez et al., 2015a). The most proximal settings are represented by coastal, deltaic, and tidal flat facies in the eastern part of the platform (present-day eastern Peru and western Brazil, Fig. 1B) were not visited in the context of this project. Field as well as biostratigraphic and thin section data from the late Albian to early Cenomanian lead to the definition of seven facies, representing three main depositional environments.

The facies types recognized at Uchucchacua show important similarities with those previously recognized in the Cajamarca region (Northern Peru, Navarro-Ramirez et al., 2015a, 2015b) and are only briefly described here. At Uchucchacua, the shallow subtidal inner ramp setting comprises three facies types: 1a, 1b and 1c. The most proximal deposits are composed of grey argillaceous mudstone with scarce fauna (facies 1a), whereas in more distal sub-environments, facies 1b is characterized by packstones and argillaceous wackestones, exhibiting a nodular fabric due to an increased argillaceous content. Facies 1c is characterized by grainstone and occasionally floatstone, typically rich in mixed and fragmented shell debris, mainly of oysters and gastropods and echinoderms.

The facies representing the open marine middle ramp setting are mainly characterized by packages of oyster bioherms, interbedded by rare grainstone units yielding various skeletal elements and intraclasts (Facies 2a). Facies 2b exhibits skeletal elements which are locally imbricated, elsewhere bioclasts are arranged in lenses and display a chaotic fabric of non- sorted and fragmented bioclasts. Basin-wards, facies 2c display thinning and are represented 105

by bioturbated mud- to wackestones, less commonly by packstones. Faunal elements consist of gastropods, echinoderms and planktonic foraminifera and of unspecified shell debris. Finally in the outer ramp setting, facies 3a consists of mudstones of dark grey in weathering colour, containing planktonic foraminifera and echinoderms.

4.2 Facies associations and depositional environments of the late Cenomanian to early Turonian

Integrating previous work (e.g., Jaillard, 1986, 1987) with our field data, three main depositional environments, each with a number of standard facies types are here established for the subtidal to outer ramp (Fig. 6). The documentation of these fundamental sedimentological data is important due to the lack of published previous work.

4.2.1 Shallow subtidal inner ramp setting

Facies 1a — Peloid-bearing grainstones: This facies is composed of light grey peloidal grainstones, showing cm to dm-thick lamination (Fig. 3A). Fauna presents some disperse miliolids, ostracods and isolated Perouvianella peruviana (Fig. 4A). Detrital material is scarce but oncoids are present. Given predominance of peloids, disperse miliolids and the low diversity of microfauna, facies 1a is interpreted to evidence a restricted lagoon environment (Gräfe, 2005) with limited water circulation (Masse et al., 2003) and low trophic levels (Hüneke et al., 2001).

Facies 1b — Miliolid-bearing pack-grainstones: This facies is made of bluish grey packstones to grainstones, displaying dm-thick lamination in unis of two to five metres thick (Fig. 3B). Bioturbation features are abundant. The fauna is dominated by miliolids, being associated to P. peruviana, ostracods, numerous dasycladales, and isolated echinoids (Fig. 4B). Non-skeletal peloids and oncoids also occur. Associations of coated grains, skeletal elements such as miliolids, P. peruviana, dasycladales, and echinoids indicate deposition in the nearshore shallow zone (Flügel, 2004; Lézin et al., 2012; Hallock and Glenn, 1986). The predominance of miliolid foraminifera indicates high salinity conditions due to restriction conditions (Hallock and Glenn, 1986). Modern miliolids live in warm, clear water lacking significant fresh-water influx and in some places they are associated with reefal settings (Fang, 2003).

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Fig. 4A): Thin section images documenting grainstones of facies 1a, represented by peloids, miliolids and isolated P. peruviana. B) Miliolids and P. peruviana in facies 1b. C) P. peruviana in facies 1c. D) P. peruviana grainstone to floatstones of facies 1c. E) Diverse fauna pack to grainstones of facies 1d, characterized by a chaotic fabric. F) Crinoid-bearing pack to grainstones of facies 2a. G) Echinoid facies 2b showing fragments of P. peruviana and crinoids. H) Echinoid facies 2b.

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Fig. 5. Regional sequence stratigraphic interpretation and carbon-isotope stratigraphy of the Uchucchacua section for the late Albian. Correlation with the Quebrada Chinchin section (northern Peru, Navarro Ramirez et al., 2015b) and with Tethyan composite curve of Herrle et al. (2015) is shown. Biostratigraphic data of the late Albian have been obtained from oyster fauna associations, taken from Benavides-Caceres (1956) and Jaillard (1986, 1987).

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Fig. 6. Legend for figs. 5, 7 and 8, denoting colour codes for different facies types.

Facies 1c — Large benthic P. peruviana-bearing grain-rud-floatstones (Figs. 4C–D): This facies is characterized by grainstones, rudstones, and floatstones containing numerous P. peruviana. These limestones represent thickening upwards sequences, each one being 20 to 80 m thick, and are predominantly situated in highstand deposits. They occur together with miliolids and peloids (Figs. 4C–D) and occasionally with gastropods and crinoids. According to previous workers, this facies may indicate deposition in the external part of the inner platform (<20 m; Stein et al., 2012) and specifically clear and well-illuminated waters and low trophic levels (Hallock and Glenn, 1986; Lézin et al., 2012).

Facies 1d — Diverse fauna-bearing pack-grainstones: This facies is built by dm to m-thick bedded grey pack to grainstones and occasionally by floatstones (Fig. 3C). At their base, beds present discontinuous horizontal laminae with scours and bioturbation (Fig. 3D). Facies 1d includes mixed and fragmented shell debris mainly consisting of P. peruviana, miliolids, gastropods, crinoids, echinoderms, and planktonic foraminifera (Fig. 4E). Following previous workers, this type of deposits indicate deposition in high-energy, shallow subtidal shoals and storm event beds (Elrick et al., 2009), just above the fair-weather wave base (Lézin et al., 2012) and are linked to transgressive intervals (Flügel, 2004).

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4.2.2 Open marine middle ramp setting

Facies 2a — Crinoid-bearing pack-grainstones: This facies is built by dm to m-thick bedded, bluish-grey packstones and occasionally by grainstones, characterized by large rounded bioclasts composed of crinoidal remains (Fig. 4F). The main benthic foraminifera are P. peruviana and miliolids whose tests are often abraded. Occasionally peloids are present as well as gastropods, echinoids and planktonic foraminifera. The lack of mud suggests a relatively high hydrodynamic level underneath the fair-weather wave base (Stein et al., 2012).

In Central Peru, facies types 2b, 2c and 3a are near-identical to those identified at Piedra Parada (Northern Peru) describe in detail in Navarro-Ramirez et al., 2015b). For the sake of brevity, we only report the most important characteristics. Facies 2b (echinoid-bearing packstones to wackestones) is characterized by cm-thick laminations in units of two to five metres thick (Fig. 3E) and is rich in echinoids and crinoids. These features suggest a middle ramp environment and sediment deposition beneath the fair-weather wave base (Flügel, 2004; Stein et al., 2012). Facies 2c (bryozoans-bearing wackestones to mudstones) is characterized by 5 to 10 cm-thick horizontal (Fig. 3F) and variably sized fragments of bryozoans, echinoids, and planktonic foraminifera. Deposition is assigned to open marine settings, underneath the fair-weather wave base (Amini et al., 2004; Stein et al., 2012). Facies 3a (planktonic foraminifera-bearing mudstones) is characterized by mm to cm-thick lamination, horizontal bedding and consists of dark grey mudstones. The fauna includes planktonic foraminifera and small echinoid fragments. These evidence an open marine setting below the reach of the effective storm wave base (Lézin et al., 2012; Stein et al., 2012).

4.3 Sequence stratigraphic interpretation

The Uchucchacua section is described as the representative case example for the late Albian to early Turonian in the study area (Jaillard, 1986; Jaillard and Arnaud-Vanneau, 1993; Figs. 2A–C). This choice is motivated by the fact that the outcrop conditions are excellent and stratigraphically most complete (Fig. 2B). Moreover, the Lauricocha section is also of high outcrop quality and allows comparing the late Cenomanian to early Turonian transition with respect to the Uchucchacua section (Figs. 2C–D). A basic interpretation of the sequence stratigraphic scheme of Central Peru was presented in Jaillard (1986, 1987) and 110

Jaillard and Arnaud-Vanneau (1993). Navarro-Ramirez et al. (2015b) subdivided the Albian– Turonian of the northern Peru into ten medium-scales sequences (Figs. 5–6). Building on this work, we here subdivide the late Albian to early Cenomanian in four medium-scales sequences (MsSs3–6; Fig. 7), and the late Cenomanian to early Turonian into two medium- scales sequences (MsSs9–10; Fig. 8). This is possible due to the existing biostratigraphic data of the northern and central carbonate ramp of Peru (Benavides-Caceres, 1956; Wilson 1963; Von Hillebrandt, 1970; Romani, 1982; Jaillard 1986, 1987; Jaillard and Arnaud-Vanneau, 1993).

The sequence stratigraphy approach followed here is based on the work of Embry (2009) and Andrieu et al. (2015) for shallow marine environments. This implies that the depositional sequences recognized here are bounded by sequences boundaries (SB), which represent either maximum regressive surfaces or transgressive surfaces (Van Wagoner et al., 1988). These surfaces indicate a shift from a shallowing-upward to a deepening-upward pattern (Embry, 2009; Andrieu et al., 2015). Moreover, a depositional sequence is a depositional cycle that comprises: transgressive deposits (TD) that mark a change to upwards deepening facies (retrograding structures); maximum flooding interval (MFI), a portion of the section that marks the transition from deepening-upward to shallowing-upward patterns; and a highstand deposit (HD), representing shallowing facies grading upwards into more proximal facies (prograding structures).

4.3.1 The late Albian to early Cenomanian of the Jumasha Formation

The lowermost medium-scale sequence 3 (Fig. 7) at Uchucchacua consists of thickening- upwards units of argillaceous, nodular limestone built by facies 2a rich in oysters. This pattern may represent a highstand system tract. These levels were attributed to the late middle Albian due to ammonite associations such as Lyelliceras ulrichi, Oxytropidoceras douglasi (Benavides-Caceres, 1956; Wilson, 1963; von Hillebrandt, 1970; Jaillad, 1986). The overlying medium-scale sequence 4 (Fig. 7) comprises a thickening-upward transition through facies types 1d to facies 2c, perhaps indicating transgressive deposits. This unit is followed by an abrupt deepening trend to outer ramp sedimentation typified by facies 3a with abundant planktonic foraminifera, defining the maximum flooding interval. Further upsection, a thickening-upward succession from facies 2c to 2a, increasingly enriched in quartz grains, oysters and serpulids evidences the highstand.

The overlying medium-scale sequence 5 (52 to 122 m; Fig. 7) is represented at the base by a proximal storm bed (facies 1d), rich in diverse heterozoan fauna. As sea level continued 111

to rise, these relatively shallow deposits passed gradually upwards into progressively deeper depositional environments characterized by brownish marly echinoid- and bryozoa-bearing limestones (facies 2c and 2d). With upward reduced creation of accommodation space, thickly-bedded oyster-bearing limestones (facies 2a), indicate the systems tract. At section metre115 m measured from the base of the Jumasha Formation at Uchucchacua, Romani (1982) reported Exogyra costagyra cf olisoponensis and Merlingina cretacea of Cenomanian age. Echinoids and oyster limestones characterize the last sequence (122 to 156 m; Fig. 7). This interval is perhaps best assigned to the transgressive system of medium-scale sequence 6.

