Cretaceous Oceanic Red Beds (Corbs): Different Time Scales and Models of Origin

Earth-Science Reviews 115 (2012) 217–248

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Cretaceous oceanic red beds (CORBs): Different time scales and models of origin

Xiumian Hu a,⁎, Robert W. Scott b, Yuanfeng Cai a, Chengshan Wang c, Mihaela C. Melinte-Dobrinescu d a State Key Laboratory of Mineral Deposits Research, School of Earth Sciences and Engineering, Nanjing University, Nanjing 210093, China b University of Tulsa, 149 West Ridge Road, Cleveland, OK 74020, USA c Research Center for Tibetan Plateau Geology, China University of Geosciences, Beijing 100083, China d National Institute of Marine Geology and Geo-ecology, RO-024053 Bucharest, Romania article info abstract

Article history: The Cretaceous oceanic red bed (CORB) is a newly opened window on global oceanic and climate changes Received 22 January 2012 during the Cretaceous greenhouse world. As a result of the International Geoscience Programmes 463, 494 Accepted 16 September 2012 and 555 (2002–2010), CORBs have been documented in many places by numerous publications. The princi- Available online 5 October 2012 ple goal of this paper is to summarize scientific advances on CORBs including chronostratigraphy, sedimentology, mineralogy, elemental and isotopic geochemistry, and their relationship to oceanic anoxic events Keywords: (OAEs), palaeoclimate and palaeoceanography. Cretaceous oceanic red beds (CORBs) fi Palaeoceanography We propose a new geochemical classi cation of the CORBs using CaO, Al2O3 and SiO2 values, which lithologically Palaeoclimate refer to marly, clayey, and cherty CORBs respectively. Detailed mineralogical studies indicate that hematite, Oceanic anoxic event (OAE) goethite and Mn2+-bearing calcite are the minerals imparting the red color of CORBs. Hematite clusters of several to tens of nanometers in the calcite structure are the main cause of the red coloring of limestones, and the Mn2+-bearing calcite gives additional red color. Goethite was thought to form originally with hematite, and was subsequently transformed to hematite during late diagenesis. Chronostratigraphic data allow the distinction of two groups of CORBs by their durations. Short-term CORBs are generally less than 1 myr in duration, and seem to be on the scale of Milankovitch cycles. During the deposition of Cretaceous reddish intervals from ODP cores 1049 and 1050, low primary productivity and relatively high surface temperature resulted in low organic carbon flux into the sediments which reduced oxygen demand and produced oxidizing early diagenetic conditions. In such an oxic environment, iron oxides formed imparting the reddish color. The long-term CORBs' depositional events lasted longer than 4 myr, and may be a possible consequence of the OAEs. Enhanced amounts of organic carbon and pyrite burial during

and after the OAEs would have resulted in a large and abrupt fall in atmospheric CO2 concentration, which probably induced signiﬁcant global climatic cooling during and after the OAEs. Global cooling would have enhanced formation of cold deep water, increasing its oxidizing capacity due to the greater content of dissolved oxygen and would promote formation of oceanic red beds. Sedimentological, mineralogical and geochemical data indicate that CORBs were deposited under highly oxic, oligotrophic conditions probably at a low sedimentation rate. The Cretaceous red and white limestones from Italy have similar compositions of terrestrial input-sensitive elements (Al, Ti, K, Mg, Rb, Zr), higher contents

of Fe2O3, and depleted redox-sensitive elements (V, Cr, Ni, and U) and micronutrient elements Cu, Zn, indicating similar provenance sources but red limestones were deposited under more oxic conditions at the sediment–water interface than white limestones. The Cretaceous red shales such as those from the North Atlantic and Tibet have similar mineralogy and geochemistry as the Late Cenozoic red clays in the Pacific Ocean and the environment where both are formed was well-oxidizing at a very low sedimentation rate. We compiled seventeen published stratigraphic examples of Phanerozoic oceanic red beds including the Late Cenozoic red clays in the Pacific. Different hypotheses explain the origin of red pigmentation of limestones and shales including (1) detrital origin of iron derived from continental weathering; (2)iron-bacterial mediation at the time of sedimentation; and (3) iron oxidation in oligotrophic, highly oxic environment. Additional research on Phanerozoic oceanic red beds is needed in order to better document their origin and palaeoceanographic and palaeoclimatic significance. © 2012 Elsevier B.V. All rights reserved.

⁎ Corresponding author. Tel.: +86 25 8359 3002; fax: +86 25 8368 6016. E-mail address: [email protected] (X. Hu).

Contents

1. Introduction ...... 218 2. Geochemical classiﬁcation of CORBs ...... 219 2.1. Ca-CORBs ...... 219 2.2. Al-CORBs ...... 221 2.3. Si-CORBs ...... 221 3. Chronostratigraphy of CORBs ...... 222 3.1. Long-term CORB deposition ...... 222 3.2. Short-term ORB deposition ...... 222 4. Mineralogical composition and color origin of CORBs ...... 225 4.1. Mineral composition of CORBs ...... 225 4.2. Iron oxides in CORBs ...... 226 4.2.1. Quantitative analysis of iron oxide with XRD ...... 226 4.2.2. DRS data of the CORBs ...... 226 4.2.3. Quantitative analysis of iron oxide concentrations in ODP Hole 1049C ...... 228 4.3. Origins of the red color of CORBs ...... 229 4.3.1. Determination of staining minerals ...... 229 4.3.2. Red coloring mechanism of red limestone ...... 229 4.3.3. Red coloring mechanism of red shale ...... 232 4.4. Summary ...... 232 5. Elemental geochemistry ...... 232 5.1. Data source and method ...... 232 5.2. Ca-CORBs (red carbonates) ...... 232 5.2.1. Major elements ...... 232 5.2.2. Trace elements ...... 235 5.2.3. Rare earth elements ...... 235 5.3. Al-CORBs (red shales) ...... 235 5.3.1. Major elements ...... 235 5.3.2. Trace elements ...... 235 5.3.3. Rare earth elements ...... 237 5.4. Summary ...... 237 6. OAE–ORB transition and stable carbon isotopes ...... 237 6.1. OAE 1a–ORB1 transition ...... 237 6.1.1. Gorgo a Cerbara section, Italy ...... 237 6.1.2. Yenicesihlar section, Turkey ...... 237 6.2. OAE 2–ORB 9 transition ...... 237 6.2.1. Umbria–Marche, Italy ...... 237 6.2.2. Buchberg section, Austria ...... 238 6.2.3. Onset of the ORB9: an Early to Middle Turonian global event? ...... 238 6.3. Summary on the stable carbon isotopes during the OAE–ORBtransition...... 238 6.4. Model for the OAE–ORB transitions ...... 238 7. Orbital-cycle control CORBs on Blake Nose, North Atlantic ...... 239 7.1. Data ...... 239 7.2. Model ...... 239 8. CORBs and the Cretaceous ocean and climate ...... 239 9. Perspective: Phanerozoic oceanic red beds ...... 240 9.1. Distribution of oceanic red beds in the Phanerozoic Eon ...... 240 9.1.1. Red shales ...... 240 9.1.2. Red carbonates ...... 241 9.2. Late Cenozoic red clays ...... 243 9.3. Origin of Phanerozoic marine red beds ...... 243 9.3.1. Detrital origin of the iron derived from continental weathering ...... 243 9.3.2. Iron-bacterial mediation at the time of sedimentation ...... 243 9.3.3. Iron oxidation in oligotrophic, highly oxic environment ...... 243 10. Conclusion ...... 244 Acknowledgments ...... 244 References ...... 245

1. Introduction Cretaceous marine red beds in the Chuangde section, southern Tibet that were deposited in the Eastern Tethys Ocean. Following studies of Cretaceous marine red beds have been known for at least 150 years, well-known Cretaceous chalks (see Hay, 2008) and black shales since Štur (1860) and Gűmbel (1861) first described them from the (Schlanger and Jenkyns, 1976; Jenkyns, 1980), Cretaceous oceanic red Pűchov beds in the Carpathians and the Nierental beds in the Eastern beds (CORBs) became a study focus since the 2000s within the frame- Alps. A few biostratigraphic and sedimentological studies followed, par- work of the International Geoscience Programme (IGCP) projects. ticularly in European countries. During late 1990s Chinese colleagues Transnational interdisciplinary research was conducted on their global (Chengshan Wang and Xiumian Hu) discovered and studied the Upper distribution, correlation, and significance of the oxidation of these X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 219 deposits for palaeoceanographic reconstructions. Their relationships and the Romanian Carpathians (Melinte-Dobrinescu and Roban, 2011). were contrasted to the distinctly different, interbedded mid-Cretaceous Neuhuber and Wagreich (2011) discussed the inorganic geochemistry black shales. of the CORBs. The term CORB was born in 2001 discussions in Lhasa, Tibet and Here we propose a new geochemical classification of the CORBs,

Chengdu when Luba Jansa, Massimo Sarti, Chengshan Wang and Xiumian using CaO, Al2O3 and SiO2 values and the mineral composition and el- Hu meet together to revise IGCP463 proposal. The term CORB was first emental geochemistry of different types of CORBs. The relationships published by Wang et al. (2004, 2005) and Hu et al. (2005),andbecame with OAEs and orbitally controlled oceanic red beds (ORBs) are con- widely used in global correlation. CORBs are reddish to pinkish to brown- sidered. Finally, we discuss the geological history and different origins ish sedimentary rocks (generally limestone, marl, shale, and/or chert), of of oceanic red beds. Cretaceous age, deposited in pelagic marine environments (Hu et al., 2005; Scott et al., 2009). A number of publications were published both 2. Geochemical classification of CORBs in international and regional journals during the past years as achieve- ments of IGCP projects, including the IGCP 463 “Upper Cretaceous Ocean- The CORBs are generally classified into three end members mainly ic Red Beds: Response to Ocean/Climate Global Change” (2002–2006 led based on types of sedimentary petrology (Wagreich et al., 2009; by Chengshan Wang, Massimo Sarti, Robert Scott, and Luba Jansa), the Wang et al., 2009): (1) red hemipelagic and pelagic carbonates IGCP 494 “Dysoxic to oxic change in ocean sedimentation during (Fig. 1A and B); (2) deep-water red claystones deposited below the mid-Cretaceous: a study of the Tethyan realm” (Young Scientists Project, carbonate compensation depth (CCD) (Fig. 1C); and (3) red cherts 2003–2006 led by Xiumian Hu, Krzysztof Bak, Jens Wendler and Natalyia and radiolarites deposited below the CCD (Fig. 1D). In this paper,

Tur), and the IGCP555 “Rapid Environmental/Climate Change in the Cre- we use geochemical CaO, Al2O3 and SiO2 values to classify the taceous Greenhouse World: Ocean-Land Interactions” (2007–2010 led by CORBs based on the compiled data of CORBs (Table 1, Figs. 2 and 3). Chengshan Wang, Robert Scott, Hugh Jenkyns, Michael Wagreich, Wil- liam Hay and Yuri D. Zakharov). Papers related to CORBs were mainly 2.1. Ca-CORBs published in four special volumes: two in the journal “Cretaceous Re- search” (Wan and Sarti, 2005; Hu et al., 2012a), one in the journal “Sed- The CaCO3 contents of the Ca-CORBs vary from 48% (Brandenberg, imentary Geology” (Wagreich et al., 2011), and another in the Society of Austria) to 50% (Chuangde, Tibet) and up to the highest at 95% (Vispi,

Sedimentary Geology (SEPM) Special Publication 91 (Hu et al., 2009b). Italy) (Table 1). The Al2O3 values range from 1% (Vispi, Italy) to 10% Wang et al. (2009) summarized the results of the IGCP studies, including (Brandenberg, Austria). The SiO2 value varies from 3% (Vispi, Italy) to temporal–spatial distribution, biostratigraphy, and sedimentology. Sev- 30% (Chuangde, Tibet), except for one sample from DSDP 417D which eral papers summarized CORBs in several regions including the Northern has 58% SiO2, and can be classiﬁed as siliceous limestone. Ca-CORBs Atlantic (Jansa and Hu, 2009), the Eastern Alps (Wagreich et al., 2009) are mainly pelagic limestones like the Italian Scaglia Rossa. Biogenous

A B

C D

Fig. 1. (A) Field photos of red limestone from the Scaglia Rossa Formation (Campanian–Maastrichtian) at the Bottaccione section (Gubbio, central Italy); (B) photomicrographs of red limestones (sample YL13, Campanian) from the Yongla section near Gyangze, southern Tibet; (C) field photo of red shales (Santonian–Early Campanian) at the Chuangde section, southern Tibet; and (D) field photo of red cherts (Aptian) in the Yarlung Zangbo suture at the Xialu section, southern Tibet. Photo by Dr. Jiangang Wang. (For interpretation of the references to color in this figure, the reader is referred to the web version of this article.) 220

Table 1 Average contents of CaO, Al2O3 and SiO2 in different types of CORBs.

a b No. Place Section Type of Number of CaCO3 CaO Al2O3 SiO2 References CORB samples Average Stdev Max Min Average Stdev Max Min Average Stdev Max Min

1 Tibet, China Chuangde Ca-CORB 6 49.9 28.10 5.88 37.69 20.98 8.62 2.56 11.61 4.04 29.55 5.27 36.12 22.52 This study 2 Tibet, China Baergang Ca-CORB 2 65.3 36.80 1.08 37.56 36.03 3.26 0.60 3.68 2.83 25.71 0.58 26.12 25.30 This study 3 Tibet, China Tianba Ca-CORB 7 82.9 46.73 6.45 51.65 33.53 2.75 1.48 5.85 1.69 10.32 8.10 27.67 4.52 Chen (2008) 4 Turkey Unaz Ca-CORB 11 74.3 41.87 4.19 47.00 35.50 2.81 1.41 5.00 1.00 17.68 4.90 25.00 11.70 Eren and Kadir (1999) 5 Greece Pindos Ca-CORB 1 84.7 47.73 ––– 2.00 –––15.73 –––Baltuck (1982) 6 Austria Rehkogel Ca-CORB 4 58.3 32.85 ––– 8.65 –––26.25 –––Neuhuber and Wagreich (2009) 7 Austria Buchberg Ca-CORB 7 56.0 31.52 ––– 8.73 –––28.64 –––Neuhuber et al. (2007) 8 Austria Brandenberg Ca-CORB 4 47.9 26.99 –––10.41 –––26.70 –––Neuhuber and Wagreich (2011)