In general, medium-scale sequences 3–5, display a stacking pattern including parasequences topped by marine transgressive surfaces (sequences boundaries 3–5; Fig. 6). This parasequence set may indicate a long-term regressive trend as reflected by successively more restricted marine environments, less fossiliferous and increasingly argillaceous-rich facies until the top of medium-scale sequence 5. Near-identical patterns were recognized at Quebrada Chinchin (Northern Peru, Figs. 6 and 8; Navarro-Ramirez et al., 2015b). There, the equivalent depositional sequences represented a regressive trend of the large-scale sequence 1 (LsS1). The regressive trend during the Albian–Cenomanian has been documented in several studies, indicating a regional event (e.g., Sahagian et al., 1996; Gale et al., 2002; Miller et al., 2003; Immenhauser, 2005; Koch and Brenner, 2009; Scott et al., 2013). Evidence given by these authors suggested glacioeustasy as the driving mechanism (Gale et al., 2002; Miller et al., 2003; Immenhauser, 2005; Bornemann et al., 2008; Koch and Brenner, 2009).

4.3.2 The late Cenomanian to early Turonian of the Jumasha Formation

The late Cenomanian–early Turonian portion at Uchucchacua and Lauricocha is characterized by the third marly ledge of the Jumasha Formation (Jaillard, 1986; Fig. 2A and C). Here, a prominent stratigraphic feature is displayed in the form of sequence boundary 9 (Fig. 3D), that separates medium-scale sequence 9 from sequence 10 (Figs. 6 and 8) and marks a major change in depositional style from a regressive to a major transgression event during the M. nodosoides Zone (Jaillard and Arnaud-Vanneau, 1993; Fig. 8). In the late Cenomanian to early Turonian transition interval, five small-scale sequences (SsS1–5), corresponding to the five depositional sequences Ro1–3 and Co1–2 sensu Jaillard and Arnaud-Vanneau (1993) are here defined. Small-scale sequences 1–4 (SsS1–4; Fig. 8) represent the medium-scale sequence 9 with thicknesses of about 40 to 70 m.

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Fig. 7. Regional sequence stratigraphic interpretation and carbon-isotope stratigraphy of the Uchucchacua and Lauricocha sections for the late Cenomanian. Correlation with the Piedra Parada section (northern Peru, Navarro Ramirez et al., 2015b) is shown. Data points shaded grey correspond to OAE2 based on bio- and chemostratigraphic data. Biostratigraphic data have been obtained from ammonite associations (Benavides- Caceres, 1956 and Jaillard and Arnaud-Vanneau; 1993). 113

At the base of both sections, an abrupt deepening trend from nearshore restricted (facies 1a–1b) to outer ramp facies is observed through small-scale sequence 1 into 2 and may define a transgressive system tract in terms of the medium-scale sequence 9. The maximum of this deepening interval, situated at around section metre 35 in both sequences, is marked by an increase in facies associations 2c and 3a. This is followed by a shallowing-upward trend recorded in the small-scale sequences 3 and 4, indicated by a thickening-upward transition facies 2b upwards 1a. These two sequences represent the regressive trend of the medium- scale sequence 9 and yield abundant P. peruviana. This relative sea-level change is associated to the late Cenomanian–early Turonian and coincides also with the relative sea- level recorded at the Piedra Parada section (Navarro-Ramirez et al., 2015b), as well as a coeval sea-level change present in sections and cores across the Anglo-Paris Basin (Wilmsen et al., 2005; Gale et al., 2005; Voigt et al., 2006). Concluding, despite the limitations of the time framework shown here, the relative sea-level change recorded in the medium-scale sequence 9 is at least regional, and perhaps beyond, in significance.

Small-scale sequence 5 (SsS5, Fig. 8) is at least 90 m thick (but not fully exposed). This sequence includes the early Turonian and the medium-scale sequence 10. Facies 1d through 2c are well-defined in the lower part of both sections (Figs. 4C and 5C), showing an inner- middle ramp environment grading upwards into outer ramp facies 3a rich in planktonic foraminifera (Fig. 5H), i.e. the maximum flooding. The transition from facies 3a into the overlaying thickly-bedded limestones of facies 1a is gradual, marking the highstand system tract. The small-scale sequence 5 yields abundant echinoids, bryozoans, and planktonic foraminifera. The deepening trend of medium-scale sequence 10 marks the reduction of neritic carbonate production in the M. nodosoides Zone at Piedra Parada and coincides well with a transgressive pulse in the Northern and Southern Tethyan platforms during the M. nodosoides Zone (Philip and Airaud-Crumiere 1991; Drzewiecki and Simo, 1997; Gertsch et al., 2010a, 2010b; Lézin et al., 2012). There, transgression resulted in condensed carbonate facies. This relative sea-level rise is most likely related to the opening of the Equatorial Atlantic Seaway (Trabucho-Alexandre et al., 2010), and/or climatic changes impeding the development of corresponding carbonate platforms (Philip and Airaud-Crumiere 1991; Drzewiecki and Simo, 1997; Gertsch et al., 2010a, 2010b; Lézin et al., 2012; Lebedel et al., 2013).

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Fig.8. Regional correlation of the sequence stratigraphic interpretation and carbon-isotope stratigraphy between the northern and central Peru for the early Albian – Turonian. Biostratigraphic information data (Benavides- Caceres, 1956; Robert et al., 2009; Jaillard and Arnaud-Vanneau, 1993), denoting the main OAEs. Late Albian to Cenomanian zones have been obtained from oyster fauna associations. Lower to middle Albian as well as Turonian zones are based on ammonite fauna.

4.4 Carbon-isotope stratigraphy

4.4.1 Late Albian isotope excursion

The late Albian to early Cenomanian bulk micrite and organic carbon isotope stratigraphy at Uchucchacua is documented in figures 6 and 8. At the lowermost part of the

13 13 curve, the δ Ccarb and δ Corg values vary between –0.3 and +2.7‰ and between –27.4 and –

13 25.3‰, followed by a significant positive shift from –0.2 to +3.1‰ (δ Ccarb) and from –27.4

13 to –24.2‰ (δ Corg), showing a prominent positive chemostratigraphic feature. Taking the proposed biostratigraphy and correlation of Jaillard (1987) for central and northern Peru into account, this positive shift is here interpreted to represent the positive excursion recorded in the late Albian in northern Peru (Navarro-Ramirez et al., 2015a). Pending that this chemostratigraphic and biostratigraphic framework is correct, this interval may comprise the OAE1d (Fig. 6; Kennedy et al., 2004; Jarvis et al., 2006; Herrle et al., 2015). From about

13 13 section metre 90 onwards (Fig. 6), C-values decrease reaching 1.3‰ in δ Ccarb and –26.8‰

13 in δ Corg, followed by rapidly enriched values reaching +2.9‰ and –24.2‰, respectively. The top of the prominent positive isotope interval ends with a declining trend to a base level

13 13 of +1.1‰ in δ Ccarb and –26.7‰ in δ Corg. The good match between the Uchucchacua and the Quebrada Chinchin locations (Northern Peru, 500 km away) documents the over- regional significance of these patterns.

4.4.2 Late Cenomanian isotope excursion

In the epeiric-neritic ramp of Peru, the Cenomanian–Turonian isotope excursion linked to OAE2 was previously recorded in the Piedra Parada section (northern Peru; Navarro- Ramirez et al., 2015b; Fig. 7). The δ13C curve matches well with global published high- resolution data for coeval successions such as those reported from the English Chalk (Jarvis et al., 2011) and the Portland # 1 core (Sageman et al., 2006). Based on existing biostratigraphic data in central Peru (von Hillebrandt, 1970; Romani, 1982; Jaillard and Arnaud-Vanneau, 1993) and the chemostratigraphic terminology used for OAE2 interval in

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the Piedra Parada section including terms such as onset, trough, plateau, and recovery (Paul et al., 1999; Tsikos et al., 2004), and datum levels A, B, and C as used in Tethyan and North Atlantic sections (Du Vivier et al., 2014, 2015). We characterize specific sub-intervals of the Cenomanian–Turonian isotope excursion for the Uchucchacua and Lauricocha sections as shown in Figures 7 and 8.

In terms of their δ13C amplitudes and trends, the Uchucchacua and Lauricocha isotope curves are directly comparable. In both sections, the lower part of the prominent carbon isotope excursion related to OAE2, was not covered due to a gap in exposure. However, the OAE2 interval at Uchucchacua displays a significant rise from mean values of +3.7‰ to a

13 13 first peak of +4.6‰ in δ Ccarb and a high ratio –22.6‰ in δ Corg, whereas at Lauricocha the

13 δ Ccarb curve shows depleted values. This specific segment is here interpreted to represent the upper part of the OAE2 onset.

Upsection at both locations, both bulk micrite and organic carbon ratios are characterized

13 13 13 by decreasing C-values of 3.3‰ in δ Ccarb and –25‰ in δ Corg, followed by a rapid trend to 13C-enriched values reaching +4.2‰ and –22.1‰, respectively. This pattern corresponds with a `trough´ according to the Tethyan scheme (Du Vivier et al., 2014; Fig. 8). Thereafter, a

13 13 plateau interval of δ Ccarb and δ Corg values ranging between +1.78 to +3.3‰ and –23 to – 20.9‰ is recorded. The late Cenomanian carbon isotope excursion is topped by a marine hardground termed SB9, perhaps representing a hiatus interval in both settings. At the top of the positive excursion, δ13C values are characterized by a rapid decline to a base level of

13 13 +2.5‰ in δ Ccarb and –25.9‰ in δ Corg.

5 Discussion

5.1 Environmental implications of oyster and P. peruviana mass occurrences

Data showed here document the predominance of oyster-rich mixed siliciclastic- carbonate deposition during the late Albian (Fig. 9A) and Perouvianella-rich carbonate deposition during the late Cenomanian (Fig. 9B).

A key feature of the medium-scale sequences 3–6 recorded at the Uchucchacua location is the presence of oyster facies. This oyster facies is also present at the Quebrada Chinchin location (Figs. 6 and 8), indicating a regional facies belt in Peru (Fig. 9A). The lack of fossil remains of reef constructers (Jaillard, 1987) suggests the establishment of a heterozoan ramp 117

devoid of any morphological feature such as a rimmed morphology. The absence of a rim or barrier may have caused rather uniform facies belts across wide portions of the ramp, where sediment and organic material constituents were transported and re-distributed by storm action and currents (Navarro-Ramirez et al., 2015a). Oyster bioherms may have separated the inner from the mid-zone of the epeiric-neritic ramp (Fig. 9A). The presence of argillaceous material in facies 1a and 1b points to continent-derived sediments probably transported by rivers. According to Jaillard (1987), during the late Albian to early Cenomanian, a diachronic progradation of terrigenous supratidal facies was accentuated, resulting in a NE-oriented deltaic system fed by runoff from the Brazilian continental basement (Fig. 9A). It is also conceivable that nutrient-rich surface currents from the southern margin of the proto-Atlantic reached the western platform via the epeiric-neritic ramp from the north (Trabucho-Alexandre et al., 2010). Nevertheless, the Huarmey-Trough (Atherton and Webb, 1989; Jaillard, 1987; Soler and Bonhomme, 1990), as well as Paracas structural high may have isolated the central Peruvian realm from vigorous exchange and mixing with the water masses of the Pacific Ocean, developing a less stratified water column setting at what is now the Uchucchacua section (central Peru; Fig. 9A). In northern Peru, however, the installation of the Lancones-Celica Basin during the mid-Cretaceous (e.g., Winter et al., 2010) may have formed a seaway for currents between the Pacific and the Peruvian ramp, favouring the mixing of water masses from the proto-Atlantic and Pacific. This may have led to the formation of a stratified water column in the western platform domain represented by the Quebrada Chinchin section.