9 Austria Vircgdirf (AT03) Ca-CORB 1 84.8 47.76 ––– 2.58 ––– 8.63 –––This study 217 (2012) 115 Reviews Earth-Science / al. et Hu X. 10 Italy Pia da Stua Ca-CORB 3 90.3 50.89 1.82 52.95 49.48 1.41 0.53 1.88 0.84 4.71 1.60 6.00 2.92 This study 11 Italy Vispi Quarry Ca-CORB 9 94.7 53.34 0.84 54.51 52.05 0.59 0.27 1.14 0.27 3.29 1.10 5.39 1.79 Hu et al. (2009a) 12 Atlantic Aptian ODP1049C Ca-CORB 5 70.1 39.49 3.33 43.23 36.32 5.23 1.07 6.51 3.98 15.84 3.09 19.03 12.07 Chen (2008) 13 Atlantic DSDP Site 417D–19 Ca-CORB 1 60.0 33.80 ––– 2.07 ––– –––Murray et al. (1992) 14 Tibet, China Baergang Al-CORB 4 6.1 3.45 1.54 5.15 1.53 14.88 3.54 18.14 11.00 55.97 5.29 63.64 51.74 This study 15 Tibet, China Chuangde Al-CORB 12 7.2 4.08 5.65 15.29 0.37 16.39 2.71 19.22 11.60 52.04 4.97 57.78 43.10 This study 16 Tibet, China Tianba Al-CORB 2 7.0 3.97 4.35 7.04 0.89 16.81 2.19 18.36 15.26 53.05 4.53 56.25 49.85 Chen (2008) 17 Greece Pindos, Red shale Al-CORB 1 36.1 20.33 ––– 7.44 –––47.71 –––Baltuck (1982) 18 Greece Pindos, Maroon shale Al-CORB 1 2.2 1.22 –––12.36 –––58.65 –––Baltuck (1982) 19 Poland Silesian Nappe Al-CORB 13 1.3 0.74 0.07 0.85 0.64 13.62 1.06 15.08 11.31 65.43 1.66 68.54 62.53 Bak (2007) 20 Cezch Republic Marzak Al-CORB 8 0.8 0.43 0.04 0.52 0.39 14.73 0.93 16.40 13.90 67.63 1.74 69.20 64.70 Jiang et al. (2009) 21 North Atlantic DSDP site 105 Al-CORB 24 1.0 0.57 0.14 0.81 0.23 18.94 4.65 27.78 8.03 63.58 9.40 82.19 43.52 Lancelot et al. (1972) 22 North Atlantic Albian ODP1049C Al-CORB 3 21.5 12.12 1.97 13.67 9.91 14.40 0.56 14.92 13.80 41.59 2.45 44.41 40.02 Chen (2008) 23 Mid Paciﬁc Leg 129-800A Al-CORB 10 11.9 6.68 6.32 22.50 0.66 8.27 2.26 10.60 2.90 50.62 12.07 73.30 40.60 Karl et al. (1992) 24 New Zealand Manganotone Al-CORB 7 2.3 1.30 1.07 3.72 0.74 17.90 1.00 18.84 15.96 65.15 0.93 66.58 64.14 Hikoroa et al. (2009) 25 Romania Audia Al-CORB – 1.9 1.08 –––19.72 –––60.29 –––Melinte-Dobrinescu et al. (2009)

26 Romania Tarcău Al-CORB – 1.7 0.95 –––12.81 –––70.12 –––Melinte-Dobrinescu et al. (2009) – 248 27 Romania Marginal Fold Al-CORB – 8.3 4.67 –––12.46 –––66.2 –––Melinte-Dobrinescu et al. (2009) 28 Tibet, China Baergang Si-CORB 2 16.8 9.47 6.68 14.19 4.75 3.84 0.78 4.39 3.29 71.53 14.30 81.64 61.42 This study 29 Tibet, China Yongla (0517B) Si-CORB 1 0.2 0.14 ––– 2.59 –––91.44 –––This study 30 Tibet, China Xialu Si-CORB 8 0.4 0.23 0.10 0.39 0.08 1.66 0.87 2.94 0.80 89.84 3.62 93.42 84.12 Zhou et al. (2008) 31 Greece Pindos Si-CORB 3 1.4 0.76 0.61 1.47 0.41 0.71 0.28 0.95 0.40 92.94 0.29 93.11 92.61 Baltuck (1982) 32 Pacific DSDP leg 62 Si-CORB 3 0.3 0.19 0.07 0.26 0.12 0.70 0.04 0.73 0.65 96.42 0.41 96.67 95.95 Hein et al. (1981) 33 Pacific ODP leg 129-800A Si-CORB 10 0.8 0.44 0.24 0.93 0.05 3.05 1.56 5.33 0.57 85.76 4.89 95.70 79.80 Karl et al. (1992) 34 Pacific DSDP legs 20, 30, 32 Si-CORB 3 0.9 0.49 0.08 0.55 0.40 1.27 0.73 1.90 0.47 96.20 1.71 98.10 94.80 Murray et al. (1992) a Standard Average carbonate –– 75.1 42.30 ––– 0.79 ––– 5.13 –––Turekian and Wedepohl (1961) b Standard Deep sea carbonate –– 77.6 43.71 ––– 3.78 ––– 6.85 –––Turekian and Wedepohl (1961) c Standard Average shale –– 3.9 2.20 –––16.70 –––58.90 –––Wedepohl, 1971 (1991) d Standard Late Cenozoic Pacific red clay –– 1.2 0.70 –––16.60 –––54.90 –––Glasby (1991) e Standard Bonarelli black shales (Gubbio, Italy) – 38 3.4 1.90 ––– 4.20 –––72.00 –––Turgeon and Brumsack (2006)

a Indicated in Fig. 2. b Calculated from the CaO values using an oxide conversion factor of 1.7751. X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 221

CaO Atlantic) to 70% (Tarcău Nappe, Romania). Terrigenous clay and silt-sized quartz are the main grain types of the Al-CORBs and are 11 mainly derived from continental erosion and transported either by 10 a ocean currents or wind.

9 2.3. Si-CORBs 3 b 5 4 The CaCO3 contents of the Si-CORBs are extremely low (0.2 to 1.4%) except the samples from Baergang (Tibet) with 17% (Table 1). 12 2 The Al2O3 values are from 1% (Pindos, Greece) to 4% (Baergang, fi Ca-CORB Tibet). The SiO2 values vary from 86% (ODP site 800A, Paci c) to fi 13 96% (ODP legs 20, 30, 32, Paci c), except the samples from Baergang 6 7 (Tibet) which have 72% of SiO2. The cherts from the Baergang (Tibet) 8 1 are associated with red limestones and deposited on the northern Indian continental margin, above the CCD (Hu et al., 2006a). Other 17 Si-CORB examples are all pelagic red cherts in open oceanic environments Al-CORB 22 28 including those from the Pacific ODP sites, Xialu section in Tibet 15 23 16 14 18 and Pindos in Greece (Wang et al., 2009). This is the reason why d c 27 e 33 29 34 the cherts from Baergang have relatively low SiO2 and high CaCO3 25 21242019 26 30 3132 contents. Si-CORBs consist of biogenous SiO , which is mainly from Al O 5 SiO 2 2 3 2 radiolaria. Depositional environments of most CORBs were pelagic, deep Fig. 2. CaO–Al2O3–SiO2 ternary diagram on the classification of the CORBs. Numbers refer to Table 1. Green dots are Ca-CORBs; red rhombuses are Al-CORBs; blue triangles water in oceanic basins where they were generally far from shore- are Si-CORBs. Red crosses are standard rocks, see Table 1. (For interpretation of the ref- lines (Wang et al., 2009; Fig. 4). Outer shelf to upper/middle bathyal erences to color in this figure, the reader is referred to the web version of this article.) CORBs comprise mainly red pelagic limestones and marlstones Data are listed in Table 3 of the Appendix. (Ca-CORBs) characterized by high amounts of planktonic foraminifera, and constitute a widespread facies type in low palaeolatitudes, especially in the Tethyan palaeogeographic realm (e.g. Spain, Italy, carbonate is mainly derived from calcareous nannofossils and plankton- Switzerland, Austria, Poland, Greece, Turkey, Indian Himalaya). To- ic foraminifera. wards deeper bathyal depths, carbonate content decreases as the depositional area approaches the calcite lysocline. Red pelagic limestones and cyclic CORBs become scarce, and marls and marly shales 2.2. Al-CORBs become the dominant lithology/facies type. Below the CCD, in abyssal depths, either red claystones and carbonate-free shales (Al-CORBs) or

The CaCO3 contents of the Al-CORBs are generally lower than 10%, red cherts (Si-CORBs) were deposited, depending on the proportion ranging from 1% (Marzak, Czech Republic) to 21% (Albian ODP1049C, of terrigenous input. Such lithologies are widely distributed in the North Atlantic). The exception is a calcareous shale sample from the North Atlantic Ocean (Jansa et al., 1979). The sediment input of

Pindos basin (Greece) with a CaCO3 content of 36% (Table 1). The Al-CORBs and Si-CORBs is controlled largely by aeolian transport or Al2O3 values range from 7% (Pindos, Greece) to 20% (Audia Nappe, oceanic bottom currents, and siliceous biogenic production, which is Romania). The SiO2 values vary from 42% (Albian ODP1049C, North comparable to siliceous oozes in modern oceans, respectively.

80 Ma

19 12 20 21 6-9 17-18 13 10-11 5 25-27 22 31 4 32

34 33 23 28-30 14-16 1-3

Ca-CORB 24

Al-CORB Si-CORB

Fig. 3. Cretaceous (80 Ma) global palaeogeographic map (Hay et al., 1999) showing the positions of three types of CORBs (data refer to Table 1). Numbers refer to Table 1. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) The base map is downloaded from the www.odsn.de website. 222 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248

offshore winds

clacn sti i put aeolian dust exchange of CO current transport 2 of silt and clay

planktic carbonate production well oxygenated clastic nearshore facies depth range of Ca-CORB storm waves oxygen minimum zone ng elli possible upw

dissolution of cyclic Ca- or Al-CORB calcite lysocline carbonate Ca-CORB CCD

Al- or Si-CORB

Fig. 4. Overview of main depositional environments and main factors inﬂuencing CORB pelagic sedimentation. Revised from Wang et al. (2009). (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.)

3. Chronostratigraphy of CORBs brown clay overlies black clay at the Cenomanian/Turonian boundary, and the Upper Turonian to Santonian CORBs are brown clays The ages and rates of sediment accumulation of the CORBs provide interbedded with calcareous oozes. In DSDP 398 Turonian red to yellow constraints on hypotheses of origin of these deposits. Bimodal distri- brown mudstone is interbedded with rhythmic thin silt intercalations bution of CORB durations based on numerous sections suggests two and is overlain by Campanian–Maastrichtian red to brown marly chalk, groups (Fig. 5A–C; Scott, 2009). Short-term CORBs generally are less marl and claystone. These sections indicate that for most of the Late than 1 myr in duration and range up to 3 myr. Some short-term Cretaceous oxic bottom waters were long lasting throughout the North CORBs approximate the duration of Milankovitch cycles. Long-term Atlantic and periodic influx of coarse siliciclastic and calcareous CORB deposition was longer than 4 myr (Fig. 5B and C) and the red nannofossil oozes interrupted marine red bed deposition. Meanwhile, intervals may be composed of composite short-term CORB cycles. in the equatorial Atlantic and parts of the Caribbean Sea organic- carbon-rich strata were deposited during the Coniacian–Santonian, 3.1. Long-term CORB deposition which was interpreted as OAE 3 (see Jenkyns, 1980; Wagreich, 2009). Other examples of long-term CORB deposition are on the northern Long-term CORBs are several to tens of meters thick and have var- Indian shelf and in Alpine basins. For example, the Santonian to Cam- iable beddings. Long-term CORB deposition in the Tethys began dur- panian Chuangde Formation in Tibet was deposited on the northern ing the Late Aptian very soon after deposition of the organic beds of margin of the Indian plate and is composed of homogeneous red OAE 1a (Hu et al., 2005, 2006a). The Upper Aptian ORB 1 at Gorgo a shale interbedded with red limestones deposited near the base of Cebara (Italy) is 15 m thick and composed of cyclically interbedded the continental shelf below the CCD (Hu et al., 2006b; Chen et al., red calcareous marls (Hu et al., 2006a). ORB 1 is 2.6 m above the 2011). West of the Indian plate CORBs were deposited in the small Selli Level and overlies gray carbonate; and deposition began about partly isolated Penninic Ocean basin (Wagreich and Krenmayr, 890–500 kyr after the Selli bed and ORB duration was about 2005; Wagreich et al., 2009), where pelagic and hemipelagic and 5.12 myr (Fig. 6A). At Ouzon, Switzerland Upper Aptian pelagic red turbiditic red beds were deposited beginning in the Early Turonian beds that interbedded with gray limestone are 8.4 m thick and span and continued at different places into the Maastrichtian. At fl – at least 2.3 Ma in duration (Hable, 1997). In the deep-water Pindos Plagers ue, Switzerland, Turonian Santonian CORBs are red, thin Basin of Greece, siliceous CORBs 25 to 100 m thick extend from bedded pelagic limestones (Guillaume, 1986). In the Carpathian Valanginian to Coniacian (~50 myr in duration). Mountains of the Czech Republic and Romania, CORB deposition Late Cretaceous ORB deposition began in the Tethys and North began in the Late Albian and continued into the Early Palaeocene Atlantic basins soon after the Cenomanian/Turonian black shale of OAE (Melinte and Jipa, 2005; Melinte-Dobrinescu et al., 2009; Skupien et 2(Hu et al., 2005, 2006a; Scott, 2009; Yilmaz et al., 2010). CORB deposi- al., 2009). In the Sakarya Zone of northwestern Turkey Upper – tional durations ranged from about 2 to 26 Ma from Early Turonian to Cenomanian Upper Turonian red beds alternate with black shales Maastrichtian. In the Umbria–Marche basin, ORB 9 (Turonian), ORB10 and bracket OAE2 strata (Yilmaz et al., 2010). In the Pontide of (Coniacian to Santonian) and ORB 11 (Campanian to Maastrichtian) Turkey, the CORBs of the Late Santonian Unaz Formation are 5 to ű ű are about 19 m, 63 m and 147 m thick, respectively, with durations of 20 m thick (T ys z et al., 2012). ~2 myr, ~5 myr and ~15 myr, respectively (Coccioni and Premoli Silva, 2012). Along the western Atlantic margin close to North America in 3.2. Short-term ORB deposition DSDP 385, 386 and 603B, thin bedded yellowish red to brown clay alter- nates with greenish gray clay and dark red zeolitic clay extending from Short-term ORBs 1–7 are precisely dated as the Upper Aptian– Turonian to Maastrichtian (Jansa and Hu, 2009).Inthemiddleofthe Albian in the Piobbico core, the Monte Petrano section and the Vispi North Atlantic basin on the proto-Mid-Atlantic ridge, thin bedded, Quarry section in central Italy (Hu et al., 2006a; Fig. 6A–C). dark brown to red brown clay is interbedded with volcaniclastics in Correlation of these ORB intervals with the Albian stage boundaries DSDP 382 where CORB deposition lasted about 4 myr during the latest is based on traditional biozones of planktonic foraminifers and calcare- Campanian and earliest Maastrichtian. On the eastern margin of the ous nannofossils defined in the ORB sections and on graphic correlation North Atlantic near the coast of Iberia in ODP 641A lower Turonian to the CRET1 Database (Scott, 2009). Both methods result in a degree of X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 223