In the late Cenomanian towards the early Turonian, a key feature of medium-scale sequences 9–10 recorded at Uchucchacua and Lauricocha is the presence of Perouvianella- rich carbonate facies (Fig. 7). Perouvianella peruviana is an endemic form that is only known from these rocks throughout the late Cenomanian–early Turonian transition (Jaillard, 1986 and references therein). During this time, central Peru was isolated from the Pacific and from eastern deltaic influx of the Brazilian continental basement due to the uplift of the Marañon Massif. Evidence for the presence of a morphological barrier is given by the lack of an argillaceous material in these sections (Jaillard, 1986, 1987). It is at least conceivable that this isolated setting resulted in a less stratified water column in what is now central Peru. This later palaeoceanographic feature might have triggered – or at least not inhibited – the mass occurrence of P. peruviana and associated miliolids. In contrast to central Peru, northern Peru (Piedra Parada section) was exposed to nutrient-rich runoff waters and Brazilian shield-derived sediments (Jaillard, 1987) and the influx of proto-Atlantic and Pacific water masses. This confluence of different aquafacies is seen as an argument to propose a stratified water column setting at the Piedra Parada location (Fig. 9B). However, the postulated palaeoceanographic differences between the central and the northern

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Peruvian portions of the western platform are most obvious in inner-middle ramp facies, whilst differences in vertical water mass organization seem less prominent in the outer ramp settings. Evidence for this comes perhaps from the inner-middle ramp deposition of peloids, miliolids, P. Peruviana, and crinoids in central Peru. Ostracods, gastropods, oysters, oncoids, and argillaceous material typify the inner-middle ramp portion in northern Peru. Conversely, the outer ramp domain in both settings is dominated by uniform echinoid facies.

5.2 Carbon cycle disturbances

Shallow water carbonates have been shown to record, under favourable conditions, first order patterns of the global marine carbon isotope signature (Immenhauser et al., 2005; Föllmi et al., 2006; Gertsch et al., 2010a, 2010b; Lezin et al., 2012; Stein et al., 2012; Huck et

13 al., 2013, 2014; Krencker et al., 2014; Andrieu et al., 2015). The global seawater δ CDIC preserved in the rock record are related to changes in the ratio of marine carbonate carbon and organic carbon pools (Scholle and Arthur, 1980; Weissert, 1989; Immenhauser et al., 2003). With reference to the Peruvian sections documented here, we argue that the mid- Cretaceous low-amplitude sea-level fall was insufficient to expose portions of the carbonate ramp studied. Reasons for that might include rapid basement subsidence, a feature that is supported by the stratigraphically thick successions represented for example in the expanded chemostratigraphic through feature at the end of the positive excursion of the OAE2 interval (Ucchucchacua = 150 m and Piedra Parada = 38 m; Fig. 8). Further evidence for rapid basement subsidence comes from the remarkable lack of karst-related features, bleaching, isotope shifts or other patterns assigned to meteoric diagenesis (e.g., Allan and Matthews, 1977; Christ et al., 2012; Huck et al., 2014) beneath discontinuity surfaces recorded in the sections studied. This may imply only weak or no meteoric diagenesis and suggest that the carbonates studied here experienced mainly marine and subsequent burial diagenesis. Judging from the direct comparability of chemostratigraphic sections in Peru with such in many other locations worldwide (e.g., Sageman et al., 2006; Jarvis et al., 2011), we tentatively suggest that most of these carbonates stabilized in the presence of marine pore waters. This is turn implies that at least the chemostratigraphic patterns, but perhaps not absolute isotope values, represent palaeoceanographic features as opposed to diagenetic resetting. The question to which degree these patterns reflect global features versus regional water mass properties merits discussion.

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Fig. 9A). Palaeogeographic interpretation of the western South America platform during the late Albian, characterized by oysters-rich mixed siliciclastic-carbonate deposition. It shows the regional oyster facies deposition, diachronic progradation of the NE deltaic system fed from the Brazilian continental basement, nutrient-rich surface currents from the southern margin of the proto-Atlantic, as well as the Paracas structural high isolating the central Peru from the Pacific Ocean. The northern Peru, in contrast, was open to the mixing with Pacific currents. B) Palaeogeographic interpretation of the western South America platform during the late Cenomanian, characterized by Perouvianella-rich carbonate deposition in central Peru. It shows continuous insulation from the Pacific, and too from eastern deltaic influx of the Brazilian continental basement due to the activation of the Marañon Massif. Interpretations based on Benavides-Caceres (1956), Jaillard (1986, 1987), Atherton and Webb (1989), Soler and Bonhomme (1990), Jaillard et al. (2000), Callot et al. (2008), Winter et al. (2010) and Robert et al. (2009) works.

5.2.1 Impact of environmental changes (OAEs 1d and 2) on the Peruvian neritic carbonate factory

OAE1d is associated to carbon isotope shifts with amplitudes between 0.5 and 1.5‰ in the Tethyan Ocean (e.g., Bornemann et al., 2005; Gambacorta et al., 2015 and references

13 therein). Despite the limitations of the biostratigraphic control used here, the δ Ccarb and

13 δ Corg records at Uchucchacua indicated a 3‰ shift during the late Albian and are higher

13 13 than those recorded at Quebrada Chinchin (1.5‰ δ Ccarb and 2.5‰ for δ Corg). Here the carbonates are, within the limitations of facies characterization, near-identical in both areas (Fig. 6). This implies that facies-control on chemostratigraphy is weak to absent. It is likely

13 that the regional offset in δ Ccarb ratio amplitude may have been induced by differences in the water mass residence time or differential 13C sources (Lloyd, 1964; Holmden et al., 1998; Saltzman et al., 2004; Immenhauser et al., 2008) in the case of the less stratified water column setting at Uchucchacua (central Peru) versus that of the potentially stratified water column setting at Quebrada-Chinchin (northern Peru).

Over-regional oyster facies deposition as recorded in northern and central Peru, however, may to some degree witness the environmental patterns related to OAE1d. Oyster facies in Peru might be analogous with oyster-rich limestones previously reported from Tethyan neritic environments (e.g., Gertsch et al., 2010a, 2010b and references there in). This specific facies type is rarely observed at times of high sea levels and probably points to more humid climate (Gertsch et al., 2010a, 2010b) with all corresponding changes in sediment runoff and type. Increasingly humid climate is reported for the OAE1d (Bornemann et al., 2005). Moreover, the association of oysters with gastropods and echinoids in Peru, which were deposited in the sub-equatorial belt, might suggest mesotrophic conditions (e.g., Andrieu et al., 2015) as reflected by the installation of the Equatorial Humid Belt across Gondwana (Hay and Flögel, 2012).

13 13 For OAE2, the contrast-comparison of δ Ccarb and δ Corg records from the Uchucchacua and Lauricocha sections (situated 36 km apart) agrees with the notion of mainly 121

palaeoceanographic patterns and a weak diagenetic overprint (Fig. 7). However,

13 13 discrepancies exist in the amplitudes of the δ Ccarb and δ Corg chemostratigraphic curves when these are compared with the reference section of Piedra Parada situated approximately

13 13 500 km to the north (Fig. 7). The differential amplitudes of δ Ccarb and δ Corg records are more pronounced at the Piedra Parada location than in sections at the Uchucchacua and Lauricocha sites. Probably, differences in palaeogeographic settings significantly affected or

13 overprinted global patterns at Uchucchacua and Lauricocha (Fig. 9B). Furthermore, δ Ccarb

13 and δ Corg ratios were probably influenced by sea-level changes, which dictate spatial and vertical differences, a feature that is particularly well manifested in the medium-scale

13 sequence 9 (MsS9). Transgressive deposits show relatively heavy δ Ccarb and isotopically

13 depleted δ Corg values and are characterized by P. peruviana, crinoids, and miliolids facies deposition, perhaps at the level of the OAE2 onset `A´ (Fig. 7). The maximum flooding

13 13 interval is characterized by isotopically depleted values in both δ Ccarb and δ Corg records and by echinoids-dominated facies. This interval is likely contemporaneous with the trough

13 feature or the `B´ peak of OAE2. Highstand deposits show relatively heavy δ Corg and

13 depleted δ Ccarb values at the plateau of OAE2, represented by the return of the monospecific P. Peruviana facies. The recovery stage of OAE2 is characterized by peloid-miliolid facies.

The abundance of benthic fossils in these sections supports well oxygenated water masses throughout OAE2 and ongoing carbonate deposition. Two different palaeoecological ecosystems are documented as well. These are characterized by a peculiar biotic pattern showing almost monospecific mesotrophic (P. peruviana and miliolids mass occurrences) assemblages at Uchucchacua and Lauricocha (Figs. 7, 8 and 9B) and a biodiverse mesotrophic (e.g., echinoderms, oysters, gastropods, ostracods) at Piedra Parada.

5.2.2 OAEs 1d and 2: Significance and comparison to other reference curves

Changes in trophic levels are conceived as one of the most important mechanism driving the changes in the geometry, depositional, and carbonate-secreting organisms shaping the evolution of the northern Tethyan platform (e.g., Weissert et al., 1998; Immenhauser et al., 2005; Bodin et al., 2006; Föllmi et al., 2006; Andrieu et al., 2015). These changes coincided well with phases of platform demise linked to major palaeoceanographic change during the Cretaceous (e.g., Föllmi et al., 2006; Immenhauser et al., 2005; Bodin et al., 2006; Andrieu et al., 2015). Other controlling mechanisms include upwelling of oxygen-rich basinal waters forming a link between palaeoceanographic and trophic change (Föllmi et al., 2006). Because upwelling carried waters rich in nutrients and dissolved CO2 onto the platform, culminating of upwelling events episodically lead to anoxic conditions and drowning episodes (Föllmi et 122

al., 2006). Moreover, it was suggested that changes in δ13C correlate with the evolution of the northern Tethyan carbonate producing organisms (e.g., Föllmi et al., 2006; Andrieu et al., 2015). Examples include phases of photozoan deposition associated to positive trends in chemostratigraphy whereas heterozoan facies is often associated with more depleted values (Föllmi et al., 2006). Often, phases of platform drowning show an initial increase in δ13C, followed by a long-term decrease in δ13C values (e.g., Föllmi et al., 2006; Andrieu et al., 2015).

Most of these Tethyan patterns differ from what is observed in the western South American carbonate ramp. Specifically, evidence of trophic changes associated to a palaeoceanographic perturbation is lacking. Conversely, the evolution of the Peruvian carbonate ramp during the mid-Cretaceous was characterized by ongoing Albian–Turonian heterozoan carbonate production (Navarro-Ramirez et al., 2015a, 2015b). With respect to possible upwelling patterns in western South America, this has most likely happened along portions of the Pacific and the southern margins of the North Atlantic (Trabucho-Alexandre et al., 2010), leading to the accumulation of organic-rich black shale in deeper-open marine environments near to the low latitude in the proto-Atlantic and Tethys Ocean (Trabucho- Alexandre et al., 2010). Peruvian δ13C records related to carbonate production does not reflect any changes in δ13C related to distinctive carbonate producer turnover events. Instead, the continuous heterozoan deposition is linked to both positive and negative excursions in carbon isotopes suggesting a significant degree of decoupling of the two parameters.