Duration Millions of Years 2 3 1 4 5

A 30 C 90

25 ORB9 OAE 2 20 95

15 ORB8 CENOMANIAN

10 ORB7 ALBIAN OAE 1d 100 OAE 1c 5

CORB Duration in Millions of Years 0 ORB6

1 3 5 7 9 1113151719212325272931333537 smunnA-ageMegA 105 CORB Stratigraphic Occurrences

ORB5 B ORB4 110

b1EAO ORB3

10 ALBIAN ORB2 APTIAN 115

ORB1

5 120

Number of Occurrences

0-12345678910>10 OAE 1a CORB Duration in Millions of Years 125

Fig. 5. CORB durations. (A) Durations in millions of years of stratigraphic and geographic occurrences of Tethyan CORB. (B) Number of CORBs in one million year increments showing the difference in durations of orbital-cycle control and long-term CORB deposition. (C) Chronostratigraphic succession of ORBs in Italian sections and their chronologic relation with oceanic anoxic events and the Aptian/Albian, and Albian/Cenomanian stage boundaries. Arrows mark positions of minor black shale beds. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) Panel A and Panel B data are from Hu et al. (2005); Scott (2009).

imprecision and uncertainty. The recognition of zones assumes that the Section and Point (GSSP) for the Aptian/Albian boundary has been ac- first and last appearances of species are synchronous, although species cepted. This criterion, however, was rejected by Owen (2002),as ranges relative to each other differ in various sections so zone bound- being of local value and not associated with any microfossil or aries tend to be diachronous. chemostratigraphic events. He supported the first occurrence of the The CRET1 Database is composed of the ranges of nearly 3500 taxa geographically widespread nannofossil Prediscosphaera columnata as and chronostratigraphic events in 161 globally distributed outcrops the base of the Albian stage (Owen, 2002). The first occurrence of and cored sections (Scott, 2009). CRET1 integrates the ranges of plank- L. tardefurcata projects into the Vispi Quarry section at 6.1 m by tonic and benthic foraminifera, nannofossils, dinoflagellates, ammo- graphic correlation to the CRET1 Database (Fig. 6C). nites, inoceramids, magnetochrons, radiometric dates, geochemical The Albian/Cenomanian boundary is defined at its GSSP at Mont events, and selected sequence stratigraphic markers defined in pub- Risou (France), by the first occurrence of Rotalipora globotruncanoides lished reference sections. Stage boundaries as defined by Ogg et al. (Gale et al., 1996; Kennedy et al., 2004). The first occurrence of (2004) were integrated into this database using the same criteria and R. globotruncanoides in the Monte Petrano section (Fiet et al., 2001) calibrated to the new age scale except for the base of the Cenomanian. (Fig. 6B) projects into the Vispi Quarry section between 58.4 and The age of the base Cenomanian is interpolated at 99.6 Ma by Ogg et 74.5 m (Fig. 6C). The numerical ages of ORB intervals in the Piobbico al. (2004) and at 97.13 Ma by Scott et al. (2009) based on chrono- core, the Monte Petrano and Vispi Quarry sections are interpolated by stratigraphic data from two different basins. The ages of ORB strata de- the lines of correlation, which are constrained by the bioevents and fined by Hu et al. (2006a) are interpolated from the Piobbico core, eccentricity cycles (Fig. 6A–C). Monte Petrano and Vispi Quarry sections (Fig. 6A–C). The Piobbico core cuts seven ORB beds (Fig. 6A) and yields a diverse Kennedy et al. (2000) proposed that the base of the Albian stage assemblage of Lower Aptian to Upper Albian planktonic foraminifers can be defined by the first occurrence of the ammonite Leymeriella and calcareous nannofossils (Erba, 1988; Tornaghi et al., 1989). These tardefurcata in the Col de Pré-Guittard section of the Vocontian biozones correlate ORBs 1–7 with Upper Aptian through the Upper Basin (France) at 66.5–66.9 m (Fig. 6D) and Niveau Paquier laminat- Albian substages (Hu et al., 2006a). The base of the Albian as approxi- ed black shale at 66.5–68 m. The first occurrence of L. tardefurcata matedbythefirst occurrence (FO) of P. columnata is between ORBs 2 was also used by Hoedemaeker et al. (2003) and Reboulet et al. and 3 in eccentricity cycle 1. The Albian/Cenomanian boundary was (2009) as the basal Albian zone, but no Global Boundary Stratotype above the core top and above ORB 7 (Fig. 6A). Secondary correlation 224 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248

10 70 m m .mFacnaiB B R. brotzeni

A ailgacS

0 ORB 7 E25 60 R. appenninica E. turriseiffelii R. appenninica T. ticinensis P. buxtorfi E25 P. praebuxtorfi ORB7 -10 E20 T. ticinensis 50 E. turriseiffelii LOA. verdieri ORB 6 T. subticinensis -20 E15 B. breggiensis E20 ORB 5 40

E10 -30 ORB6 A. albianus E15 ORB 4 30 D. cladoides C. ehrenbergi E5 Urbino bed ORB5 -40 T. breggiensis T. primula ORB 3 P. columnata

2BRO LO P. spinosum

E1 .mFdiocuFaitsicS 20 E10 FO L. conuspinum -50 P. cheniourensis L. arundum N. bucheri Marnes à Fucoïdes Fm. S. cretacea A. albianus 10 ORB4

1BRO T. primula

Monte Petrano, Central Italy (MIDK55) E5 -60 G. algerianus H. gorbachikae C. campanulata Urbino bed S. cretacea

Piobbico Core, Marche, Italy, Apt-Alb (MIDK13) C. litterarius P. columnata L. cabri ORB3 -70 G. blowi H. trocoidea 119.6 Ma 0 O. hinzii 71.4 m H. peridictya E1 OAE1a 121.8 Ma C. psygma H. planispira Maiolica Fm. Albian -80 Aptian Albian -10 LateAptian Early Mid Late m m 130 Ma 125 120 115 110 105 100 95 Ma 120 Ma 115 110 105 100 95 Ma MIDK3 DATABASE (10/30/2010) MIDK3 DATABASE (10/30/2010) m 100 180 m C R. asper D Top of Sampled Section aiblA 90 A. albianus H.similis 90 m = 90.45 Ma to 95.67 Ma 160

elddi 80 Lower Cenomanian E. monechiae

Mn 74.5 m = 97.20 Ma 140 T. orionatus 70 P. furtiva UncertainAge 120 G. meddii 60 H. steinmanni naiblA 58.4 m = 97.20 Ma

UpperAlbian Niveau Leenhardt noitamroFseuelBenraM A. salomoni 50 100 H. milletioides rewoL H. anglicus T. primula D. mammillatum FO B. convergens P. columnata FO H. trivialis Urbino bed Eiffellithus 40 H. albiensis FO S. hugoniotii 36.01 m = 101.974 Ma 80 monechiae S. achylosus FO S. longifurcatum S. glaber Niveau Paquier ORB 6 Niveau Paquier S. primitivum D. leightonense 112.64 Ma

.mFassoRailgacS 65.9 30 ORB 5 M. hoschultzi FO L. tardefurcata 113.15 60 UpperAlbian Hayesites S. glabra irregularis

20 MidAlbian naitpA Contessa Vispi Quarry, Italy - Albian Axopodorhabdus albianus 13.5 m = 108.85 Ma Tranolithus orionatus 40 Niveau Kilian 13.5 m = 108.85 to 108.60 Ma ORB 4 Pre-Guitard, France, Aptian-Albian (MIDK58) 10 Cretarhabdus Zeugrhabdotus erectus reppU striatus Watznaueria ovata L.Albian 2m ORB 3 C. granulatum 20 -0.3 m = 113.96 Ma P. spinosum U.Aptian T. bejaouaensis Hayesites albiensis T. roberti 0m Scaglia Prediscosphaera columnata Niveau Jacob Bianca Fm. LateAptianEarlyAlbianMidAlbian LateAlbian -10 m 0 m 120 Ma 115 110 105 100 95 90 Ma 120 Ma 118 116 114 112 110 108 106 104 102 Ma Aptian Early Mid LateAlbian Cenomanian MIDK41 DATABASE (05/15/2011) CRET1 DATABASE

Fig. 6. Cross plots A–C correlate the Albian ORB sections (Y-axes) with biostratigraphic databases (X-axes) and plot D correlates the candidate base Albian GSSP at Pre-Guitard, France (Kennedy et al., 2000) with MIDK41 database. The MIDK3, MIDK41 and CRET1 databases on the X-axes are successive developmental stages of Cretaceous chronostratigraphic data by graphic correlation of multiple sections to compile a global database (Scott, 2009). Y-axis is meters of logged section; X-axis is mega-annums. First occurrences of taxa marked by □; last occurrences by +. Red bars on Y-axes are positions of ORB strata. E-numbered events are eccentricity cycles. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.)

ties are the Selli Level in the lower part of the core (Erba, 1988)and frequencies at about 406 ka and 95–100 ka, an obliquity frequency the Urbino bed. Cyclostratigraphy is a third tool. The rhythmically of about 41 ka, and a weak precession signal of about 19–23 ka is interbedded marl, marly limestone and limestone of the Scisti a Fucoidi also recorded (Fischer et al., 1991; Grippo et al., 2004; Huang et al., Formation were deposited with orbital-cycle controlled eccentricity 2010). X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 225

The Monte Petrano section records continuous pelagic deposition ORB 4 at 109.11–108.67 Ma is slightly older than the FO of the Middle from Late Aptian to earliest Cenomanian and spans ORBs 3 through Albian marker Hoplites dentatus at 107.60 Ma. ORB 5 at 105.48– 7(Fig. 6B) (Hu et al., 2006a). The biostratigraphy is calibrated by 104.94 Ma is within the age range of Dipoloceras cristatum from planktonic foraminifera, calcareous nannofossils, dinoflagellates, and 105.54 to 104.16 Ma, which is the basal zone of the Upper Albian. ORB the Urbino bed. Fiet et al. (2001) correlated the Albian base at 6 at 103.97–103.21 Ma is about the same age as the Upper Albian the base of the measured section by the FO of the nannofossil Hysteroceras varicosum ammonite subzone at 103.03 to 102.27 Ma. P. columnata and the dinoflagellate Hystrichosphaeridium atlasiense, Within this ammonite subzone, included in the Mortoniceras inflatum and above the last occurrence (LO) of the foraminifer Ticinella Zone (Owen, 1999), Bown et al. (1999) reported the FO of E. monechiae. bejaouaensis, which is slightly below the base of ORB 3. The Albian/ The co-occurrence of the Upper Aptian–Albian ORBs and eccen- Cenomanian boundary based on foraminifera was correlated with tricity cycles defined by lithology suggests that ORBs may be related the FO of Rotalipora brotzeni as the equivalent to R. globotruncanoides to eccentricity cycles. The duration of ORBs is clearly highly variable that is the primary marker of the base of the Cenomanian stage and not an obvious climatic cycle frequency (Appendix Table 2). But (Kennedy et al., 2004). ORB 7 is just below this boundary. Twenty- do they begin on a Milankovitch beat or are they composites of mul- nine marl–limestone bundles represent the two eccentricity frequen- tiple cycles? To test these hypotheses the Upper Aptian to Upper cies of about 400 ka and 100 ka (Fiet et al., 2001) and support the Albian ORB beds are plotted with the eccentricity cycles on a time correlation of Monte Petrano with the Piobbico core. scale (Fig. 7). The ages of the eccentricity cycles are interpolated by The Vispi Quarry section is composed of limestone interbedded plotting their positions in the Piobbico core and the Monte Petrano with ORBs 3 to 6. The limestones yield low diversity, moderately section to the CRET1 Database by means of the pelagic microfossils preserved calcareous nannofossil assemblage diagenetically altered (Fig. 6). The ages of the ORBs in the same sections are interpolated (Appendix Table 1). The most abundant species is Watznaueria in the same manner. The ages of ORB onset vary relative to eccentric- barnesiae, and Eprolithus floralis, Zeugrhabdotus embergeri, R. asper ity cycles. In the Piobbico core ORB 3 began near the end of cycle 31 and R. angustus are consistently present. Correlation of this section and ended at the end of cycle 29. ORB 4 coincided with but was with the Albian stage and substages is based on new analyses of cal- shorter than cycle 25, but ORBs 5, 6 and 7 straddle cycle boundaries. careous nannofossils (Fig. 6C; Appendix Table 1). ORB 3 at the base The thicknesses of ORBs 1–7 range from 17.5 to 0.6 m and the dura- of the section is in the basal Albian stage, which is approximated by tions range from 5.51 to 0.14 myr. In the Piobbico core the Albian the FO of P. columnata, the nannofloral event at the base of NC8A ORBs 3 to 7 are either shorter or somewhat longer than the eccentric- zone (Sissingh, 1977; Bralower et al., 1993, 1995). This event is ity cycles but are not composites (Grippo et al., 2004; Huang et al., followed by the FO of the nannofossil Hayesites albiensis, which 2010). So no direct relationship is evident between the Upper Aptian marks the base of the NC8B subzone (Bralower et al., 1993, 1995) to Upper Albian ORBs and eccentricity cycles. also in the Early Albian. The top of ORB 4 is just below the base of the Middle Albian as defined by the FO of the nannofossil Tranolithus 4. Mineralogical composition and color origin of CORBs orionatus, which defines the base of Subzone NC8C (Bralower et al., 1993, 1995). In the Middle Albian, the FO of Axopodorhabdus albianus The origin of the red color is one of the keys to understanding the defines the base of NC9A nannofloral subzone (Bralower et al., 1993, formation of CORBs. In this section, we discuss mineral compositions, 1995), which is between ORBs 4 and 5. The top of ORB6 is slightly iron oxides and spectral features of the CORBs in three areas: Tibet of below the base of the Upper Albian defined by the base of subzone China, Italy and the Atlantic ODP1049C, in order to interpret the ori- NC9B and the FO of Eiffellithus monechiae. At Col de Palluel, gin of the CORBs' colorations. Hautes-Alpes, France, the FO of E. monechiae is in the lower part of the Upper Albian coincident with the last occurrence of H. albiensis 4.1. Mineral composition of CORBs (Gale et al., 2011). The graphic interpretation of the Albian interval in the Vispi Based on the qualitative analysis, we use the whole-pattern fitting Quarry section is constrained by the nannofossil ranges and the posi- technique under the support of a general structure analysis system tions of ORB 3–6(Fig. 6C). The age at the base of the section is inter- program (Larson and Von Dreele, 2004) based on Rietveld refinement polated at 113.96 Ma, slightly younger than the FOs of P. columnata to calculate the weight fraction of every X-ray diffractometry and H. albiensis; the latter species defines NC8B Zone. So, the oldest (XRD)-detected mineral for limestones from Italy and shales from part of NC8 Zone, Subzone NC8A is not sampled. The base of Subzone Tibet. The red limestone in Scaglia Rossa Formation from the Vispi NC8C at the FO of T. orionatus is at 13.5 m and here projected at Quarry section is mainly composed of calcite, quartz and boehmite 108.85 Ma at the base of the Middle Albian. An unconformity or (Fig. 8), with minor hematite, illite, montmorillonite, and albite in normal fault is posited at 36.01 m because the age of the FO of the residues after the removal of carbonates (Cai et al., 2009). The E. monechiae is slightly offset above the termination of ORB 6. The quantitative analysis shows that the calcite weight fraction ranges age of this bioevent is projected at 101.97 Ma in the lower part of from 92 to 99%; boehmite is 0.1 to 0.9%; and the quartz ranges from the Upper Albian. The uppermost interval from 36.01 m to 90 m is 0.5 to 6.5% (Fig. 8). poorly constrained and two interpretations are proposed. The base The red shales from the Chuangde section are composed of chlo- of the interval is pinned on the FO of E. monechiae and the top of rite, illite and kaolinite, quartz, albite, calcite, and hematite. They dif- the steeper line of correlation (LOC) is at the LO of E. monechiae and fer from red limestone in the same section by the higher contents of the alternative LOC is at the LO of A. albianus. The position of the hematite and chlorite (Fig. 9). The weight fractions of chlorite, illite Albian/Cenomanian boundary defined by the FO of R. globotruncanoides and kaolinite vary in the range of 6–23%, 26–57% and 3–13%, respec- at 97.2 Ma in the CRET1 Database is projected into the Vispi Quarry sec- tively (Fig. 9). The weight fractions of quartz, albite and calcite are in tion between 58.4 m and 76.6 m. Additional data are needed to con- the range of 3–30%, 5–24% and 0–33%, respectively (Fig. 9). strain this interpretation more narrowly. The Upper Albian–Lower Turonian red shales of the Eastern The ages of the ORBs 3–6 in the Vispi Quarry section appear to be Carpathian inner Moldavides nappes are dominated by around 50% il- coeval with those beds in the Piobbico core and the Petrano section. lite followed by up to 11% kaolinite and around 3% each of chlorite The numerical ages are projected from the database by the correlation and hematite, and quartz and chalcedony together make around lines and the ORBs are correlated with the Albian ammonite zones 25% (Papiu et al., 1983; Melinte-Dobrinescu et al., 2009). Dolomite by the same lines. ORB 3 at 111.73–110.61 Ma is within the range of and calcite are up to 1%. In the outermost part of the Eastern Early Albian Douvilleiceras mammillatum at 111.60 to 110.26 Ma. Carpathians Outer Moldavides, quartz and chalcedony together are 226 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248