In the Atlantic, the western Tethys, and the Pacific Ocean, the OAE1d has been associated to high turnover rates, anoxia, increased sea-surface temperatures, and black-shale deposition (e.g., Nederbragt et al. 2001; Wilson and Norris, 2001; Erbacher et al., 2001; Strasser et al., 2001; Giraud et al. 2003; Bornemann et al., 2005; Gale et al., 2011; Schröder- Adams et al., 2012; Scott et al., 2013; Melinte-Dobrinescu et al., 2015; Gambacorta et al., 2015). Evidence for ramp demise, shutdown, drowning, or organic-rich sediment deposition and anoxia related to OAE1d is not observed in Peru. In contrast, phases of enhanced carbonate deposition dominated by high-mesotrophic are recorded. On the level of a tentative working hypothesis, it is here argued that these changes in environmental patters as reflected in heterozoan mesotrophic facies are linked to the impact of the transient carbon cycle perturbation during OAE1d.

During the late M. geslinianum Zone time, a progressive flooding was reported in relation to the OAE2 interval, specifically in its trough feature (or `B´ segment) in the Anglo-Paris Basin (Voigt et al., 2006). According to the authors, this led to a shift in surface water conditions from mesotrophic to distinctly oligotrophic due to the breakdown of the shelf edge frontal system, indicated by significant ecological changes (Gale et al., 2000). Data for the

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OAE2 interval in central and northern Peru presented here, do not signify evidence of morphological or ecological adaptation of the carbonate producers. The endemic P. peruviana fauna constitutes an exceptional case of large benthonic foraminifera that thrive throughout the OAE2 interval without discernible adaptation (Jaillard and Arnaud-Vanneau, 1993). In this context, the late stages of Gondwana breakup during the mid-Cretaceous (Moulin et al., 2010) might have played an important role on the configuration of the palaeogeography in the western South America. Because, this may have mitigated the impact of oceanic deeper upwelled currents against the Peruvian carbonate producer community, precluding carbonate crisis in the ramp. Here, it was thus demonstrated that local signatures in carbon cycling were superimposed on global patterns (Robinson et al., 2004), as shown by the records of OAEs1d and 2 in Peru.

6 Conclusions

The late Albian–Cenomanian neritic sections in Peru document a predominance of oyster-rich mixed siliciclastic-carbonate facies. Regional evidence suggests that this rather uniform facies belt expanded across wide portions of the carbonate ramp in northern and central Peru. The Huarmey trough and the Paracas structural high may have resulted in a less stratified water column setting in the central Peruvian realm. In contrast, in northern Peru evidence for a potentially stratified water column is present. In the late Cenomanian and the early Turonian, the deposition of a very specific large benthic foraminifera facies (P. Peruviana) is a key element that thrive throughout the OAE2 interval without discernible adaptation. During this time, central Peru was isolated from the Pacific and from deltaic influx from the emerged shield of Brazil in the east. This later palaeoceanographic feature might have triggered – or at least not inhibited – the mass occurrence of P. peruviana and associated miliolids.

Despite the complex palaeogeographic setting of the middle Cretaceous carbonate ramp in Peru and no clear evidence for a carbonate crisis OAEs 1d and 2 are recorded in central Peru. The main line of evidence comes from detailed carbon isotope stratigraphy. The mid- Cretaceous ramp evolution was characterized by ongoing Albian–Turonian heterozoan carbonate production with phases of enhancing mesotrophic heterozoan deposition during OAEs 1d and 2. Shallow water masses productivity may have allowed carbonate producers to respond or adapt to carbon disturbances, as well as keeping pace with high subsidence, sea- level changes and transient intervals of increased continental run-off.

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Acknowledgements

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Scott, R.W., Formolo, M., Rush, N., Owens, J.D., Oboh-Ikuenobe, F., 2013. Upper Albian OAE1d event in the Chihuahua Trough, New Mexico, U.S.A. Cretaceous Research 46,136–150.

Soler, P., Bonhomme, M.G., 1990. Relations of magmatic activity to plate dynamics in Central Peru from late Cretaceous to present. Geological Society of America Special Paper 241, p. 173−192.

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Weissert, H., Lini, A., Föllmi, K.B., Kuhn, O., 1998. Correlation of Early Cretaceous carbon isotope stratigraphy and platform drowning events: a possible link? Palaeogeography, Palaeoclimatology, Palaeoecology, 137, 189–203.

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Wilmsen, M., Nielbuhr, B., Hiss, M., 2005. The Cenomanian of northern Germany: facies analysis of a transgressive biosedimentary system. Facies 51, 242–263.

Wilson, J.J., 1963. Cretaceous stratigraphy of Central Andes of Peru. - American Association of Petroleum Geologists Bulletin 47, 1-34.

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CHAPTER 6

SYNTHESIS

The main objectives of this study were i) to investigate the impact of global mid- Cretaceous palaeoenvironmental perturbation on the neritic carbonate platforms in the sub- equatorial eastern Pacific setting of Peru; ii) to understand if the mid-Cretaceous palaeoceanographic patterns found in Peru are comparable to those found in the proto- Atlantic and the Tethyan realm; iii) to examine if global mid-Cretaceous perturbations of the carbon cycle left a significant record in Peru or not, and if yes, how would this be expressed within the epeiric-neritic domain, and iv) to test the hypothesis that the oceanic anoxic event (OAE2) is as scarcely represented in the eastern sub-equatorial Atlantic as suggested from e.g., Site 465 (Hess Rise) by means of land-based sections.

In order deal with the above aims, a multidisciplinary approach has been used in this thesis. Geochemical analyses are coupled with a detailed sedimentological and palaeoecological assessment of carbonate platform successions. The data have been collected in six sections of the northern and central part of the Andes Mountain in Peru. Below I address the above aims and summarize the information gained during this thesis:

1. What was the impact of global mid-Cretaceous palaeoenvironmental perturbation on the neritic carbonate platforms in the sub-equatorial eastern Pacific setting of Peru?

13 13 In Peru, the Albian–Turonian geochemical evidence (manly δ Corg and δ C) points to four oceanic anoxic events associated to carbon cycle disturbances. More precisely, these events are the OAE1b set (early Albian), the OAE1d (Albian–Cenomanian), the MCEI (MCEIa-b; mid-Cenomanian), and the OAE2 (Cenomanian–Turonian). Moreover, other events that might have been associated to a perturbed carbon cycle have been recognized in Peru. These are the mid-Albian (OAE1c?) event, the early Turonian event (M. nodosoides Zone) and the early–middle Turonian event (Round Down Event). These are concomitant

13 with fluctuation of the δ C measured on both bulk micrite and bulk organic matter. Similar features have been recognized in numerus Tethyan and proto-Atlantic hemi-pelagic and 133

pelagic settings and thus, are interpreted as evidence for mid-Cretaceous palaeoenvironmental perturbation (chapters 3, 4 and 5).

From a sedimentological point of view, the siliciclastics-rich ramp of the Lower Cretaceous of Peru (Goyllarisquizga Group) is abruptly replaced by a well-developed heterozoan carbonate ramp in the earliest Albian. On the field, this pattern is translated by an abrupt lithological contrast marked by the complete disappearance of siliciclastic facies. This featured is unique within the OAE1b set and is specifically attributed to the Kilian Level (chapter 3). Within the OAE1b set, in terms of carbonate production, an incipient phase of platform demise is recognized in Peru. Based on previous work from other section, this pattern is assigned to global perturbation of the carbon cycle during the Paquier Level. The isotopic signature of the Leenhardt level in Peru correspond to a major demise of neritic carbonate production, marked in the field by a major discontinuity surface associated to a transgressive surface.

Following the drowning related to the Leenhardt Level, the Peruvian sections reflect condensed sedimentation typified in the field by heterozoan biofacies in dark-grey mudstones. These condensed beds represent maximum flooding stages during the mid- Albian. In the Albian–Cenomanian transition, phases of enhanced carbonate deposition dominated by heterozoan mesotrophic facies are recorded along the Peruvian carbonate ramp. Perhaps, these assumed changes are linked to the impact of the transient carbon cycle perturbation during the OAE1d (chapter 5). Upsection in the early to middle Cenomanian, the MCEI interval is associated to a maximum sea-level fall that resulted in increased argillaceous influx and an overall reduced carbonate production (MCEIa), followed by a rapid rise and heterozoan high mesotrophic deposition (MCEIb; chapter 4).

The Cenomanian–Turonian transition in Peru is marked by a pronounced positive carbon isotope excursion assigned to the OAE2 interval (chapter 4). The top of the onset `A´ in the OAE2 interval is present in a transgressive deposit, characterized by a heterozoan carbonate factory. The trough `B´ of the OAE2 interval lies in a maximum flooding interval, characterized by outer-ramp heterozoan sedimentation. The plateau in the OAE2 interval is recorded in highstand deposit, represented by a return of shallow-ramp heterozoan sedimentation. The recovery sub-segment of OAE2, displays an increased influx of argillaceous facies and reduced carbonate production.

In the Turonian, a reduction in neritic carbonate production coincides well within the M. nodosoides Zone, associated with a regional transgressive pulse. Shallow marine carbonate production is re-established during a highstand interval in the early–middle Turonian and may be associated to the impact of the Round Down Event.

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2. Are the mid-Cretaceous palaeoceanographic patterns found in Peru comparable to those found in the proto-Atlantic and the Tethyan realm?

In the Tethyan realm, episodes of major palaeoceanographic perturbations coincided well with trophic level changes, culminating in phases of platform demise and in marine biota modifications (Föllmi et al., 2006; Immenhauser et al., 2005; Bodin et al., 2006; Andrieu et al., 2015). In basinal and deeper settings localized in the proto-Atlantic and Tethyan realm, OAEs are represented mainly by anoxic conditions and organic-rich sediments deposition (e.g., Schlanger and Jenkyns, 1976; Jenkyns, 2010; Bodin et al., 2015). Most of these Tethyan and proto-Atlantic patterns differ from what is observed in the western South American carbonate ramp. For instance, along the northern Tethyan margin (Helvetic Platform), the onset of the OAE1b set is associated with the final demise of the neritic carbonate factory (Föllmi et al., 2006) giving way to siliciclastic sedimentation and phosphogenesis. Conversely in Peru, the onset of the OAE1b set is linked with the turnover phase from a mainly argillaceous sedimentation to a heterozoan carbonate factory. Furthermore, the middle Albian interval in Peru is assigned to a reduced carbonate production and episodic shutdown of ramp carbonate deposition. Interestingly, this interval seems to be more pronounced in the Pacific sections (including Japan) and in Mexico compared to those in the Tethys domain. This may point to a circum-Pacific pattern not affecting the Tethyan realm.

Moreover, the OAE1d interval was associated to high turnover rates, anoxia, increased sea-surface temperatures, and black-shale deposition in the Atlantic, the western Tethys and the Pacific Ocean settings (e.g., Bornemann et al., 2005; Schröder-Adams et al., 2012; Scott et al., 2013; Gambacorta et al., 2015 and references therein). In contrast to Peru, this interval is characterized by facies of heterozoan mesotrophic fauna deposition, associated to the Equatorial Humid Belt across Gondwana (Hay and Flögel, 2012). On the other hand, in shallow marine environments in France, a change in trophic conditions has been suggested by Andrieu et al. (2015) for the MCEI interval, as reflected by phases of turnover in carbonate secreting organisms. In contrast to the MCEI interval in Peru, its base displays inner-ramp heterozoan sedimentation (MCEIa), followed by successively less fossiliferous and increasingly argillaceous-rich facies in its middle portion. This ends with outer-ramp heterozoan sedimentation (MCEIb). This facies variation may have been caused by a regional sea-level fall associated to MCEI (Gale et al., 2002, Wilmsen et al., 2005; Gertsch et al., 2010; Giraud et al., 2013).

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For the OAE2 interval, a shift in surface water conditions from mesotrophic to distinctly oligotrophic indicated by significant ecological changes, was reported in the Anglo-Paris Basin (Gale et al., 2000). Data for the OAE2 interval presented in this thesis do not signify evidence of morphological or ecological adaptation of the carbonate producers (chapters 4 and 5). The endemic P. peruviana fauna constitutes an exceptional case of large benthonic foraminifera that thrive throughout the OAE2 interval without discernible adaptation (Jaillard and Arnaud-Vanneau, 1993).