sedionacnurt

NC10B 96

-obolg.R

97.13

acininneppa.R OAE 1c - NC10A Amadeus Event 98.96 - 98.23 7

sisnenicit.T

100.52 100

sisnenicitbus.T

NC9B 101.83

102

103.14

sisneiggerb.B

104

NC9A 5 105.75

106

108

4 Urbino Event

Mega-annums

5.901-9.111 109.8 - 109.5 011 NC8C b1EAO alumirp.T Niveau Leenhardt

110.3 - 110.0 3 111.56

211411 Niveau Paquier

112.7 - 108.2 selcyCyticirtneccE NC8B oegA

aripsinalp.H Niveau Kilian 114.9 - 113.4

gninnige

113.64

Niveau Jacob 1BRO 117.5 - 115.5

NC8A sisneauoajeb.T

811 611

119.50 NC7C

120

Fig. 7. Correlation of ages between oceanic red beds, black shale beds and eccentricity cycles. Stratigraphic data are from the Italian sections (Kennedy et al., 2000; Grippo et al., 2004; Hu et al., 2006a). Planktonic foraminifera zones are from Ogg et al. (2004); and nannofossil zones are from Bralower et al. (1995). Numerical age calibrations are from the CRET1 Database (Scott, 2009).

43.4%, followed by illite (35.6%); other components (kaolinite, chlo- as sequential leaching of the sediments. When the iron oxide content rite, hematite, calcite, and feldspar) are in minor amounts, each is over its lowest detection limitation, the XRD method is one of the representing less than 6%, and combined amounting up to 15%. most efficient and frequently used. Compared to the XRD method, According to the XRD analyses, the less than 2 μm grain-size frac- the DRS is considered to be a useful, rapid and effective detection tion of the Aptian–Albian red beds in ODP Hole 1049C in the North tool that can detect the presence of iron oxide as low as 0.01% Atlantic mainly consists of quartz, albite and clay minerals, which (Deaton and Balsam, 1991). The transmission electron microscopy is are mainly illite, with subordinate chlorite, minor kaolinite and mont- extremely effective in identifying the very fine (nanno-scale) mineral morillonite or mixed-layer illite–montmorillonite (Han et al., 2008). grains. The well crystallized illite and chlorite were mainly included in the μ larger than 2 m grain size fractions. The weight fractions of chlorite, 4.2.1. Quantitative analysis of iron oxide with XRD – – – illite and kaolinite vary in the range of 0 30%, 15 75% and 0 20%, re- The XRD fitting results suggest that the content of hematite in the spectively (Fig. 10). The weight fractions of quartz and albite are in red shales from the Chuangde section varies from 3.81 to 12%wt.% – – the range of 0 40% and 0 48%, respectively (Fig. 10). (Fig. 9)(Li et al., 2009; Cai, unpublished).

4.2. Iron oxides in CORBs 4.2.2. DRS data of the CORBs The DRS data of red limestones from the Vispi Quarry section Species and amount of iron oxides, which are key issues in evalu- suggests that hematite was composed of different crystalline types ating the redox condition, can be determined by XRD, diffuse reﬂec- (Cai et al., 2008). A broad spectral band is present in the range of tance spectroscopy (DRS), transmission electron microscopy as well 520–620 nm of ﬁrst derivative spectra and is centered at 560 nm, X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 227

Calcite % Boehmite (%) m 92 96 0.4 0.8 assoRailgacS 11.0

10.8

10.6

10.4

acinaiBailgacS 10.2

10.0

9.8 2 4 6

Limestone Chert Reddish Quartz (%)

Fig. 8. Distribution of mineral components based on whole-pattern ﬁtting of the XRD pattern in the Turonian part of the Vispi Quarry section, Italy. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.)

with a shoulder present at 574 nm and/or 587 nm (Fig. 11). The greater Two DRS spectra in red shales from the Chuangde section were mea- the wave-length, the higher is the crystallinity. For example, the peak sured in bulk rock and residual after the CBD (citrate–bicarbonate– centered at 587 nm suggests the presence of well-crystalline hematite, dithionite) treatment (Fig. 12C and D). Only a single hematite peak such as specularite and mica hematite (Cai et al., 2008). The peak was present in the ﬁrst derivative spectra of both bulk rock and samples centered at 575 nm was assigned to weakly-crystalline hematite and with CBD treating (Fig. 12C; Hu et al., 2006b; Li et al., 2009). Three DRS rhodochrosite (Cai et al., 2008). spectra (Fig. 12A, and B) were collected for each sample from ODP Hole

Illite (%) Kaolinite (%) Calcite (%) Quartz (%) 30 40 50 4 8 12 0 10 20 30 10 20 30 16

)m(htpeD

0 10 20 10 20 4 8 Chlorite (%) Albite (%) Hematite (%)

Fig. 9. Distribution of mineral components based on whole-pattern ﬁtting of the XRD pattern in the Santonian–Campanian part of the Chuangde section, southern Tibet, China. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) 228 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248

Depth Albite (%) Chlorite (%) (mbsf) 0 25 50 0 20 40

139.3 1 2

3 4 5 141.8

146.0 6 7 8

148.1 0 25 50 15 45 75 0 10 20 Cycle Quartz (%) Illite (%) Kaolinite (%) No.

Fig. 10. Distribution of mineral components based on whole-pattern ﬁtting of the XRD pattern in the Aptian–Albian part of the ODP 1049C core, northern Atlantic Ocean (after Li et al., 2011). (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.)

1049C before and after CBD treating, and from a synthetic mixture stain minerals in these brown and orange sediments are hematite and with different fractions of hematite and goethite, respectively. Two goethite (Li et al., 2011). peaks, centered at 435 and 565 nm, were present in the spectra of samples with brown and orange colors (Fig. 12A) (Han et al., 2008; 4.2.3. Quantitative analysis of iron oxide concentrations in ODP Hole Li et al., 2011). After treatment by CBD, the brown and orange sam- 1049C ples changed to green or white, and the peaks assigned to hematite To obtain the absolute contents of iron oxides in these sediments, and goethite disappeared in the ﬁrst derivative spectra (Fig. 12B). Li et al. (2011) performed a quantitative analysis using DRS with mul- Then, with the different hematite and goethite contents that were tiple linear regression (see Fig. 23). The results suggest that the he- added to the CBD treated samples, the DRS analysis suggests the resto- matite content in the Albian brown samples ranges from 0.13 to ration of brown and orange colors (Fig. 12B). This indicates that the 0.82% and averages 0.51%, and the goethite content is in the range

0.5 Red Non-red 0.4 ODP Hole 1049C

0.3 445nm 0.2

0.1

0.0 565nm 575nm 0.3 Chuangde 0.2

0.1

First Derivative Value 0.0

0.3

0.2 Vispi Quarry 0.1

0.0

400 450 500 550 600 650 700 400 450 500 550 600 650 700

Fig. 11. DRS data of CORBs from the Aptian–Albian ODP1049C (data from Li et al., 2011), Chuangde section of Tibet (data from Hu et al., 2006b), and Vispi Quarry section of Italy (data from Hu et al., 2009a). (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 229

0.50 0.50 Black Orange (after CBD) 0.45 AB0.45 Orange Iron oxides added 0.40 Brown 0.40 Green 0.35 White 0.35 0.30 0.30 0.25 0.25 0.20 0.20 0.15 0.15 0.10 0.10

First derivative Value 0.05 0.05 0.00 0.00 -0.05 -0.05 400 450 500 550 600 650 700 400 450 500 550 600 650 700

0.35 0.35 before-CBD CDafter-CBD 0.30 0.30 18 19 0.25 0.25 17 20 S s T 0.20 0.20 r R t 18 s 19 r 16 R S 0.15 17 U 0.15 21 q 20 u q t 16 Q Q T V 21 0.10 15 v 0.10 22 15 5 14 p AE p u 6 13 22 4 P Ww AC AD P U 12 23 AB ac ae 3 7 1011 24Xx 25 AA27 ad 23 2 8 9 Yy 26Zz aa 28 5 14 v 0.05 o ab 29 0.05 First derivative Value 7 24 31 30 31 4 9 13 o V 2930 n O 6 8 10 11 12 O 25 27 e B2 3 26 28 ad ae d f g i m c e w AE j k l N b d f g h i n W x ac AD b h C F H I m N X y z aa ab AC c L M D E j k l Y Z AA AB I K 575nm G J L M 0.00 E G J 0.00 K C D F H B 575nm -0.05 -0.05 400 450 500 550 600 650 700 400 450 500 550 600 650 700 Wavelength (nm) Wavelength (nm)

Fig. 12. First derivative spectral patterns of specimens Pre-CBD (citrate–bicarbonate–dithionite) and post-CBD. Panels A and B are from ODP1049C (after Li et al., 2011); panels C and D are from Tibet, China (after Li et al., 2010). (For interpretation of the references to color in this figure, the reader is referred to the web version of this article.) of 0.22–0.81% and averages 0.58% (see Fig. 23). The content of hema- 4.3.2. Red coloring mechanism of red limestone tite in Aptian orange samples is in the range of 0.19–0.46% and aver- Hematite originates in red Turonian limestone of the Vispi Quarry ages 0.35%, and the content of goethite is in the range of 0.29–0.67% section in two ways. The first is continental origin by seasonal winds to- and averages 0.50% (see Fig. 23). gether with boehmite and quartz, and the other is authigenic hematite origin which is precipitated from sea water (Cai et al., 2009). The grain 4.3. Origins of the red color of CORBs size of continentally-derived hematite is more than several microns and is heterogeneously distributed in the limestone. Autogenic hematite is 4.3.1. Determination of staining minerals in very low concentrations but homogenously distributed within the in- Different stain minerals such as hematite and goethite behave dif- terstices of the calcite microlites, calcite grains, holes of the microfossils, ferently in terms of determining the color of sediments. Hematite ap- and entrapped in the calcite structure. The continentally-derived hema- pears blood-red in color (3.5R–4.1YR, Munsell hue), imparting a red tite cannot be the major cause that imparts the red color to limestone color to sediment. In contrast, goethite is bright yellow (8.1YR–1.6Y, because of its heterogeneous distribution and low proportion (Cai et Munsell hue), imparting a bright yellow color to sediment (Torrent al., 2009, 2012). and Schwertmann, 1987; Deaton and Balsam, 1991; Li et al., 2011). High resolution transmission electron microscopy (Cai et al., Mn2+-doped calcite can be excited to radiate visible red light by 2012) suggests that the nanometer-sized hematite grains are in the means of thermal, cathodoluminescent, ultraviolet or laser energy interstice of calcite grains. This is verified by selected areal electron sources (e.g. Macedo et al., 1999; Cai et al., 2008). However, the diffraction patterns. The hypothesis is that iron was originally trapped color was significantly lighter after it had been ground to a powder in the calcite structure in two forms, the first was ferrous iron in the than before it had been in the crystals, which suggests that calcite lattice, and the second was iron in the interstices of the calcite 2+ Mn -bearing calcite may only give the limestone a light red tinge structure, which replaces the oxygen from the CaO6 octahedron to 2− (Cai et al., 2008). Dyeing of kidney-red hematite is 10 times greater form Fe2O3 clusters. These clusters were then oxidized to form he- than rhodochrosite, which has a similar coloring mechanism to matite clusters by releasing the electron to dissolved O2 in groundwa- Mn2+-doped calcite (Cai et al., 2008). ter. These hematite clusters grow to several tens of nanometers in the For hematite minerals, different degrees of crystallinity may cause calcite structure, and was a major factor imparting the homogeneous different redness of sediments. Cai et al. (2008) showed that crystal- red limestone color (Cai et al., 2009). linity of specularite is best, mica hematite is middle and kidney-red The electron-spin resonance spectra of limestones from the Vispi hematite is the poorest of crystallinity. The DRS peak of hematite Quarry clearly show that manganese is present as octahedral cations was centered at 577, 585 and 586 nm for kidney hematite, mica hein the calcite structure (Cai et al., 2009) and was incorporated during matite, specularite, respectively, which suggests that the DRS peak calcite formation (Calderoni and Ferrini, 1984). This suggests that the shifts to a longer wave-length when its crystallinity increases. DRS calcite in the red limestone is a type of manganese-bearing calcite, data showed that the dyeing ability decreased as the degree of hema- which may give additional red color to those limestones. tite crystallinity increased, with the order of kidney-red hematite, DRS work suggests that both goethite and hematite are present in mica hematite and specularite. the brown and orange beds (Fig. 11), but are not found in black and 230 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248

A 10

FE 1 2 3 Fe Mg Na K Mn P Ti 6 7 0.1 8 B 9 10 10 11 12 e FE 1

V CrCoNiCuZnRbSrZrMoBaU

0.1

Fig. 13. Enrichment factors (relative to the Average Shale of Wedepohl, 1971) of major and trace elements of the Ca-CORBs. Numbers refer to Table 1. (A) major elements; (B) trace elements. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) Data source is listed in Table 3 of the Appendix.

white beds from ODP 1049C (Han et al., 2008; Li et al., 2011; Fig. 11). considerable diagenesis (Arthur and Fischer, 1977), whereas the red The CBD treatment of brown and orange samples recovered synthetic marls in ODP 1049C are largely unaffected by diagenesis (Erbacher hematite and goethite (Fig. 12) suggesting that ferric oxide is in the et al., 2001). Therefore we propose that goethite originally formed to- form of nanograins of poorly crystallized or amorphous hematite gether with hematite. Subsequently, the goethite was transformed to and goethite. Consider the following facts: (1) goethite is not found hematite during late diagenesis, as already documented in other sed- in red limestone from the Vispi Quarry section, and (2) red lime- imentary rocks (Berner, 1969; van Houten, 1973; Channell et al., stones from the Vispi Quarry are lithiﬁed and have experienced 1982).

1 14

0.5 15 Fe Mg Na K Mn P Ti 19

20 5 22

FE 24 1

0.5 VCrCoNiCuZnRbSrZrMoBaU

Fig. 14. Enrichment factors (relative to the Average Shale of Wedepohl, 1971) of major and trace elements of the Al-CORBs. Place number refers to Table 1. (A) Major elements; (B) trace elements. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) Data source is listed in Table 3 of the Appendix. X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 231

50 Vispi Quarry, Italy

MKaNgMeF n iTP ViCoCr N CuZn Rb Sr Zr Mo Ba U

Red limestone White limestone Bonarelli black shale 0.1

Fig. 15. Enrichment factors (relative to the Average Shale of Wedepohl, 1971) of major and trace elements of the red limestones, white limestones and black shales of Bonarelli level from the Vispi Quarry section of Italy. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) Data source is listed in Table 4 of the Appendix.

10 Red shale nonred shale Albian, ODP 1049 C

iTPnMKaNgMeF ViCoCr N Cu Zn RbSr Zr Mo Ba U 0.3

5 Chuangde,Tibet

iTPnMKaNgMeF ViCoCr N Cu Zn RbSr Zr Mo Ba U 0.3 5

Marzak, Czech Republic

iTPnMKaNgMeF ViCoCr N Cu Zn RbSr Zr Mo Ba U 0.3

Fig. 16. Enrichment factors (relative to the Average Shale of Wedepohl, 1971) of major and trace elements of the red shales and non-red shales from (A) Albian ODP1049C of the North Atlantic; (B) Chuangde section of Tibet; and (C) Marzak section of the Czech Republic. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) Data source is listed in Table 4 of the Appendix. 232 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248

4.3.3. Red coloring mechanism of red shale 100 Red shales from the Chuangde section have 3.81 to 12% in weight hematite which is much higher than that of the limestones from Vispi 80 Quarry (~0.1–0.2%, Hu et al., 2009a). Hematite in the Albian red shale

)%tw(OCaC in ODP1049C is no more than 0.13–0.82% (Li et al., 2011). Hematite grain size ranges from several tens of nanometers to several tens of 60 microns in red shales based on scanning electron microscopy. The 3 red color of shales in the Chuangde section is imparted by both 40 nanograin hematite and larger size (tens of microns) of hematite R² = 0.876 based on CBD treatment (Fig. 12). 20 4.4. Summary 0 (1) Hematite, goethite and Mn2+-bearing calcite impart the red 0 100 200 300 color of CORBs. Hematite imparts a red color; goethite imparts ∑ REE (ppm) a bright yellow color; Mn2+-bearing calcite gives a light red Σ – tinge. The red color decreases when either the degree of hema- Fig. 17. REE CaCO3 content crossplot for the Ca-CORBs showing a strongly negative tite crystallinity or grain size of hematite mineral increased. relationship. Data source is listed in Table 6 of the Appendix. (2) Both hematite and goethite are responsible for the brown and orange colors of the Albian–Aptian sediments from ODP 1049C. Hematite clusters of several to tens of nanometers in

the calcite structure are the main cause of the red coloring of Shale (Wedepohl, 1971) with EFMn ranging from 0.67 to 4.71 limestones, and manganese-bearing calcite may give addition- (Fig. 13A, Appendix Table 4). P is enriched in red limestones from

al red color to limestones in the Vispi Quarry section. Goethite Pia da Stua and Vispi Quarry of Italy with EFP of 9.15 and 5.90, respec- was suggested to be originally formed together with hematite, tively. Na is depleted from red limestones in Austria with EFNa ranging but was subsequently transformed to hematite during late dia- from 0.28 to 0.47. genesis. In comparison with Bonarelli black shales (Turgeon and Brumsack, (3) The red color of shales in the Chuangde section resulted from 2006), the white limestones and red limestones in the Vispi Quarry

hematites of different grain sizes, ranging from several tens are enriched in Mg and Mn, and depleted in Na (Fig. 15). Fe2O3 in of nanometers to several tens of microns in size. the red limestones are from 0.12 wt.% to 0.41 wt.% with an average of 0.22 wt.%, which is signiﬁcantly higher than in the white lime-

5. Elemental geochemistry stones, where Fe2O3 values range from 0.02 wt.% to 0.19 wt.% with an average of 0.12 wt.% (Hu et al., 2009a; Fig. 15). 5.1. Data source and method

Elemental geochemical data of CORBs are available from numerous publications. In this study, we compiled elemental geochemical 10 data of both Ca-CORBs (red limestones/marls) and Al-CORBs (red A shales) in the world (Appendix Tables 3–6). Geochemical data from the Si-CORBs (red cherts) are rare and not discussed in this paper. Elemental values were normalized to Al in order to account for dilu- 1 tion effects by potential biogenic components such as carbonates, silica, and phosphorites. Concentrations are compared to the average values of the Average Shale (Wedepohl, 1971) and expressed as relative 0.1 enrichment factors (EFElement). Enrichment factors were calculated Sample/PAAS using: EFElement =(element/Al)sample/(element/Al)average shale.Based on the elemental concentrations within the sediments, we use the term Vispi Quarry Chuangde ‘enrichment’ for EFElement ≥5, and we consider the samples ‘depleted’ when EFElement ≤0.5 (Turgeon and Brumsack, 2006). La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu We normalized the concentrations of rare earth elements in samples 10 to the REE concentrations of post-Archean Australian sedimentary rocks B (PAAS, McLennan, 1989). The Ce-anomaly is calculated according to

Wilde et al. (1996):Ce-anomaly(δCe)=log [2Cen /(Lan +Prn)], where n is the shale-normalized concentration. 1

5.2. Ca-CORBs (red carbonates)

Sample/PAAS 0.1 5.2.1. Major elements Major elements Fe, Mg, K, and Ti from red limestones and marls are relatively similar to the Average Shale (Wedepohl, 1971) with Chuangde Marzak

EFFe,EFMg,EFK,EFTi ranging from 0.62 to 1.47, 0.86 to 4.16, 0.58 to 1.65, 0.62 to 1.42 respectively (Fig. 13A, Appendix Table 3). Mn is La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu enriched in red limestones from Tibet (EFFe is 7.14 in Chuangde, Fig. 18. The post-Archean average Australian shale (PAAS, McLennan, 1989)-normalized 25.65 in Baergang, and 8.89 in Tianba) and Italy (EF is 6.48 in Pia Fe REE patterns. (A) Ca-CORB; (B) Al-CORB. (For interpretation of the references to color in da Stua and 15.67 in Vispi Quarry). Mn values of red limestone/marl this ﬁgure, the reader is referred to the web version of this article.) in Austria and Atlantic ODP1049C site are comparable to the Average Data source is listed in Table 6 of the Appendix. Gorge a Cerbara, Italy Yenicesihlar, Turkey Roter Sattel, Switzerland La Bédoule, France 13 13 δ C carb, ‰, PDB 13 δ C carb, ‰, PDB T. primula 13 δ C carb, ‰, PDB B. δ C carb, ‰, PDB breggiensis 1.0 2.0 3.0 4.0 2.0 3.0 4.0 G. alg. 2.0 4.0

43 m C9

G. ORB1

algerianus C8

(NC7)

42 G. C9

G. barri G.

ferrolensis

R. angustus R. .H ta./ErhSineRves15(02 217 (2012) 115 Reviews Earth-Science / al. et Hu X. L. cabri L. C8

s 41

G. ferrolensis G.

C7 G. C8

ferrolensi

10m

40 study this

10m C7

. cabri

10m

L C7 C7

39 cabri L.

L. cabri L.

(NC6)

OAE1a 2011 al., et Stein C3-C6 –

C6 248 C6 C5 C5 C4 C4 38 C6

G. blowi G. C. litterarius C. C4 C3 C5 C2 C3

G. blowi G. C2 C3

37 blowi G. C1-C2 blowi . C3-C6 C1 G H. sigali C1

C2 2.0 4.0

Fig. 19. Chemostratigraphic correlation by stable carbon isotopes of the Aptian interval from OAE 1a to ORB1 of Gorge a Cebara, Italy (Stein et al., 2011 and this study) with the Yenicesihlar section, Turkey (Hu et al., 2012b), the Roter Sattel section, Switzerland (Menegatti et al., 1998), and the La Bédoule section, France (Kuhnt et al., 2011). (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) 233 234 Vispi Quarry, Italy Halle-Oerlinghausen/ Buchberg, Austria English Chalk 13 Culver Cliff, Germany 13 13 Ccarb (‰ PDB) Ccarb (‰ PDB) carb (‰ PDB) 13 C Ccarb (‰ PDB) 400 kyr 1.5 2.0 2.5 3.0 cycle 2.0 2.4 Pewsey 3.0 4.0 5.0 7 m 7 2.0 3.0

20 Glynde 91

UC8a

CC12 M. coronata 6 6

tessa)

(Con 60

m i d d l e

5 iddle Turonian

5 15 M .H ta./ErhSineRves15(02 217 (2012) 115 Reviews Earth-Science / al. et Hu X.

NC14

elvetic

Tu10 (low-woollgari) i

H. h H. 50

4 i

ORB9 cuvieri 4

woollgar

Tu8 (Round Down) T u r o n i a n

picalis-

a 5 Tu7 10 40 92 3

3 Tu6

UC7

CC11

30 Tu5 (Lulworth) Tu4 –

Lowerronian Tu

2 248 2 subhc./herc. 5 10 Tu3

NC13**

nodosoides 20 Tu2

e a r l y

1 mytiloide

bio S2)

Helvetoglobotruncana helvetica

1 archaeocretacea

? (Gub Tu1 (Holywell)

W. W. 93

vonense

10 hattini Ma

de 0 c

CC10

UC6b

NC12* m 15 OAE2 ju. m OAE2 0 b a a’ m pictus

geslin

NC11 0

R. cushmani R.

Fig. 20. Chemostratigraphic correlation by stable carbon isotopes of the Turonian interval from OAE 2 to ORB9 of the Buchberg section, Austria (Wendler et al., 2009) with the Vispi Quarry section, Italy (Stoll and Schrag, 2000; Tsikos et al., 2004), and the Halle-Oerlinghausen/Culver Cliff section, Germany (Voigt et al., 2007), compared to the Upper Cretaceous carbon isotopic curve from the English Chalk (Jarvis et al., 2006). (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 235

13 13 Ccarb (‰ PDB) Ccarb (‰ PDB) 123456Ma 1 2 3 4 5 6 Ma 121 90.0

‘Pewsey’

122 Glynde 91.0

Middle

Late

ORB1 ORB9 Round Down 123 92.0 Lulworth

Aptian

Turonian

ylraE

Early 124 Holywell 93.0

b a OAE1a 125 OAE2 94.0

Cenomanian

Barremian

126 95.0

Fig. 21. Stable carbon isotopic changes during the OAE–ORB transitions. (A) OAE 1a–ORB1 transition with data from the Yenicesihlar section, Turkey (Hu et al., 2012b); (B) OAE 2– ORB9 transition with data from the English Chalk (Jarvis et al., 2006). (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.)

5.2.2. Trace elements significantly less REE than detrital clay and heavy minerals (Piper, 1974). Σ The element V is depleted in red limestones from Baergang (EFV = The differences in the REE values between red and white limestones are 0.31) and Rehkogel (Austria) (EFV =0.08; Fig. 13B), but it is close to minor (Hu et al., 2009a). the Average Shale (Wedepohl, 1971) in other places. The elements Overall, the PAAS-normalized REE patterns of studied limestones Cr, Co, Ni are similar to the Average Shale (Wedepohl, 1971) with exhibit seawater-like REE distribution patterns characterized by EF values ranging from 0.6 to 3.2 (Fig. 13B). The element Cu is small enrichments in heavy REE (HREE) relative to light REE (LREE) enriched at Baergang (Tibet) (EFCu =6.09) and is depleted at with negative Cerium anomalies (Fig. 18). δ Buchberg (Austria) (EFCu =0.38). The element Zn is enriched at the Red limestones have 0.35 of averaged Ce (Appendix Table 6). The δ Vispi Quarry with EFZn =5.94. The element Rb is depleted from Ce values of Cretaceous red limestones are very similar to those of δ b Tianba (Tibet) and Brandenberg (Austria) with EFRb =0.4. The ele- typical oceanic seawater ( Ce values ranging from 0.1 to 0.4, fi ment Zr is highly depleted from Rehkogel (Austria) with EFZr =0.04 Elder eld and Greaves, 1982). Also, in the Vispi Quarry section, the δ and depleted at Buchberg (Austria) with EFZr =0.46. The element Sr red limestones show very similar REE pattern and Ce values to is enriched at Tianba (EFSr =15.42) and the Vispi Quarry with those of the white limestones (Hu et al., 2009a). EFSr =49.44. The element Mo is enriched at the Vispi Quarry with EF =16.56 and is depleted at Chuangde with EF =0.47. The ele- Mo Mo 5.3. Al-CORBs (red shales) ment Ba is comparable with the Average Shale (Wedepohl, 1971) with EF ranging from 0.51 to 4.71, with one exception at the Vispi Ba 5.3.1. Major elements Quarry where Ba is highly enriched with EF =19.75. The element Ba The major elements Fe, K, P, and Ti from red shales are relatively U varies from place to place, being enriched at Pia da Stua, Italy similar to the Average Shale (Wedepohl, 1971) with EF ,EF,EF, (EF =6.90) and depleted at Tianba, Rehkogel, Brandenberg, and Fe K P U EF (Fig. 14, Appendix Table 4). Mg is enriched in red shales from ODP1049C with EF ranging from 0.05 to 0.38. Ti U Pacific ODP 800A. Na is depleted in red shales from the Silesian In comparison with the Bonarelli black shales (Turgeon and Nappe of Poland and Marzak of the Czech Republic. Mn is depleted Brumsack, 2006), the white limestones and red limestones in the in red shales from the Silesian Nappe of Poland and Albian ODP1049C. Vispi Quarry are enriched in Co, depleted in V, Zn and U (Fig. 15). De- The red shales at Marzak have about two times higher Fe O con- trital input-sensitive elements (Al, Ti, K, Mg, Rb, Zr) show no signifi- 2 3 tent than the gray shales, and Mn concentration tends to be higher in cant difference between red and white limestones. However, the the gray shales than in the red shales (Jiang et al., 2009; Fig. 16). Sim- redox-sensitive elements V, Cr, Ni, and U, and the micronutrient ele- ilar trends can be found in the gray and red calcareous shales from the ments Cu, Zn are slightly enriched in the white limestones (Fig. 15 Albian ODP1049C with about three times higher Mn concentration in and Appendix Table 4). the gray shales than in the red shales (Chen, 2008; Fig. 16).