3. Did mid-Cretaceous perturbations of the carbon cycle leave a significant record in Peru or not, and if yes, how would this be expressed within the epeiric-neritic domain?

Evidence of trophic changes associated to mid-Cretaceous carbon cycle disturbances on the neritic carbonate platforms in Peru is lacking. Conversely, the heterozoan organisms recorded in the Western Platform in Peru seem better adapted to respond to environmental change related to carbon cycle disturbances. This is reflected by the evolution of the Peruvian carbonate ramp during the mid-Cretaceous (chapter 3), characterized by ongoing Albian– Turonian heterozoan carbonate production (chapter 4) with transient phases of enhanced carbonate deposition during the OAE1d and the OAE2 (chapter 5). The implication of this patter found in Peru, indicates that the ramp water masses remained oxic throughout the Albian–Turonian interval. It is also noticed that the carbonate ecosystem kept pace with the long-term subsidence (Jaillard, 1987), the mid-Cretaceous long-term eustatic sea-level rise (Haq, 2014) and transient intervals of continental run-off due to humid climate, as well. As detailed in chapter 5, it is likely that the late stages of Gondwana breakup during the mid- Cretaceous (Moulin et al., 2010) might have played an important role on the configuration of the palaeogeography in western South America. In consequence, this may have mitigated the influence of oceanic deeper upwelled water masses on the Peruvian carbonate producer community, precluding carbonate ecosystem crisis in the ramp.

4. Specifically, is the OAE2 in Peru as scarcely represented and its organic content stratigraphically expanded (as opposed to being confined to a narrow interval) as suggested from e.g., Site 465, Hess Rise?

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The absence of organic-rich deposits in the Western Platform sedimentary rocks indicates that the production, distribution and preservation of organic matter was either controlled by local signatures in carbon cycling superimposed on global patterns (Robinson et al., 2004) or the deposition (or preservation) of organic matter was preferentially limited in deeper settings (chapter 5). Obviously, given that these are comparably shallow depositional settings, the absence of a basinal sub-oxic water mass has clearly affected deposition and preservation of organic matter. On the other hand, transient periods of more humid conditions during the Albian and Cenomanian affected spatially large source area of continental runoff, delivering freshwater, nutrients, and terrigenous influx of the exposed Brazilian Shield onto the Western Platform. It is possible that this shallow water masses led to the export of nutrient-rich waters into the southern margins of the North Atlantic, enhancing the deposition of organic‐rich sediments in the North Atlantic localities (e.g., Trabucho-Alexandre et al., 2010). Summing up, the OAE2 is equally organic-matter lean in Peru as has been reported from ODP sites in the eastern Pacific. Nevertheless, given that the Peruvian ramp setting is not a direct analogue of the Pacific basinal settings, the comparison of these depositional domains must be dealt with care.

5. Outlook

The establishment of a heterozoan ramp in Peru, devoid of any morphological feature such as a rimmed morphology is strongly suggested in this thesis. This is indicated by sedimentary and organic material flux on this vast ramp most likely transported and re- distributed by storm and tidal current action. Sedimentological analyses tentatively suggested that the water depths of the heterozoan ramp, from the shallow subtidal inner ramp towards the outer ramp setting, were in order of near-sea level to about 50 m bathymetry. However, in this thesis, a quantitative insight into the links between storm action and OAEs was not reached for the sections studied. Characterizing the depth at which the storm wave base was situated during the mid-Cretaceous OAEs is an important but hitherto unresolved aspect. Reasons for this include the fact that links have been made between global warming and storm intensity as suggested by the Pliensbachian–Toarcian OAE (Krencker et al., 2015). It is suggested that future work should aim at (semi- )quantitative estimates (Immenhauser, 2009) of temporal changes in wave base depths in proximal and distal sections in the Peruvian ramp.

In chapter 5, it has been suggested that in central Peru, the heterozoan ramp was isolated from vigorous exchange and mixing with the water masses of the Pacific Ocean, developing a 137

less stratified water column setting during the late Albian. Whereas during the late Cenomanian to early Turonian, the heterozoan ramp in central Peru, was isolated from the Pacific and from eastern deltaic influx of the Brazilian Shield due to the uplift of the Marañon Massif. In my thesis work, I did not succeed to study the lower to middle Cenomanian successions at Uchucchacua to assess the timing of this palaeogeographic change. In order to gain more insight, it is essential to perform geochemical and sedimentological analyses in this early– middle Cenomanian successions in central Peru. This should be done to constrain the potential spatial and bathymetric variability in epeiric-neritic DIC patterns and carbonate platform ecology for the mid-Cenomanian event interval in Peru.

In chapter 5, it is argued that the western Platform of South America was dominated by an Albian to Turonian heterozoan ecosystem. This should be tested by expanding the present data set from the epeiric Western Platform into the Eastern Basin with a special focus on changes in carbonate ecosystem. In order to test, if the Heterozoan ecosystem underwent spatial changes (e.g., from heterozoan to photozoan or to microbe-dominated ecosystem) these data are needed but its compilation has been beyond the time schedule of my thesis. This would allow associating the carbonate ecosystem to specific palaeoceanographic conditions (nutrient levels, sea-surface temperature) within proximal-distal transects.

Moreover, in my thesis, I have argued that the mid-Cretaceous OAE1b, OAE1c, OAE1d, MCEI, and OAE2 were recorded in the Western Platform of Peru, based essentially on

13 13 chemostratigraphic evidence (δ Ccarb and δ Corg, and Sr isotopes). However, the precise timing of these features with regards to the mid-Cretaceous palaeoenvironmental disturbances is not well established yet. This is due to the less-than-ideal biostratigraphic resolution of the Peruvian heterozoan ramp. The age framework could be improved by applying well-preserved oyster shells, in proximal and distal sections from the Peruvian transect presented in chapter 3, in order to perform 87Sr/86Sr isotope analyses. This approach could also help to refine the existing mainly benthic fossil-based chronostratigraphy.

Lastly, the intensification of transient humid conditions during the Albian and Cenomanian is suggested in this thesis. Enhanced nutrient fluxes from the continent accompanied by increased siliciclastic shedding due to accelerated hydrological cycling in the Mesozoic greenhouse have been proposed by various authors (e.g., Föllmi et al., 2006; Immenhauser et al., 2005; Bodin et al., 2006). It is possible to test these hypothesis by exploring source areas of nutrients eastwards across the Maranon Massif into the Eastern Basin, i.e. towards the exposed Brazilian Shield. This would allow for a proximal-distal transect resulting in a chemostratigraphic interpretation of Cretaceous oceanic anoxic events from the vast epeiric Western Platform of Peru into the Eastern Basin. By doing this, a critical

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knowledge bridge of continent-ocean dynamics during intervals of a perturbed carbon cycle in the middle Cretaceous world could be built.

References

Bornemann, A., Pross, J., Reichelt, K., Herrle, J. O., Hemleben, C., Mutterlose, J., 2005. Reconstruction of short term palaeoceanographic changes during the formation of the Late Albian Niveau Breistroffer black shales (Oceanic Anoxic Event 1d, SE France). Journal of the Geological Society, London 162, 623–639.

Gertsch, B., Keller, G., Adatte, T., Berner, Z., Kassab, A.S., Tantawy, A.A.A., El-Sabbagh, A. M., Stueben, D., 2010. Cenomanian–Turonian transition in shallow water sequences of the Sinai, Egypt. International Journal of Earth Sciences (Geol. Rundsch.) 99, 165–182.

Haq, B.U., 2014. Cretaceous eustasy revisited. Global and Planetary Change 113, 44–58.

Hay, W.W., Floegel, S., 2012. New thoughts about the Cretaceous climate and oceans. Earth-Science Reviews 115, 262–272.

Immenhauser, A., 2009. Estimating palaeo-water depth from the physical rock record. Earth-Science Reviews 96, 107–139.

Jaillard, E., Arnaud-Vanneau, A., 1993. The Cenomanian–Turonian transition on the Peruvian margin. Cretaceous Research 14, 585–605.

Jaillard, E., 1987. Sedimentary evolution of an active margin during middle and upper Cretaceous times: the North Peruvian margin from Late Aptian up to Senonian. Geologische Rundschau 76, 677–697.

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Jenkyns, H.C., 2010. Geochemistry of oceanic anoxic events. Geochemistry, Geophysics, Geosystems 11, 1–30.

Krencker, F.N., Bodin, S., Suan, G., Heimhofer, U., Kabiri, L., Immenhauser, A., 2015. Toarcian extreme warmth led to tropical cyclone intensification. Earth and Planetary Science Letters 425, 120–130.

Moulin, M., Aslanian, D., Unternher, P., 2010. A new starting point for the South and Equatorial Atlantic Ocean. Earth-Science Reviews 98, 1–37.

Schlanger, S.O., Jenkyns, H.C., 1976. Cretaceous oceanic anoxic events: causes and consequences. Geologie en Mijnbouw 55, 179–184.

Scott, R.W., Formolo, M., Rush, N., Owens, J.D., Oboh-Ikuenobe, F., 2013. Upper Albian OAE1d event in the Chihuahua Trough, New Mexico, U.S.A. Cretaceous Research 46,136–150.

Wilmsen, M., Nielbuhr, B., Hiss, M., 2005. The Cenomanian of northern Germany: facies analysis of a transgressive biosedimentary system. Facies 51, 242–263.

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141

ANNEX

GEOCHEMICAL DATA

13 13 Annex data includes all the geochemical data (δ Ccarb and δ Corg, and Sr) collected in six sections of the northern and central part of the Andes Mountain in Peru.

Appendix A: Pulluicana and Huameripashga composite section bulk micrite carbon isotope data.

Sample 13 Sample 13 Sample 13 Locality δ C Locality δ C Locality δ C no. carb no. carb no. carb (m) (‰ V-PDB) (m) (‰ V-PDB) (m) (‰ V-PDB) Huameripashg Pulluicana P 20.8 1.5 Pulluicana P 143.3 2.0 H 64.8 -1.9 a P 23.9 1.8 P 144.0 2.0 H 65.0 -1.0

P 26.0 1.8 P 146.0 1.7 H 75.0 -2.6

P 34.3 0.3 P 150.0 1.8 H 75.3 -2.9

P 40.0 0.1 P 152.0 1.5 H 79.5 -0.9

P 45.0 0.8 P 154.0 1.5 H 84.0 -0.2

P 46.0 0.6 P 156.0 2.4 H 95.0 0.7

P 47.0 0.8 P 158.0 2.0 H 100.0 0.4

P 48.0 0.1 P 159.0 1.8 H 105.0 0.9

P 51.0 0.4 P 160.0 1.9 H 110.0 0.2

P 54.0 1.1 P 160.5 1.7 H 115.0 0.0

P 55.4 1.7 P 161.0 1.7 H 120.0 0.2

P 58.0 1.4 P 161.2 1.6 H 124.0 0.4

P 59.9 1.7 P 164.0 1.9 H 125.0 0.6

P 63.3 1.8 P 170.0 0.9 H 130.0 0.9

P 67.0 1.6 P 172.0 1.1 H 135.0 0.6

P 75.0 1.3 Huameripashga H 1.0 1.1 H 140.0 0.9

P 78.1 1.7 H 5.0 0.8 H 150.0 1.2

P 81.0 2.0 H 6.7 2.8 H 155.0 1.3

P 83.0 1.8 H 13.0 1.1 H 160.0 1.3

P 84.2 1.7 H 14.4 0.6 H 165.0 1.8

P 85.0 1.6 H 14.5 1.3 H 170.0 1.6

P 90.0 1.3 H 15.2 -0.2 H 174.8 2.0

P 95.0 1.2. H 17.0 -0.2 H 184.0 1.7

P 97.0 1.2 H 21.0 0.7 H 190.0 1.7

P 100.2 1.3 H 25.0 0.5 H 194.0 1.5

P 102.0 1.6 H 30.0 0.3 H 200.5 1.6

P 104.0 1.7 H 31.0 -0.5 H 205.0 1.6

P 106.0 1.5 H 33.0 0.0 H 210.0 1.7

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P 118.5 1.3 H 35.0 1.0 H 215.0 1.6

P 120.0 1.4 H 37.0 0.5 H 220.0 1.3

P 122.0 1.1 H 39.0 -0.3 H 225.0 1.6

P 124.0 0.5 H 40.0 0.7 H 228.5 2.0

P 126.0 2.0 H 40.0 0.1 H 235.0 1.9

P 130.0 2.2 H 45.0 0.9 H 240.0 1.6

P 134.0 1.5 H 46.0 0.1 H 244.5 1.6

P 136.0 1.8 H 50.0 1.2 H 250.0 1.4

P 138.0 1.7 H 51.5 0.2 H 255.0 1.8

P 140.0 1.8 H 55.0 0.4 H 260.0 1.9

P 142.0 1.8 H 60.0 -1.2 H 263.0 1.5

Appendix B: Pulluicana and Huameripashga composite section bulk organic matter isotope data.

Sample 13 13 Sample 13 Locality δ C Locality Sample no. δ C Locality δ C no. org org no. org (‰ V- (‰ V- (‰ V- (m) (m) (m) PDB) PDB) PDB)

Pulluicana P 0.5 -24.6 Pulluicana P 120.0 -26.8 Huameripashga H 33.0 -26.4

P 1.8 -24.2 P 122.0 -27.1 H 37.0 -27.0

P 5.7 -26.2 P 124.0 -26.9 H 42.0 -26.7

P 6.9 -25.5 P 128.0 -26.6 H 50.0 -24.1

P 20.8 -25.9 P 130.0 -27.1 H 50.7 -26.4

P 26.0 -25.9 P 132.0 -27.2 H 56.5 -27.5

P 31.0 -25.5 P 134.0 -26.8 H 60.0 -26.7

P 33.3 -24.9 P 136.0 -27.1 H 63.0 -27.6

P 35.0 -25.7 P 142.0 -26.5 H 70.0 -27.2

P 36.0 -25.8 P 144.0 -26.9 H 75.3 -27.3

P 38.0 -26.0 P 150.0 -26.7 H 84.0 -26.6

P 42.0 -26.0 P 156.0 -26.9 H 99.0 -26.9

P 46.0 -26.1 P 161.0 -26.7 H 124.0 -27.4

P 47.0 -26.3 P 162.0 -28.0 H 135.0 -27.0

P 51.0 -26.4 P 164.0 -27.3 H 145.0 -27.2

P 54.0 -26.9 P 166.0 -27.0 H 160.0 -27.5

P 59.9 -25.6 P 168.0 -27.0 H 165.0 -27.2

P 62.0 -26.7 P 170.0 -26.4 H 170.0 -27.2

P 72.0 -26.1 P 172.0 -26.5 H 174.8 -27.4 Huameripashga P 75.0 -25.9 H 1.0 -27.2 H 180.0 -27.0

P 81.0 -26.3 H 5.0 -27.0 H 184.0 -25.5

P 85.0 -26 H 7.0 -26.5 H 194.0 -27.2

P 92.5 -27.1 H 9.0 -27.0 H 205.0 -27.6

P 97.0 -26.8 H 13.0 -26.5 H 215.0 -26.8

P 100.2 -26.5 H 17.5 -27.0 H 228.5 -27.1

P 104.0 -26.3 H 21.0 -27.2 H 240.0 -27.2

P 108.0 -26.7 H 25.0 -26.6 H 255.0 -26.9

P 118.5 -27 H 30.0 -26.1 H 263.0 -27.0

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Appendix C: Piedra Parada and Quebrada Chinchin composite section bulk micrite carbon isotope data

Sample 13 13 13 13 13 δ C Sample no. δ C Sample no. δ C Sample no. δ C Sample no. δ C no. carb carb carb carb carb (m) (‰ V-PDB) (m) (‰ V-PDB) (m) (‰ V-PDB) (m) (‰ V-PDB) (m) (‰ V-PDB)

(Piedra Parada) T 33,00 4,20 T 137,0 2,76 I -57,0 2,54 I -244,0 0,47

T 0,00 1,02 T 34,10 4,37 T 139,0 2,86 I -59,0 2,63 I -246,0 0,00

T 0,20 3,36 T 35,00 4,26 T 141,0 2,91 I -65,0 2,74 I -252,0 0,31

T 1,20 3,82 T 36,00 3,97 T 143,0 2,86 I -68,0 2,04 I -254,0 0,44

T 1,30 4,06 T 37,30 3,60 T 145,0 2,74 I -70,0 1,94 I -258,0 0,35

T 1,50 3,95 T 39,00 3,17 T 147,0 2,55 I -76,0 2,98 I -264,0 -0,04

T 1,70 3,98 T 40,00 2,88 T 149,0 2,37 I -78,0 2,32 I -266,0 0,05

T 2,00 4,09 T 41,00 2,86 T 151,0 2,33 I -82,0 2,97 I -266,0 -0,05

T 2,20 4,23 T 42,00 2,81 T 159,0 2,76 I -87,0 2,83 I -268,0 -0,05

T 3,00 4,43 T 42,40 2,50 T 163,0 3,01 I -93,0 2,40 I -270,0 2,65

T 3,80 4,63 T 45,00 2,16 T 167,0 3,17 I -95,0 2,64 I -303,0 -0,25

T 5,00 4,76 T 47,50 2,03 T 171,0 3,17 I -97,0 2,91 I -305,0 -0,80

T 5,50 4,73 T 49,00 2,19 T 175,0 3,14 I -100,0 2,76 I -310,0 -1,13

T 5,80 4,68 T 51,50 2,43 T 179,0 3,11 I -104,0 2,12 I -318,0 -0,02

T 6,00 4,52 T 53,00 2,57 T 179,0 3,04 I -107,0 3,07 I -320,0 -0,51

T 6,50 4,27 T 55,00 2,62 T 183,0 2,98 I -111,0 2,55 I -325,0 0,12

T 7,00 4,23 T 58,00 2,54 T 187,0 3,00 I -116,0 3,21 I -327,0 0,39

T 7,20 4,47 T 61,00 2,33 T 191,0 3,09 I -118,0 3,33 I -327,0 0,31

T 8,00 4,83 T 64,00 2,06 T 195,0 3,03 I -120,0 2,95 I -330,0 0,12

T 8,50 5,25 T 66,00 1,90 T 199,0 2,83 I -124,0 2,97 I -336,0 0,79

T 8,90 5,37 T 68,00 1,94 T 203,0 2,73 I -127,0 2,66 I -338,0 0,54

T 9,20 5,46 T 69,00 2,14 (Quebrada Chinchin) I -133,0 3,15 I -342,0 -0,05

T 9,40 5,45 T 71,00 2,43 I 0,0 1,50 I -140,0 1,59 I -346,0 1,54

T 12,80 5,38 T 73,00 2,67 I -1,0 1,61 I -142,0 0,20 I -348,0 1,38

T 13,00 5,36 T 75,50 2,72 I -2,0 1,47 I -145,0 1,61 I -350,0 1,43

T 13,20 5,46 T 79,00 2,54 I -3,0 1,89 I -150,0 1,25 I -383,0 -1,79

T 14,00 5,63 T 81,00 2,34 I -4,0 0,54 I -152,0 1,61 I -387,0 0,85

T 14,50 5,71 T 84,00 2,29 I -10,0 1,01 I -154,0 -1,59 I -391,5 0,84

T 15,00 5,61 T 87,00 2,41 I -13,0 0,93 I -162,5 1,58 I -394,0 -1,50

T 16,00 5,36 T 90,00 2,56 I -16,0 0,83 I -166,0 2,17 I -397,0 1,25

T 17,10 5,15 T 93,00 2,63 I -16,0 0,76 I -168,5 1,25 I -402,0 0,55

T 18,00 4,96 T 95,00 2,63 I -18,0 1,49 I -172,0 0,65 I -404,0 0,73

T 18,50 4,74 T 97,00 2,64 I -20,0 2,02 I -179,0 2,36 I -406,0 1,11

T 19,20 4,57 T 102,00 2,61 I -22,0 2,58 I -182,0 1,11 I -410,0 1,49

T 20,00 4,47 T 104,00 2,44 I -24,0 2,56 I -184,0 1,80 I -416,0 1,17

T 21,30 4,44 T 106,00 2,18 I -26,0 2,04 I -188,0 1,50 I -422,0 1,55

T 22,00 4,47 T 108,40 2,05 I -30,0 2,90 I -190,0 2,03 I -428,0 1,53

T 22,70 4,62 T 112,00 2,13 I -34,0 3,14 I -196,0 1,14 I -430,0 1,15

T 24,00 4,85 T 114,00 2,32 I -36,0 1,25 I -200,0 1,72 I -434,0 1,89

T 24,50 4,92 T 115,50 2,46 I -38,0 1,75 I -207,0 0,98 I -440,0 1,22

T 25,00 4,72 T 117,50 2,56 I -40,0 1,28 I -214,0 1,73 I -446,0 1,43

T 27,00 4,48 T 119,50 2,59 I -43,5 3,23 I -220,0 0,77 I -450,0 1,34

T 28,30 4,43 T 122,00 2,46 I -45,5 2,88 I -224,0 0,98 I -453,0 1,26

T 29,30 4,52 T 125,00 2,28 I -49,0 3,00 I -226,0 0,53 I -460,0 1,58 144

T 30,00 4,39 T 128,00 2,27 I -50,0 3,13 I -230,0 -0,31 I -464,0 1,14

T 30,80 4,04 T 131,00 2,44 I -52,0 2,57 I -236,0 -0,20 I -466,0 1,40

T 31,70 3,96 T 134,00 2,64 I -56,0 2,44 I -238,0 0,04 I -489,0 1,16

I -492,0 1,16 I -570,00 0,58 I -633,0 0,10 I -687,0 -0,04 I -762,00 1,32

I -494,0 0,93 I -574,00 0,51 I -646,0 -2,62 I -689,0 -0,15 I -770,00 1,53

I -500,0 1,20 I -580,00 0,17 I -650,0 -1,38 I -702,0 0,64 I -774,00 -0,15

I -508,0 1,21 I -588,00 0,00 I -652,0 0,99 I -704,0 0,24 I -778,00 0,63

I -514,0 1,07 I -588,00 0,15 I -655,0 -1,71 I -710,0 0,75 I -782,00 0,73

I -518,0 1,37 I -594,00 0,54 I -656,0 0,88 I -718,0 -0,73 I -789,00 0,88

I -523,0 1,20 I -596,00 0,27 I -660,0 -4,47 I -726,0 0,36 I -792,00 2,20

I -530,0 1,62 I -600,00 0,10 I -667,0 1,41 I -730,0 0,25 I -794,00 1,88

I -536,0 1,24 I -602,00 -0,13 I -669,0 -4,01 I -732,0 0,77 I -797,00 0,94

I -542,0 1,46 I -606,00 0,40 I -672,0 -2,36 I -734,0 -2,22

I -548,0 1,11 I -614,00 0,17 I -677,0 -1,66 I -742,0 0,46

I -555,0 0,65 I -619,00 0,34 I -679,0 -1,38 I -745,0 1,25

I -562,0 1,75 I -627,00 0,22 I -683,0 -0,15 I -750,0 0,72

I -566,0 0,76 I -633,00 0,09 I -685,0 -0,02 I -756,0 1,63

Appendix D: Piedra Parada and Quebrada Chinchin composite section bulk organic matter isotope data