5.2.3. Rare earth elements The ΣREE in the red limestones average 24.19 ppm (17.02– 5.3.2. Trace elements 41.07 ppm, n=9) in the Vispi Quarry section, 71.23 ppm (66.57– The element Cu is depleted in the red shales from Chuangde

76.58 ppm, n=4) in the Aptian ODP1049C, and 177.56 ppm in the (Tibet) (EFCu =0.45). The element Sr is depleted in the red shales Chuangde section (94.23–299.64 ppm, n=6; Appendix Table 6). The from the Silesian Nappe (EFCu =0.02) and from Marzak (EFCu = ΣREE are strongly negatively correlated (R2=0.876) with carbonate 0.26). The element Mo is depleted at Baergang and Chuangde with content (Fig. 17), because the marine carbonate phase generally contains an EFMo of 0.36–0.39. 236 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248

kaolinite/ Fe(cps) kaolinite/ chlorite quart+illite 1000 2000 0.6 0.8 0.2 0 Gg.subcarinatus Pg.palpebra fine fraction Ht.globulosa benthic

100

150 0 1.5 6000 10000 0 40 -1.0 0.0 2.0 2.5 0.1 -2.0 13 18 Ca(cps) Ti(cps) K-spar/ (warmer) δ O (cooler) δ CVPDB quartz+illite VPDB

Fig. 22. Moving average curves of elemental abundance (Ca, Fe, Ti), mineralogic composition variation, and stable carbon and oxygen isotopic results of foraminifera and ﬁne fraction from the section 16R-2 (Maastrichtian) in ODP Hole 1050C. Note abrupt changes at reddish intervals. Revised from Macleod et al. (2001). (For interpretation of the references

to color in this ﬁgure, the reader is referred to the web version of this article.) ygolohtiL

mhtpe -2

( -6 3 -1 SiO (%) MS (10 m kg ) Hematite (%) 2 Ba/Al (10 ) 20 40 020 0 0.5 1.0 0 0 4 8

139.3 naiblAylraE 140.3

enoZihcsir.H 141.3

142.3

143.3

0 0.5 0 10 20 0 40 80 CaCO (%) Goethite (%) Al 2 O 3 (%) 3

Fig. 23. Comparison of magnetic susceptibility and CaCO3 content (data from Han et al., 2008) from the Aptian–Albian ODP1049C with concentrations of hematite and goethite minerals (data from Li et al., 2011) and element values of Al2O3, SiO2 and element ratio of Ba/Al (data from Chen, 2008). MS: magnetic susceptibility. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 237

The element U is strongly depleted in the red shales compared to contents are 0.74–7.28% (Baudin et al., 1998). A 2.5 m thick strati- the gray shales (Fig. 16). The element Ba is more enriched in the graphic interval separates the Selli Level and ORB 1a (Fig. 19) and is gray shales than in the red shales from the Albian ODP1049C and lithologically characterized by bioturbated gray cherty limestones Chuangde (Fig. 16). and marly limestones. ORB1 is over 15 m thick and is dominated by dark red marlstones and marly limestones with subordinate gray 5.3.3. Rare earth elements marlstones and marly limestones, in beds 1- to 30-cm-thick. The The ΣREE in the red shales averages 450.22 pm at Chuangde, gray beds are locally interbedded with red beds in a cyclical manner. 127.46 ppm in the Albian ODP1049, and 128.57 ppm at Marzak If we apply the 2.6 m/Ma sedimentation rate for the interval from the (Appendix Table 6). The PAAS-normalized REE patterns of the base of M0 (35.6 m) to the top of Chiastozygus litterarius nannofossil Cretaceous red shales exhibit seawater-like REE distribution patterns, zone (40.8 m) of Coccioni et al. (1992), it took about 0.95 Ma for similar to those of the red limestones (Fig. 18). the change from the top of the Selli Level to the deposition of the Cretaceous oceanic red shales show negative δCe excursions with an ﬁrst reddish limestone of the ORB1 (Hu et al., 2006a). average of 0.69 of δCe, ranging from 0.54 to 0.77 (Appendix Table 6). We compiled our analyzed data with those of Stein et al. (2011) However, the red shales from Marzak show positive δCe excursions from the Gorgo a Cerbara section (Fig. 19 and Appendix Table 7). with an average of 1.34 of δCe (Appendix Table 6). The Selli Level is equal to the carbon isotopic stages of C3 to C6 of Menegatti et al. (1998) (Fig. 19). The transitional interval from OAE 5.4. Summary 1a to ORB1 is within the carbon isotopic stage C7 (Appendix Table 7). The ﬁrst occurrence of ORB1 is also within the C7 stage but a few meters (1) Compared to the Average Shale (Wedepohl, 1971), the above the base of ORB1, where the carbon isotopic values start to de- redox-sensitive element V is depleted in the red limestones crease as represented by the C8 stage. from Baergang and Rehkogel. The elements Mn, Ba, P, and Mo are highly enriched at the Vispi Quarry, probably indicating a 6.1.2. Yenicesihlar section, Turkey higher palaeoproductivity compared to the Average Shale The Yenicesihlar section documents the stratigraphic and (Wedepohl, 1971). Compared with the white limestones, the palaeoenvironmental changes from OAE 1a to ORB1 during the Aptian

red limestones have a higher Fe2O3 content. Detrital input- (Yilmaz, 2008; Hu et al., 2012b)(Fig. 19). The OAE 1a interval is ap- sensitive elements (Al, Ti, K, Mg, Rb, Zr) show no signiﬁcant proximately 2.1 m thick and consists of black to dark gray shales difference between the red and white limestones. Redox- with gray marlstones. TOC values of the black shales are up to sensitive elements (V, Cr, Ni, and U) and micronutrient ele- 2.05%. The high organic carbon content and pyritization in the black ments Cu, Zn are depleted in the red limestones indicating a shales of OAE 1a indicate an anoxic environment. The OAE 1a–ORB1 more oxic condition at the sediment–water interface. transitional interval (~20.3 m thick) displays an alternation of light (2) Red shales are similar to the Average Shale (Wedepohl, 1971) gray limestones with very thin-bedded gray marls. ORB1 is approxi- in terms of major and trace elements. Compared to the gray mately 3.5 m thick and consists of pinkish to light brownish lime- shales, the red shales from the Marzak section and Albian stones, which were deposited in an oxic environment, as indicated

ODP1049C have higher Fe2O3, lower Mn and lower Ba contents by hematite in the reddish limestones (Hu et al., 2012b). (Chen, 2008; Jiang et al., 2009). The element U is strongly de- The carbon isotopic record shows several perturbations in the OAE pleted in the red shales compared with the gray shales. Detrital 1a–ORB1 transitional interval including a negative excursion (C3, input-sensitive elements (Al, Ti, K, Mg, Rb, Zr) show no signif- 0.58 m) in the lower part of the Selli-equivalent black shales, a step- icant difference between the red and gray shales indicating a wise positive excursion (C4 to C6, 1.52 m) in the Selli-equivalent stable terrigenous input. The color change between the gray black shales, a positive carbon isotopic plateau (C7, up to 9.48 m) in or white and red shales was caused most probably by the the lower part of the transitional interval, and a carbon isotopic de- variation of redox bottom conditions in the deep ocean crease in the upper part of the transitional interval (C8, up to (Wang et al., 2005; Hu et al., 2006a; Jiang et al., 2009). 10.86 m; Fig. 19; Hu et al., 2012b). The OAE 1a to ORB1 transition in- (3) The differences in the ΣREE, δCe values as well as the PAAS- terval corresponds to carbon isotope stages C7 and C8, which lasted normalized REE patterns either between the red and white approximately 1.3 Ma (Hu et al., 2012b). This duration is very close limestones or between the red and gray shales are minor (Hu to the duration of the OAE 1a (1.1 to 1.3 myr). The C7 stage of positive et al., 2009a; Jiang et al., 2009). carbon isotopic plateau persisted for ~650 kyr, and the negative carbon isotopic shift of the C8 stage also lasted for ~650 kyr (Hu et al., 6. OAE–ORB transition and stable carbon isotopes 2012b).

Temporal–spatial distribution shows that CORBs experienced two 6.2. OAE 2–ORB 9 transition palaeogeographical expansions (Wang et al., 2009): (1) in the Aptian nannofossil zone CC7 shortly after OAE1a and (2) in the Turonian 6.2.1. Umbria–Marche, Italy zone CC11 after OAE2, respectively. To better understand the rela- In the Vispi Quarry and Moria sections, about 10.35 m above the tionship between OAEs and CORBs, we have studied two stratigraphic Bonarelli bed, white limestones of the Scaglia Bianca grade into a se- intervals in the Tethys realm in detail. First we will summarize the quence of predominantly pinkish–reddish limestones and marls of OAE 1a–ORB1 transitions at the Gorgo a Cerbara section (Italy) and the Scaglia Rossa Formation (Hu et al., 2006a; Trabucho-Alexandre the Yenicesihlar section (Turkey) and second the OAE 2–ORB9 transi- et al., 2011; and references within) (Fig. 20). Below the reddish tions in the Vispi Quarry (Italy) and the Buchberg section (Austria). limestone succession are two transitional beds, the lower at 10.35– 10.55 m and the higher at 10.63–10.83 m. Both transitional beds are 6.1. OAE 1a–ORB1 transition whitish in color at the base and gradually become increasingly pinkish color towards the top. The strata above 10.83 m are pinkish, 6.1.1. Gorgo a Cerbara section, Italy to reddish. If we apply the 7.4 m/myr sedimentation rate of the In the Gorgo a Cerbara section, OAE 1a (the Selli Level) is approx- Helvetoglobotruncana helvetica zone of Premoli Silva and Sliter imately 1.92 m thick and consists of laminated to bioturbated gray to (1994), it took about 1.5 myr for the change from the top of the black mudstones and shales, with 20 thin (1 to 3 cm) grayish silty/ Bonarelli Level to the predominantly red colored limestones of the sandy layers (Coccioni et al., 1992); and total organic carbon (TOC) Scaglia Rossa Formation (ORB9) (Hu et al., 2006a). 238 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248

Near the Cenomanian–Turonian boundary, a positive δ13C excur- sediment–water interface regains oxic conditions due to either great- sion was documented globally (Jenkyns, 1980; Schlanger et al., er content of dissolved oxygen or low productivity, organic carbon 1987; Tsikos et al., 2004). From the Early to Middle Turonian, the burial during the ORBs would greatly decrease. The gradual δ13C values stay relatively ﬂat in a so-called carbon isotopic plateau diminishing of organic carbon burial could result in a gradual de- with several slight positive or negative excursions including the crease in carbonate δ13C values, as more 12C-bearing organic matter Holywell, Lulworth, Round Down, and Low-woollgari events (Jarvis was delivered to the hydrosphere–atmosphere reservoirs and was et al., 2006; Fig. 20). Just above the Low-woollgari event (Middle not preserved in the sediments, which would be recorded in the Turonian), the δ13C values fall steeply becoming the Glynde event at coeval carbonates. the minimum value (Fig. 20). In the Vispi Quarry section, red limestones occur around the Round Down event (Fig. 20). The steeply fall- 6.4. Model for the OAE–ORB transitions ing δ13C values start around 2.5 m above the Low-woollgari event (Fig. 20). Wang et al. (2011) proposed a hypothesis that deposition of the CORBs was a possible consequence of Cretaceous OAEs. Enhanced 6.2.2. Buchberg section, Austria amounts of organic carbon and pyrite burial during the OAEs would

The Buchberg section exposes a Cenomanian–Turonian succession have resulted in a large and abrupt fall in atmospheric CO2 concentra- of limestones, marly limestones and marlstones with Cenomanian stra- tion (Arthur et al., 1988). For example, the OAE 2 would lead to a pro- ta being not fully exposed (Neuhuber et al., 2007; Wagreich et al., 2009; found decrease (26% to 40–80%) in atmospheric pCO2 (Kuypers et al., Wendler et al., 2009). Light gray limestones with dark gray mottles al- 1999; Barclay et al., 2010). The decrease in pCO2 probably induced ternate with medium and dark gray spotty marlstones at the base of significant global climatic cooling. But, it is still not known for how the section. They are overlain by gray and red limestones and marl- long the decrease in pCO2 and global climatic cooling would last as stones of Turonian age; pink marl is 2 to 3 m above beds of OAE 2 and a result of enhanced burial of organic carbon. Recently, a remarkable red limestone is 5 m above (Wagreich et al., 2009). This interval repre- large cooling (5–11 °C) was discovered within OAE 2 interval, which sents a duration of approximately 1.1–1.5 myr (Neuhuber et al., 2007; was interpreted to be a response to the associated pCO2 drop Scott, 2009; Wendler et al., 2009). In the lower part of the profile, the (Sinninghe Damsté et al., 2010). They argued that the response of red color is restricted to marl beds whereas in the upper part of the sec- pCO2 drop and climatic cooling to enhanced burial of organic carbon tion both marl and limestone beds are red (Wendler et al., 2009). The during OAEs are probably concurrent and relatively short in time. sediment accumulation rate of the red beds was about 3 m/myr Another view focused on the relatively long time after the OAEs. For (Scott, 2009). example, the Late Aptian after the OAE 1a and the Middle to Late The stable carbon isotopes from the Buchberg section display a Turonian after OAE 2 are in fact considered two cold snaps during broad plateau in the interval of 0–4.8 m with maximum values of the greenhouse Cretaceous. Direct evidence for the Late Aptian global +2.7‰ and decrease to a minimum of +1.9‰ at 5.4 m (Wendler et cooling is indicated by occurrences of glendonites and ice-rafted de- al., 2009; Fig. 20), where the general shape of the stable carbon iso- bris in the Canadian Sverdrup Basin, the West Svalbard Basin and topes is comparable to the Early to Middle Turonian isotope records in southeast Australia (Frakes et al., 1995). Isotopic evidence for a in the other regions (Fig. 20). cool Late Aptian includes data from south temperate belemnites, glendonites (De Lurio and Frakes, 1999) and carbon isotopes 6.2.3. Onset of the ORB9: an Early to Middle Turonian global event? (Weissert and Lini, 1991). Recently, the Late Aptian global cooling The ORB9 occurred after the OAE2 black shales in a time span of was suggested by a decline of Tethyan calcareous nannofossils and a 1.1–1.5 Ma in Italy and Austria (Hu et al., 2006a; Neuhuber et al., subsequent biogeographic expansion of species of high latitudinal af- 2007). In detail, the first reddish or pinkish beds occur at 590 kyr in finities (Mutterlose et al., 2009). Another “cool” period during the Italy (Hu et al., 2006a) or about 700 kyr in Austria following OAE 2 Cretaceous is the Middle to Late Turonian, with evidence from a pos- (Wagreich et al., 2009). In the eastern Atlantic, CORB deposition itive oxygen isotopic excursion in marine limestones (Voigt and began about 670 kyr following OAE 2 in ODP 641A (Scott, 2009). Wiese, 2000), brachiopods (Voigt et al., 2004) and foraminifers Wagreich et al. (2009) suggested that the coeval onset of Turonian (Bornemann et al., 2008). An incursion of boreal faunas (ammonites, red oxic sedimentation both in deep-ocean like the North Atlantic and echinoids and inoceramids) into lower latitudes (Wiese and Voigt, in the basinal and slope environments above the CCD of the Tethys 2002), together with data from organic geochemical proxy (TEX86) may indicate a widespread or even global change in the climate– (Forster et al., 2007), and data from the Late Turonian sequence on ocean system following the OAE2 event. the eastern U.S. continental shelf and elsewhere represent sea level fall of 25–30 m that may be related to glacio-eustasy (Miller et al., 6.3. Summary on the stable carbon isotopes during the OAE–ORB transition 2003). Moreover, Bornemann et al. (2008) suggested that short periods of glaciation existed during Middle Turonian, with ice sheets The stable carbon isotope signature of the OAE 1a–ORB1 transition of about half the size of the modern Antarctic ice cap. However, other shares some similar characters with the OAE 2–ORB9 transition factors such as tectonics could have lead to sea-level fall. We also note (Fig. 21). The δ13C values show a prominent carbon positive excur- that one of the possible mechanisms for global cooling after the OAEs sion during the OAEs, they either stay in an isotopic plateau (the C7 would be the general reduction in seafloor volcanic activity (Larson, stage of Menegatti et al., 1998 or Low-woollgari–Glynde transition 1991). of Jarvis et al., 2006), or return to the pre-OAE values (Fig. 21). The Global cooling would have enhanced formation of cold deep abrupt decrease of δ13C above OAE 2 is about 2.5 m stratigraphically water, increasing its oxidizing capacity due to the greater content of above the first occurrence of the ORBs at the Vispi Quarry section. In dissolved oxygen. Furthermore, enhanced organic carbon burial dur- the Buchberg section, the occurrence of the ORBs correlates to the ing the Cretaceous OAEs would result in an equal addition of O2 to 13 13 onset of the abrupt δ C decrease (Fig. 20). The abrupt δ C decreases the atmospheric oxygen reservoir, assuming a CO2-to-O2 photosyn- after OAE 1a occurs a few meters above the first occurrence of the thetic ratio of 1:1 (Arthur et al., 1988, p.716). This would also have in- oceanic red beds at the Gorge a Cerbara section. At the Yenicesihlar creased the oxidation potential of oceanic deep water. Therefore, section, the ORBs correlate to the C9 stage with carbon isotopic min- an increase in pCO2 and global climate cooling with increased cold imum values (Fig. 19). A scientific question is: Does the occurrence deep-water formation resulting from organic carbon burial during of the ORBs contribute to the abrupt δ13C decrease following the the OAEs would have increased the oxidizing capacity of oceanic OAEs? We believe that during the formation of the ORBs, when the deep water and promoted formation of the ORBs. As stated above, X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 239 the transition from OAE 1a to ORB1 during the Aptian took a relatively palaeoproductivity proxy (e.g. Dymond et al., 1992), and is higher in short time of 0.95–1.2 Ma in western Tethys (Italy), and 1.3 Ma in the the white beds than in the red beds. Thus the brown beds may have Middle Tethys (Turkey). The transition from OAE 2 to ORB9 took resulted from low primary productivity that may have produced an about 1.1 Ma in western Tethys (Italy and Austria). This model pro- oxic environment during deposition. vides an effective mechanistic explanation for the development of CORBs, but it remains speculative pending further data collection. 7.2. Model 7. Orbital-cycle control CORBs on Blake Nose, North Atlantic During the deposition of brown beds, low primary productivity would result in low organic carbon flux into the sediments, which Besides long-term CORB deposition such as in the Scaglia Rossa, would reduce oxygen demand and produce an oxidizing early diage- highly cyclic ORBs have received less attention. In recent decades, netic condition (Fig. 24A). In such an oxic environment, iron oxides high-frequency cycles consisting of CORBs have been widely recog- will form. Low primary productivity is expected to be associated nized in the Tethyan Ocean, North Atlantic, South Atlantic, Indian with low CaCO production and low Ba content, as well as high con- Ocean, and Pacific Ocean (Macleod et al., 2001; Hu et al., 2006a; 3 tents of clay minerals and the terrigenous-related elements of Al, Si, Neuhuber and Wagreich, 2009; Wang et al., 2009; Wendler et al., K, and Na, which represent high terrigenous input. During the depo- 2009; Li et al., 2011). Below we describe this type of CORB cored on sition of white beds, low terrigenous input and relatively high prima- the Blake Nose of the North Atlantic as an example. ry productivity are expected to be associated with high CaCO3 and Ba contents and low contents of clay minerals (Fig. 24B). In such a con- 7.1. Data dition, organic carbon flux into the sediments would be relatively high, which would result in relatively reducing conditions, thus nei- Cycles 30–50 cm thick in ODP1050C (Blake Nose) Maastrichtian ther hematite nor goethite would form. core (Core 16-2R) are defined by gradational color alternations between light grayish green and reddish brown (Fig. 22), which were interpreted as representing the ~21 kyr precessional signal (Norris 8. CORBs and the Cretaceous ocean and climate et al., 1998). Macleod et al. (2001) found that high planktonic δ18O values in the green intervals are correlated with high planktonic but Lower Cretaceous (Berriasian–Barremian) ORBs are known from low benthic δ13C values (Fig. 22), indicating that cooler and/or more rare outcrops (Greenland, Northern Calcareous Alps of Austria, saline surface waters were associated with higher productivity. The Cismon of northern Italy, and in the eastern Indian Ocean only from green intervals with higher productivity are also characterized by site DSDP 261). Generally, the Early Cretaceous was regarded as a rel- high CaCO3 concentration, as well as enrichment in feldspar and kao- atively cool climate possibly with transient glaciations (Price et al., linite (Fig. 22). Conversely, the reddish intervals are interpreted to 2000). Recently, the TEX86 palaeotemperature proxy suggested that have low-productivity, and together with low planktonic δ18O values Early Cretaceous sea-surface temperatures exceeded 32 °C at 15°– indicate warm and (or) less saline surface waters, high Fe and Ti con- 20°N, 26 °C at 53°S (Littler et al., 2011) and 26°–30 °C at 60°–70°S centrations, and enrichment in quartz, illite, and chlorite (Fig. 22). (Jenkyns et al., 2011). These temperatures substantially exceed mod- Potential forcing mechanisms for covariation of these properties ern temperatures at equivalent latitudes, and are incompatible with found in the Maastrichtian ODP1050C include cyclic variation in the notion of cooler conditions in the earliest Cretaceous. water column stratification, variation in continental nutrient and de- Mid-Cretaceous (Aptian–Cenomanian) ORBs are well exposed in the trital fluxes, and changes in the intensity of upwelling (Macleod et al., remnants of the Tethys Ocean, Atlantic Ocean and Pacific Ocean as well 2001). as in the Indian Ocean (Wang et al., 2009). After the Early Aptian OAE1a, The Albian sediments encountered on the Blake Nose in ODP red or brown clays, marls, and limestones became more geographically 1049C (Core 12X) in the Microhedbergella rischi planktonic foraminif- widespread. Consequently Late Aptian is thought to mark the onset of eral zone (e.g. Huber and Leckie, 2011) are characterized by global CORB deposition. Mid-Cretaceous CORBs commonly occur as dis- high-frequency cycles consisting of oceanic, brown and white calcar- crete thin beds (named ORB1 to ORB8) with short durations of hun- eous claystone. This interval is composed of five cycles of brown and dreds of thousands years (Hu et al., 2006a), except for ORB1, which white beds that are based on magnetic susceptibility (Li et al., 2011). was 5.12 myr in duration in central Italy. The greenhouse climate of The thicknesses of the cycles (from top to bottom) were 580, 420, the mid-Cretaceous was likely related to major global volcanism and as-

650, 320, 310 mm, which correspond to cycle durations of 99, 71, sociated outgassing of CO2. OAEs may be recognized as a negative feed- 111, 54, 53 ka, respectively. These durations are based on a sedimen- back in response to sudden warming episodes preventing further tation rate of 5.88 mm/ka for the Albian part, which is derived from acceleration of warming through the removal of organic carbon from palaeomagnetic data and astronomical cycles (Ogg et al., 1999). The the ocean–atmosphere (CO2), via increased productivity cycle, with or- lowest durations are comparable to the 53.6 ka cycle of the weak ganic carbon being buried in the deep ocean. However, the presence of obliquity of the Earth's axis, and the highest are comparable to the oxic sediments (CORBs) in transition zones between individual OAEs 85–140 ka cycle of the short eccentricity of Milankovitch cycles. indicates development of longer cooling-climate periods, and therefore Based on the study of the Albian sediments in the Core 12X of Hole sharp swings in the mid-Cretaceous climate. Hay (2008) proposed that 1049C by Han et al. (2008) and Li et al. (2011), we summarized charac- the ocean structure would be very different during the Cretaceous if ters of the brown–white cycles as follows (Table 2): (1) clay mineral there were no ice at the poles. He stated that the only steady ocean cur- contents as well as Al2O3 and SiO2 contents are higher in red than in rents would have been those of the equatorial systems, forced by the white beds (Fig. 23). This may indicate that terrigenous input was constant easterly wind. At mid and high latitudes the ocean would stronger in the brown beds than in the white beds. (2) Albian brown have been ﬁlled with eddies. Without polar ice, about half of the entire beds contain an average hematite value of 0.51% and an average goe- ocean surface would have almost the same density, and that the density thite value of 0.58% (Fig. 23). The white beds contain neither hematite of the ocean surface would approximately be the same as the density of nor goethite. The color of ORBs from the Albian ODP1049C is mainly the deeper ocean. Under these conditions, storm and wind-generated controlled by hematite that imparts a red color to sediment and goe- eddies could pump water up and down. Through the changes in ocean thite that imparts a bright yellow color (Li et al., 2011). (3) CaCO3 con- salinity by evaporation and precipitation in different climate modes, tents are higher in the white beds (58% on average) than in the brown the ocean circulation may be changed accordingly, which could explain beds (21% in average, Fig. 23). The Ba/Al value is considered to be a the frequently changing redox conditions during the mid-Cretaceous as 240 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248

Table 2 Comparison between the brown beds and white beds in the Albian part of the ODP1049C.

Content Brown bed White bed Interpretation

CaCO3 (%) Low, 21 High, 58 Terrigenous inputs are stronger in the red beds than those in the white beds Clay mineral (%) High, 25 Low, 8.3 Total iron (%) High, 5.23 Low, 1.02 Hematite and goethite are the minerals imparting the brownish color of Iron oxides (%) High, 1.04 Low, 0.02 sediments, forming in oxic condition. Hematite (%) High, 0.13–0.82 None Goethite (%) High, 0.22–0.81 None Magnetic susceptibility (10−6 m3 kg−1) High, 26 Low, 4.8 P enrichment factor Low, 0.66 High, 1.88 Low primary productivity was suggested in the brown beds compared to Ba enrichment factor Low, 2.37 High, 6.06 the white beds. δ18C(b45 μm whole rock) High Low Brown beds were formed in warm conditions compared to the white beds. U enrichment factor Low, 0.33 High, 2.47 Brown beds were formed in a high oxic condition, while red beds were in Mn enrichment factor Low, 0.37 High, 3.10 a normal oxic condition. V, Co, Ni, Cu Low High evidenced by the repeated changes from black shales to red beds in the basins organic-rich muds and chalks were deposited and are now im- deep ocean. portant source rocks, e.g. Austin Chalk, Niobrara Chalk, Eagle Ford/ The Upper Cretaceous (Turonian to Maastrichtian) ORBs were Boquillas Shale, and Maastrichtian in the Gulf of Suez, among others. globally deposited in the Tethyan, Indian, Atlantic and Pacific oceans, even much wider than the mid-Cretaceous CORBs. The global occur- 9. Perspective: Phanerozoic oceanic red beds rence of the Upper Cretaceous ORBs indicated that bottom waters were well-oxygenated with high dissolved oxygen, as supported by 9.1. Distribution of oceanic red beds in the Phanerozoic Eon morphologic characteristics of benthic foraminifers in the global ocean (Kaiho, 1994). Besides well-oxygenated deep water, several Oceanic red beds have been found in nearly every period from the other processes may have contributed to the formation of CORBs. Cambrian to the Quaternary as mentioned as early as Galloway During the Late Cretaceous, bio-available phosphorus was at a low (1922). We compiled the distribution of the ORBs (Ca- and Al-rich level (Főllmi, 1996), which may have resulted in lower productivity. oceanic red beds) mainly based on the published data and our own Most of the organic carbon exported from the photic zone to the field investigation from China (Fig. 25 and Table 3). deeper waters was oxidized and only a small amount was buried. The oxic, lower productivity ocean can be further explained by the 9.1.1. Red shales global climatic–oceanic change that occurred during the Late Creta- (1) Rhythmically bedded red and green claystones, siltstones, ceous Epoch. With global cooling (Huber et al., 2002) from a “hot-” sandstones and conglomerates in a Cambrian flysch in Quebec to a “cool-” greenhouse climate during Turonian to Campanian Appalachians of Canada (Lajoie and Chagnon, 1973) are (Clarke and Jenkyns, 1999), the equator-to-pole temperature gradi- 12–90 m thick. The green hemi-rhythms are 7–150 m thick, ent increased, progressively favoring the formation of large volumes and are composed of claystone with minor siltstone and sand- of deep water at high-latitudes, particularly along the Antarctic mar- stone beds. gin (Otto-Bliesner et al., 2002), and increasing the turnover rates of (2) The Lower Cambrian marine Caerfai Bay Shale is a volcaniclastic the ocean. Opening of the deep connection between the North and red-bed horizon in Dyfed, Wales (Turner, 1979). The red colora- South Atlantic Oceans in the Campanian allowed deep waters gener- tion was suggested to have been produced entirely by a diage- ated in the southern hemisphere to circulate throughout the Atlantic netic alteration of iron-silicates and volcanic ash in an oxidizing basins (Poulsen et al., 2001). Furthermore, widening and deepening diagenetic environment (Turner, 1979). of the passages between Antarctica, Africa, and India, and between (3) Lower Silurian (Middle Llandovery to Lower Wenlock) red shales India and Australia allowed deep waters generated along the Antarc- and siltstones are widely distributed in the UK, Norway, Ireland, tic margin to enter the Eastern Tethys Ocean (Hay et al., 1999). The Estonia, Latvia, and Appalachian basin (Ziegler and McKerrow, cooler high latitude deep waters introduced more oxygen to the 1975) where they were mainly deposited in offshore shelf marine deep ocean, so that more organic matter was oxidized, and iron ox- environments. In south China (Rong et al., 2012)redshales ides formed at the sediment–seawater interface. However, in some were deposited in shallow marine environments. Ziegler and

warm and humid high irradiation cold and relative dry low irradiation strong terrestrial weathering weak terrestrial input red bed white bed sea level sea level

low productivity sediment- water high productivity 2+ low CaCO3 sediment- water Fe +O2 Fe2 O3 interface low content of 2+ O2 Fe O2 dissolved oxygen interface high terrestrial input O2 O2 high content of O2 rich in clay minerals O2 weak reducting dissolved oxygen environment O2 O2 high CaCO3 O2 O2 depleted V, Co, Ni, Cu, Cr slightly enriched V, Co, Ni, Cu, Cr

high total Fe content low EFUM and EF n high EFUM and EF n Fe present in form of ferric low EFPB and EF a high EFPB and EF a oxides (goethite and hematite) very low TOC low TOC A B