Sample 13 Sample 13 13 13 13 δ C δ C Sample no. δ C Sample no. δ C Sample no. δ C no. org no. org org org org (m) (‰ V-PDB) (m) (‰ V-PDB) (m) (‰ V-PDB) (m) (‰ V-PDB) (m) (‰ V-PDB)

(Piedra Parada) 22,70 -25,47 51,50 -25,61 108,50 -25,92 -20,00 -26,19

0,00 -19,45 24,00 -25,25 53,00 -25,58 110,00 -25,64 -30,00 -25,52

1,70 -24,83 24,50 -25,79 55,00 -26,33 112,00 -26,79 -40,00 -26,31

2,20 -25,10 25,00 -21,90 58,00 -25,53 114,00 -25,95 -49,00 -23,99

3,00 -22,62 25,60 -25,17 61,00 -25,41 115,50 -25,83 -59,00 -25,16

5,00 -24,32 26,40 -26,79 64,00 -25,98 117,50 -26,46 -68,00 -26,64

6,00 -25,65 27,00 -23,35 65,00 -26,59 119,50 -26,41 -78,00 -26,46

7,00 -26,38 27,00 -23,30 66,00 -25,14 122,00 -25,93 -87,00 -23,88

8,00 -24,22 27,70 -27,18 66,50 -24,98 125,00 -24,37 -95,00 -21,94

8,50 -25,36 28,50 -21,03 67,00 -25,00 128,00 -23,34 -107,00 -24,76

8,50 -25,44 30,00 -23,07 68,00 -26,25 131,00 -25,47 -116,00 -25,21

8,90 -23,67 30,80 -21,81 69,00 -26,48 134,00 -26,26 -127,00 -26,44

9,40 -23,97 31,70 -20,78 71,00 -26,25 137,00 -23,27 -133,00 -23,16

10,00 -22,47 33,00 -21,11 73,00 -26,02 139,00 -24,88 -145,00 -26,08

12,80 -23,72 33,00 -24,05 75,50 -25,36 141,00 -26,12 -154,00 -26,79

13,00 -23,20 34,10 -22,72 77,00 -24,02 143,00 -27,63 -156,00 -22,55

14,00 -22,89 35,00 -17,95 79,00 -24,69 145,00 -26,57 -184,00 -24,25

15,00 -25,47 36,00 -22,20 81,00 -23,49 147,00 -24,43 -200,00 -24,27

16,00 -21,68 37,30 -24,32 84,00 -25,71 149,00 -25,92 -207,00 -24,14

17,10 -23,06 38,00 -25,38 87,00 -26,11 151,00 -21,98 -214,00 -24,27

18,00 -21,51 39,00 -22,46 90,00 -25,61 153,00 -26,11 -226,00 -23,87

145

18,00 -27,13 40,00 -24,78 93,00 -26,20 155,00 -25,38 -238,00 -26,19

18,50 -20,11 41,00 -24,75 95,00 -26,17 (Quebrada Chinchin) -246,00 -25,89

19,20 -24,51 42,00 -23,37 97,00 -25,91 0,00 -26,59 -258,00 -26,84

20,00 -21,29 42,40 -27,48 100,00 -24,44 -4,00 -26,40 -270,00 -24,42

20,50 -26,21 45,00 -25,55 102,00 -23,88 -10,00 -26,20 -310,00 -25,11

21,30 -24,84 47,50 -26,25 104,00 -24,47 -13,00 -27,15 -318,00 -24,79

22,00 -21,89 49,00 -25,30 106,00 -25,36 -18,00 -26,93 -325,00 -27,16

AppendixE: Uchucchacua and Lauricocha sections bulk micrite carbon isotope data 13 13 13 13 13 Sample no. δ Ccarb Sample no. δ Ccarb Sample no. δ Ccarb Sample no. δ Ccarb Sample no. δ Ccarb (‰ V- (‰ V- (‰ V- (‰ V- (‰ V- (m) (m) (m) (m) (m) PDB) PDB) PDB) PDB) PDB)

Uchucchacua_Late Albian UE 137,50 1,64 UC 24,00 3,31 Late Cenomanian LB 106,00 2,68

UE -22,00 0,11 UE 140,00 1,08 UC 26,00 2,47 Lauricocha_A_B LB 114,00 2,23

UE -20,00 -0,29 UE 142,00 1,61 UC 30,00 2,72 LA 4,00 3,91 LB 122,00 2,32

UE -15,00 2,69 UE 146,00 2,00 UC 38,00 3,18 LA 8,00 3,75 LB 128,00 2,37

UE -12,00 0,70 UE 150,00 1,91 UC 40,00 3,16 LA 12,00 3,63 LB 134,00 2,47

UE -8,00 0,78 UE 152,00 1,11 UC 44,00 2,49 LA 16,00 4,02 LB 138,00 2,47

UE -4,00 -0,23 UE 156,00 1,91 UC 48,00 2,50 LA 16,00 3,86 LB 142,00 2,71

UE -2,00 1,08 Late Cenomanian UC 50,00 2,14 LA 20,00 3,47 LB 146,00 2,14

UE 0,00 0,81 Uchucchacua_B_C_D UC 56,00 3,24 LA 24,00 3,45 LB 150,00 4,04

UE 4,00 0,89 UB 0,00 2,64 UC 60,00 4,17 LA 28,00 3,35 LB 156,00 3,71

UE 8,00 0,65 UB 4,00 2,60 UC 64,00 3,78 LA 32,00 3,47

UE 8,00 2,79 UB 8,00 2,45 UC 68,00 3,35 LA 32,00 3,40

UE 12,00 0,54 UB 12,00 2,58 UC 72,00 3,38 LA 40,00 3,41

UE 16,00 0,51 UB 16,00 2,45 UC 76,00 3,18 LA 43,00 3,89

UE 20,00 -0,25 UB 19,00 2,02 UC 78,00 3,24 LA 46,00 3,05

UE 24,00 0,19 UB 20,00 2,57 UC 82,00 3,28 LA 50,00 4,71

UE 28,00 2,45 UB 24,00 3,02 U.D 1,00 3,65 LA 54,00 4,39

UE 32,00 1,34 UB 26,73 2,74 U.D 2,50 3,05 LA 58,00 3,81

UE 36,00 2,24 UB 30,00 2,66 U.D 6,00 3,31 LA 62,00 3,42

UE 40,00 1,28 UB 36,00 2,82 U.D 12,70 3,97 LA 66,00 2,61

UE 42,00 2,18 UB 41,00 1,09 U.D 14,00 3,57 LA 70,00 3,54

UE 46,00 2,53 UB 44,00 3,25 U.D 18,00 4,56 LA 74,00 3,09

UE 48,00 2,83 UB 46,00 2,96 U.D 20,00 3,68 LA 78,00 3,24

UE 52,00 2,24 UB 50,00 3,40 U.D 24,00 4,18 LA 82,00 3,38

UE 56,00 2,99 UB 54,00 2,92 U.D 28,00 2,92 LA 86,00 3,33

UE 60,00 2,68 UB 58,00 3,08 1,89 LA 98,00 4,00

UE 64,00 2,87 UB 62,00 3,08 0,54 LA 100,00 3,38

UE 68,00 2,92 UB 66,00 2,78 1,01 LB 0,00 3,29

UE 74,00 2,83 UB 69,00 3,58 0,93 LB 4,00 2,99

UE 78,00 2,65 UB 72,00 2,35 0,83 LB 8,00 3,45

UE 78,00 2,74 UB 76,00 3,03 0,76 LB 12,00 3,77

UE 80,00 3,12 UB 80,00 3,12 1,49 LB 16,00 3,88

UE 84,00 2,82 UB 84,00 3,47 2,02 LB 20,00 3,70

UE 88,00 2,76 UB 86,00 1,62 2,58 LB 28,00 3,31

146

UE 90,00 1,63 UB 92,00 2,09 2,56 LB 32,00 2,93

UE 94,00 1,97 UB 96,00 3,09 2,04 LB 38,00 3,77

UE 96,00 1,33 UB 96,00 3,07 2,90 LB 46,00 3,24

UE 98,00 2,05 UB 100,00 3,29 3,14 LB 50,00 3,62

UE 102,00 1,59 UB 104,00 2,99 1,25 LB 54,00 3,58

UE 106,00 2,28 UB 108,00 3,21 1,75 LB 58,00 3,67

UE 110,00 2,11 UB 110,00 2,65 1,28 LB 64,00 3,30

UE 114,00 2,79 UC 0,00 3,26 3,23 LB 72,00 3,34

UE 118,00 2,88 UC 4,00 1,92 2,88 LB 78,00 2,90

UE 122,00 2,69 UC 8,00 2,94 3,00 LB 84,00 3,38

UE 126,00 2,12 UC 12,00 3,09 3,13 LB 86,00 3,44

UE 130,00 1,91 UC 16,00 3,09 2,57 LB 90,00 3,09

UE 134,00 1,69 UC 20,00 1,88 2,44 LB 98,00 3,37

Appendix F: Uchucchacua and Lauricocha sections bulk organic matter isotope data 13 13 13 13 13 Sample no. δ Corg Sample no. δ Corg Sample no. δ Corg Sample no. δ Corg Sample no. δ Corg (‰ V- (‰ V- (‰ V- (‰ V- (‰ V- (m) (m) (m) (m) (m) PDB) PDB) PDB) PDB) PDB)

Late Albian Late Cenomanian UC 48,00 -23,24 Late Cenomanian LB 64,00 -22,65

Uchucchacua_E Uchucchacua_B_C_D UC 50,00 -22,66 Lauricocha_A_B LB 66,00 -23,12 - UE 20,00 -26,55 UB 0,00 -25,73 UC 54,00 -20,68 LA 4,00 -23,75 LB 72,00 -22,22 - UE 15,00 -24,11 UB 4,00 -25,09 UC 56,00 -22,05 LA 8,00 -22,90 LB 78,00 -22,94 - UE 12,00 -25,66 UB 8,00 -25,90 UC 60,00 -21,18 LA 12,00 -22,11 LB 80,00 -23,24

UE -8,00 -26,52 UB 12,00 -25,59 UC 62,00 -22,09 LA 16,00 -23,95 LB 84,00 -23,21

UE -4,00 -27,36 UB 16,00 -25,73 UC 64,00 -22,39 LA 20,00 -24,06 LB 86,00 -25,11

UE -2,00 -26,90 UB 20,00 -26,13 UC 68,00 -23,34 LA 24,00 -23,30 LB 90,00 -23,77

UE 0,00 -25,34 UB 24,00 -25,20 UC 70,00 -24,39 LA 28,00 -24,52 LB 94,00 -23,95

UE 4,00 -26,42 UB 26,70 -23,80 UC 72,00 -23,90 LA 32,00 -24,65 LB 98,00 -23,56

UE 8,00 -26,01 UB 30,00 -23,96 UC 76,00 -24,10 LA 36,00 -24,74 LB 102,00 -24,93

UE 12,00 -26,81 UB 32,00 -24,99 UC 78,00 -24,16 LA 40,00 -24,41 LB 106,00 -25,70

UE 16,00 -26,96 UB 36,00 -24,09 UC 82,00 -24,94 LA 43,00 -24,65 LB 110,00 -26,30

UE 24,00 -27,40 UB 41,00 -24,41 U.D 1,00 -22,10 LA 46,00 -23,76 LB 114,00 -25,51

UE 26,00 -27,02 UB 44,00 -23,09 U.D 4,00 -22,82 LA 50,00 -23,59 LB 118,00 -25,57

UE 28,00 -25,77 UB 46,00 -24,08 U.D 6,00 -23,62 LA 54,00 -23,63 LB 122,00 -25,87

UE 32,00 -25,89 UB 48,00 -23,81 U.D 12,70 -22,37 LA 58,00 -22,42 LB 126,00 -25,76

UE 36,00 -24,88 UB 50,00 -22,81 U.D 14,00 -22,05 LA 62,00 -22,09 LB 128,00 -25,60

UE 40,00 -26,49 UB 54,00 -22,36 U.D 18,00 -23,37 LA 66,00 -22,68 LB 134,00 -26,11

UE 42,00 -26,68 UB 58,00 -23,04 U.D 20,00 -22,45 LA 70,00 -21,89 LB 138,00 -26,02

UE 46,00 -26,48 UB 62,00 -21,89 U.D 24,00 -22,83 LA 74,00 -21,98 LB 142,00 -25,55

UE 48,00 -24,84 UB 66,00 -23,38 U.D 25,00 -22,32 LA 78,00 -23,62 LB 146,00 -23,45

UE 52,00 -26,24 UB 69,00 -23,54 U.D 28,00 -22,60 LA 82,00 -23,38 LB 150,00 -22,77

UE 56,00 -24,43 UB 72,00 -24,52 LA 86,00 -22,49 LB 152,00 -22,74

UE 60,00 -25,79 UB 76,00 -23,13 LA 90,00 -24,02 LB 156,00 -24,33

UE 64,00 -26,11 UB 78,00 -23,60 LA 94,00 -22,98

UE 68,00 -24,67 UB 80,00 -22,43 LA 98,00 -22,13

UE 72,0 -26,21 UB 84,00 -23,62 LB 0,0 -21,82

147

UE 74,0 -26,78 UB 86,00 -20,86 LB 2,0 -22,61

UE 80,0 -26,66 UB 88,00 -22,16 LB 4,0 -20,97

UE 84,0 -25,10 UB 92,00 -21,79 LB 8,0 -22,21

UE 88,0 -26,75 UB 96,00 -22,27 LB 12,0 -22,42

UE 90,0 -23,07 UB 98,00 -22,39 LB 16,0 -24,03

UE 94,0 -23,99 UB 100,00 -22,59 LB 20,0 -22,68

UE 96,0 -23,39 UC 0,00 -23,03 LB 24,0 -24,20

UE 98,0 -25,17 UC 4,00 -22,57 LB 28,0 -22,98

UE 102,0 -23,66 UC 8,00 -23,55 LB 32,0 -22,03

UE 110,0 -24,16 UC 12,00 -22,65 LB 36,0 -23,56

UE 114,0 -24,95 UC 16,00 -22,42 LB 38,0 -22,97

UE 118,0 -24,62 UC 20,00 -22,77 LB 42,0 -23,20

UE 122,0 -24,99 UC 24,00 -22,70 LB 46,0 -22,70

UE 126,0 -25,87 UC 26,00 -22,28 LB 50,0 -23,17

UE 130,0 -23,24 UC 30,00 -21,63 LB 50,0 -25,26

UE 134,0 -25,46 UC 33,00 -24,00 LB 54,0 -22,46

UE 142,0 -26,75 UC 38,00 -22,69 LB 56,0 -25,14

UE 150,0 -26,37 UC 40,00 -22,72 LB 58,0 -22,53

UE 156,0 -26,08 UC 44,00 -22,81 LB 60,0 -23,55

148

CURRICULUM VITAE

Juan-Pablo Navarro-Ramirez

Current address: Beisingstraße 3, 44807 Bochum, Germany

44807 BOCHUM, Germany Date of birth: 12th March 1985

Phone: +4915214828958 Familial situation: Married

Citizenship: Peruvian Driving license: submitted

E-mail: [email protected] Phone: +4915214828958

GEOLOGIST

Skills and Expertise

• Palaeoceonography

• Palaeoclimatology

• Palaeoecology

• Sedimentary geology (carbonate and clastic)

• Field Sedimentology

• Sequence stratigraphy

• Geomorphology

• Microfacies analyses

• Time series analyses

• Geochemistry (carbon-oxygen stable isotope, strontium radiogenic)

• Field mapping and Structural cross-section (scales: 1:10,000 to 1:100,000)

• Basin analyses

• Regional geology

• Aerial photographic and satellite image interpretation

149

Education

11.2012 – 02.2016 Ph.D. Earth Sciences, University of Bochum (with Prof. Dr. Adrian Immenhauser and Dr. Stephane Bodin)

Research thesis: “Impact of Middle Cretaceous climatic change on sub-equatorial Pacific settings (Andean Basin, Peru): A comparison with the Tethyan Realm.

2008-2010 Diploma in Geology Engineering in the National University of Piura, Peru

Thesis: “Palaeogeographic Evolution of the upper Cenozoic Northwestern Peru”

2002-2007 B. Sc. in Geology Engineering, National University of Piura, Peru

Computer Office automation: Microsoft Office Suite

Desktop Publishing: Adobe Suite

Drawing: Illustrator

Languages Spanish: native language

English: fluent

German: fluent

Experience:

Laboratory:

2013-2015 Preparing bulk rock samples for stable isotope (carbon and oxygen) analyses, Strontium radiogenic and sample decarbonatation at the Ruhr University of Bochum (continuously).

2008-2012 Representation of 3-d structures on an arbitrary 2-d horizontal geological map, hand and digital drawings of geological maps, structural cross section interpretation, aerial photographic and satellite image interpretation, sedimentology and basin analysis at the Peruvian Geological Survey

2007 Teamwork by the geological evaluation of hydrocarbons exploratory areas in the Ene Basin (sub-andean eastern basin). Samples descriptions, logging sheets and structural cross section drawing at the Peruvian Geological Survey

2005-2006 Palaeontological cleaning and preparation of fossils at the Palaeontological institute of the National University of Piura, Peru

150

Field work:

2013-2014 Two campaigns in the Andean Basin of Peru to study the impact of climate changes on the carbonate factories during the Albian to Turonian interval. 40 days of field work done from February 2013 to September 2014. Skills developed: sedimentology, geometry of sedimentary bodies, sequence stratigraphy, tidal deposits, palaeoecology. Supervisor: Prof. Dr. Adrian Immenhauser and Dr. S. Bodin

2007-2012 Peruvian Geological Survey (INGEMMET): 21 campaigns in the Peruvian geological settings (eg.,Talara, Sechura, Lancones, Casma, Ene and Chota basins) to study the geology, geodynamic and economic implications of the North and Central Andes of Peru. 505 days of field work done from April 2007 to March 2012. Skills developed: Field mapping, structural cross-sections, field sedimentology, stratigraphic correlations, structural geology, basin analyses. Supervisor: Prof. Dr. Victor Carlotto, Dr. Mirian Mamani

2006 Two campaigns to study the Evolution of the continental ecosystems during the Plio-Pleistocene in Northern Peru” organized by the Peruvian Geological Survey. 28 days of field work done from March to November 2008. Skills developed: fluvial-marine field sedimentology. Supervisor: Prof. Dr. Jean Noel Martinez and Prof. Dr. Victor Carlotto,

Teaching:

2013, 2014 S.Bodin student Assistant for the sedimentary petrography practical course (2 hours per week, Winter semesters).

Sept. 2013 B.Sc. students fieldwork supervision in Cajamarca, Peru (3 days)

June 2013 B.Sc. students fieldwork supervision in Pötzen quarry, Germany (1 day)

Professional:

2008 to Mar. 2012 Peruvian Geological Survey (INGEMMET)

Others:

Apr.-Jul 2013 Participation in the theoretical - practical course on Sedimentary carbonate systems organized by Prof. Dr. Adrian Immenhauser

151

Sept.-Dec 2013 Participation in the theoretical - practical course on sequence stratigraphy organized by Prof. Dr. Adrian Immenhauser

Sept. 2010 Participation at the theoretical - Practical course on Applications of Ichnology in facies analyses and sequence stratigraphy organized by the 18th ISC, Mendoza

Participation in the theoretical - Practical course on Tectonics and sedimentation in Petroliferous Basins organized by the 18th ISC, Mendoza

Achievements:

Mar 2012 to Mar 2016 A DAAD scholarship for a research program at the Ruhr University of Bochum

2009 Awarded by the Geological Society of Peru as the second best thesis 2009 in Peru.

Interest and Personal Experience

Sport Football in club from 1995 to now

Hobbies Climbing, fitness,

Publication List

Articles:

Navarro-Ramirez J.P., Bodin S., Immenhauser A., 2015. Ongoing Cenomanian – Turonian heterozoan carbonate production in the neritic settings of Peru. DOI: 10.1016/j.sedgeo.2015.10.011

Navarro-Ramirez J.P., Bodin S., Immenhauser A. (in preparation). Response of western South American epeiric-neritic heterozoan ecosystem to OAE 1d and 2

Geological maps:

Jaimes, F., Santos, A., Navarro-Ramirez, J.P., Russe, E., Bellido, F., 2012. Mapa geológico del cuadrángulo de Las Lomas (10-c), hoja 10-c, cuadrantes I al IV. Updating of the National geological Maps, INGEMMET, Lima, Peru.

Jaimes, F., Navarro-Ramirez, J.P., Santos, A., Russe, (from 2008 to 2011). Geological Maps of Olmos (12-d-II, III), Pomahuaca (12-e-II, III), Huancabamba (11-e-II, III), Las Lomas (10- c-I, II, III, IV), Incahuasi (13-e-I, II, III, IV), Chiclayo (14d- I, II, III, IV), Chota (14-f-I) and San Marcos (15-g-I). Updating of the National geological Maps, INGEMMET, Lima, Peru. 152

Cueva, E., Carlotto, V.S., Acosta, H., Navarro-Ramirez, J.P., Alván, A.A. & Russe, E. (2010) - Mapa geológico del cuadrángulo de Casma (19-g), hoja 19-g, cuadrante III. Updating of the National geological Maps, INGEMMET, Lima, Peru.

Poster Presentations:

Navarro-Ramirez, J.P., Bodin, S., Heimhofer, U. & Immenhauser, A., Impact of middle Cretaceous (Albian to Turonian) climatic changes in the eastern sub-equatorial marginal Pacific region, EGU General Assembly 2015.

Navarro-Ramirez, J.P. & Alvan A., Sedimentary facies of Plio-Pleistocene successions between Chira and Piura rivers (Northwestern Peru), IAS 2010.

Navarro-Ramirez, J.P., Jaimes F., Santos A. & Alvan A., La Formación Sávila en el noroeste de Perú: equivalente occidental de la Formación Condorsinga, nuevos registros estratigráficos del Toarciano, XV CPG, Cuzco-Peru, 2010.

Navarro-Ramirez, J.P. & Martinez J.-N., Facies sedimentaria en la Serie Plio-Pleistocena entre los ríos Chira y Piura (noroeste del Perú): reconstrucción paleogeográfica, XV CPG, Cuzco-Peru, 2010.

Oral Presentations:

Navarro-Ramirez, J.P., Martinez J.-N., Análisis sedimentológico-paleogeográfico del Plio- Pleistoceno de Chusís (Provincia de Sechura, Departamento de Piura), XIV CPG, Lima-Peru, 2008.

Navarro-Ramirez, J.P., Martinez J.-N., Aspectos sedimentológicos del valle del río Chira (Noroeste del Perú, XIV CPG, Lima-Peru, 2008.

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