Fig. 24. Conceptual models for the Albian red (A) and white (B) beds in ODP1049C. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.) X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 241

A B

C D

E F

Fig. 25. Field photos of Phanerozoic Eon oceanic red beds. (A) Red limestone of the Guniutan Formation, Middle Ordovician, Shitai County, Anhui Province, China; (B) red nodular limestone of the Podona Formation, Upper Ordovician, Jiange County, Sichuan Province, China (photo by Prof. Dr. Mingcai Hou); (C) red limestone with stromatactis (sponge cavities), Frasnian (Upper Devonian), Les Bulants quarry, Neuville, Belgium (photo by Prof. Frédéric Boulvain); (D) red nodular limestone in the Lanchengquxia Formation (Lower Triassic), Pulan, western Tibet, China (photo by Dr. Wenzhong Li); (E) red limestone in the Nanlinghu Formation, Olenekian (Lower Triassic), Caohu, Anhui Province, China (photo by Dr. Kuidong Zhao); (F) red nodular limestone of the Ammonitico Rosso Formation (Upper Jurassic) near Gubbio, Italy. (For interpretation of the references to color in this ﬁgure, the reader is referred to the web version of this article.)

McKerrow (1975) suggestedthattheredoxidizedmaterialis 9.1.2. Red carbonates probably second-cycle, derived through coastal erosion of soil or Up to now, at least twelve stratigraphic levels of oceanic red lime- alluvial complexes during marine transgression. They concluded stones are known in the Phanerozoic Eon (Table 3). that “the sediments are red because they were derived from oxidized source materials and were buried rapidly in quiet marine (5) Red Lower Cambrian carbonate buildups in the Flinders environments before they could be reduced.” Obviously this state- Ranges are composed of two facies: low-energy, archaeocyath- ment needs to be re-evaluated by further investigation especially sponge-spicule mud mounds, and high-energy, archaeocyath- in light of modern integrated provenance analysis. calcimicrobe (calciﬁed microbial microfossil) bioherms (James (4) The Upper Devonian (mainly Famennian) red mudstones are and Gravestock, 1990). Mud mounds are composed of red car- a conspicuous facies type of the Variscan basins in Europe bonate mudstone and ﬂoatstone with stromatactis and abun- (Germany and England) (Franke and Paul, 1980). The environ- dant sponge spicules that form structures up to 150 m wide ment was interpreted to be pelagic based on the study of fossils and 80 m thick. Bioherms are either red or dark gray limestone and trace fossils. forming isolated small structures 2–20 m in width. 242 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248

Table 3 Phanerozoic oceanic red beds other than the Cretaceous oceanic red beds.

Lithology No. Period Stage Place Structures Environments Origin suggested References

Red 16 Palaeogene Palaeocene Central Italy Pelagic Oligotrophy, iron oxidation Arthur and Fischer (1977) limestone to Eocene in oxic environment 15 Jurassic Middle to Late Tethyan ocean–Ammonitico Nodula Pelagic Iron oxidation and low Jenkyns (1974); Rosso sedimentation rate or Préat et al. (2006, 2008) iron-bacterial mediation 14 Jurassic Middle Pieniny Klippen Belt Stromatactis Mud mound Not discussed Aubrecht et al. (2002) (Western Carpathians, Slovakia) 13 Jurassic Early Southern Alps, Switzerland Stromatactis Mud mound Not discussed Neuweiler and Bernoulli (2005) 12 Triassic Early South China and Tethyan Shelf Unknown Tong et al. (2001); Himalaya Guo et al. (1991) 11 Permian Dorashamian Sanandaj–Sirjan Province, Pelagic below Not discussed Heydari et al. (2003) Iran storm wave base 10 Carboniferous Bashkirian–Moscovian Northern Spain–Baleas Lm Stromatactis Shallow to Iron–bacterial mediation van der Kooij et al. (2007) carbonate slope 9 Devonian Frasnian–Famenian Belgium–mud mounds; Stromatactis Shallow to open Iron–bacterial mediation Boulvain et al. (2001) France–Coumiac Lm and marine Griottes Lm 8 Devonian Pragian Czech Republic–Slivenec Lm Stromatactis Shallow to open Iron–bacterial mediation Mamet et al. (1997) marine 7 Silurian Late Quebec Appalachians Stromatactis Shallow marine Iron–bacterial mediation Bourque and Boulvain (1993) 6 Ordovician Late South China and Tarim– Nodular Outer shelf? Unknown Zhan and Jin (2007); Pagoda Formation Zhang (1996) 5 Cambrian Early Flinders Range, Australia Stromatactis Shelf Unknown James and Gravestock (1990) Red shale 4 Devonian Famennian Germany and southwest Pelagic Oligotrophy, iron oxidation Franke and Paul (1980) England in oxic environment 3 Silurian Early UK, Norway, Ireland, Offshore marine; Marine transgression; Ziegler and McKerrow Estonia, Latvia, shallow marine continental input (1975); Rong et al. (2012) Appalachian Basin; South China 2 Cambrian Early Wales, UK Volcaniclastic Marine Oxidizing conditions Turner (1979) (Caerfai Bay Shales) rich 1 Cambrian Canadian Quebec Flysch type Deep marine Rapid resedimentation Lajoie and Chagnon (1973) Appalachians of red sediment

(6) The Middle to Upper Ordovician reddish nodular limestones hemipelagic outer ramp, below the storm wave base. are widely distributed in the south China Yangtze area, from Si- (10) Carboniferous red limestones are distributed in Austria and chuan to Zhejiang provinces and are over 1500 km long (Zhang, Spain and were interpreted to be deposited in an outer ramp en- 1996; Zhan and Jin, 2007)(Fig. 25A). The Guniutang Formation vironment (see van der Kooij et al., 2007). A total of seven red consists of reddish to purple and gray nodular and bedded lime- muddy, or peloidal patches or well-developed layers, rich in or- stones about 20 m thick (Zhang, 1996). The Pagoda Formation is ganic material are present on the slope of the Sierra del Cuera about 10–20 m thick and dominated by light purple–reddish platform (van der Kooij et al., 2007). thin-to-thick bedded limestones (Fig. 25B). One of its typical (11) In the Sanandaj-Sirjan Province of Iran, the 18 m thick upper- characters is the polygonal reticulate structure that is widely de- most Permian is reddish argillaceous limestones with character- veloped in the Pagoda Formation, and its origin is greatly debat- istically nodular texture and pelagic ammonites and conodonts ed. These Upper Ordovician red limestones also crop out in the (Heydari et al., 2003). These red limestones were deposited in southern part of the Tarim Basin, China (Jiang et al., 2001). pelagic, deep-water, below storm wave base. (7) The Upper Silurian red stromatactis limestones in the Gros (12) The Lower Triassic (Olenekian stage, Spathian substage) red thin Morbe Member of the West Point Formation in the Québec Ap- bedded and partly nodular limestones are 10–20 m thick in the palachians (Bourque and Boulvain, 1993) are mound-shaped, Anhui Province, southeastern China (Tong et al., 2001) finely crystalline, crudely bedded. Individual mounds are up to (Fig. 25E). In southwestern Tibet along the Yarlung Zangbo su- 115 m thick. The Gros Morbe facies is interpreted as having ture zone Lower Triassic red nodular limestone (Fig. 25D) is been deposited in quiet water, below storm wave base, and over 100 m thick and was deposited in a deep-water environ- probably below the photic zone (Bourque and Boulvain, 1993). ment (Guo et al., 1991). (8) Deep red Lower Devonian limestones (Mamet et al., 1997)crop (13) Lower Jurassic red stromatactis limestone in Switzerland out near Prague, Czech Republic. Normally they are lenses- (Neuweiler and Bernoulli, 2005) forms a massive sublithofacies shaped and deposited in an inner ramp environment. The red and is enclosed in a mound-shaped, brachiopod-rich bioclastic pigmentation of the limestones was suggested to be bacterial limestone. The thickness ranges from several tens of centimeters in origin by occurrences of hematite filling as mineralized films to several meters and limestone patches are lense- to mound- and microstromatolites (Mametetal.,1997). shaped with gradational contacts. (9) The Upper Devonian red limestones are widely distributed in the (14) A Middle Jurassic red stromatactis mud-mound is in the Belgium Ardennes and France (Boulvain et al., 2001)(Fig. 25C). Czorsztyn Unit of the Pieniny Klippen Belt (Western Carpathians, They mainly consist of condensed red argillaceous micritic car- Slovakia) (Aubrecht et al., 2002). bonates about 20–30 m thick, with many hardgrounds. These (15) Perhaps the most famous red limestone is the Jurassic limestones were interpreted to be deposited on a distal Ammonitico Rosso, which is well known for its nodular X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 243

structure, the presence of ammonites and the world-wide distri- dominantly in the trivalent state and reduction of nitrate in the bution as an ornamental stone. The Ammonitico Rosso facies are pore waters by organic carbon does not occur (Sawlan and Murray, widely distributed in Tethys, such as in Spain, Italy, Switzerland, 1983). The sediments therefore remain brown throughout the length Austria, Czech, Poland, Yugoslavia, Hungary, Bulgaria, Romania, of the sediment column, and elements such as Mn, Co, Cu, and Ni are Morocco, Algeria, Greece, Turkey (Cecca et al., 1992), as well as enriched in the authigenic fraction of the sediment (Glasby, 1991). in the central Atlantic. The Ammonitico Rosso is a stratigraphically condensed pelagic facies and consists of the 9.3. Origin of Phanerozoic marine red beds Middle to Upper Jurassic to basal Cretaceous red pelagic limestones, commonly nodular and rich in ammonites (Martire et During the past decades, numerous authors dealt with continental al., 2006)(Fig. 25F). The Ammonitico Rosso was deposited in a red beds and concluded that the red coloration commonly is due to pelagic deep-marine environment probably near the CCD, breakdown and oxidation of iron-bearing minerals (e.g., van which resulted in the formation of the nodular structure at the Houten, 1973). However, the origin of the ORBs and their geological early diagenetic stage (Jenkyns, 1974). Martire (1996) suggested significance remained a great debate. Galloway (1922) thought that that sediments accumulated at a very slow sedimentation rate some marine red beds resulted by telogenetic (post-sedimentary) su- (~3 mm/ka, even after decompaction and taking into account perficial alteration. He stated that “red limestones were not originally multiple hiatuses). However, Préat et al. (2006, 2011) reported red, but in all cases are due to atmospheric weathering”. A good ex- calcite and ferruginous microfilaments of different shapes and ample is the Lower Carboniferous Redwall Limestone in the Grand sizes from the Ammonitico Rosso in Italy and they attributed Canyon, USA in which the overlying Pennsylvanian red arenites pro- the red color to the presence of hematite (former iron vide iron oxides that superficially cover the face of the outcrops oxy-hydroxides), which resulted from the mediation of iron bac- (McKee and Gutschick, 1969). Obviously this origin model cannot teria at the poorly dysoxic–anoxic sediment–water interface. The apply to the ORBs. Up to now, three different hypotheses have been isotope δ56Fe (Préat et al., 2008) is systematically lighter in the proposed to explain the origin of red pigmentation of limestones red facies than in the gray facies. and shales with the following possible scenarios. (16) It is important to point out that many of the CORBs extend into the Palaeogene, such as in central Italy where the reddish facies 9.3.1. Detrital origin of the iron derived from continental weathering of the Scaglia Rossa range from the Cretaceous to as young as Lajoie and Chagnon (1973) studied rhythmically red and green Late Eocene. Therefore, Palaeogene ORBs are present. claystones in a Cambrian flysch sequence, Quebec Appalachians of Canada, and suggested that the red detritus was derived from an oxidiz- 9.2. Late Cenozoic red clays ing, paralic environment, and was directly and rapidly transported into the basin. Similarly, Ziegler and McKerrow (1975) suggested that a Late Cenozoic red clays, one example of the ORBs, were discovered plentiful supply of oxidized iron from land into sea was the main origin during the Challenger Expedition (1872–1875) and cover approxi- for the red color of Silurian marine red beds in Western Europe and in mately 31% of the world's ocean basins or about 20% of the total sur- the Appalachian Basin. They related this process to the Llandovery ma- face of the Earth (Bryant and Bennet, 1988). Red clays constitute up to rine transgressions. Franke and Paul (1980) suggested that hematite 49%, 26%, and 25% of the sediments of the Pacific, Atlantic, and Indian enrichment of the clay fraction was derived from land and imparted a Oceans, respectively (Anderson, 1986). The red clays in the Pacific red coloration to the Devonian red shales in Germany. They stated: Ocean were studied in detail, being often used as a model for other “As nearly all fine-grained terrigenous sediments contain several per- marine red sedimentary deposits. Generally, the red clays represent centages of ‘free’ ferric iron, the minimum of 1.5% hematite required the residual sediment type formed in deep-sea areas of low biological for red coloration is easily achieved. Hence, all claystones and silty productivity occurring below the CCD where terrigenous input is low. claystones are potentially red.” However, as documented from the

Red clay generally has a grain size of about 1 μm, a CaCO3 biogenic Cretaceous deep-water red shales intercalated with turbiditic beds in content b15%, net bulk density of 1.35–1.40 g/cm3, water content of the Gosau Group in Austria, the red shales represent as a normal back- 125%, porosity of 75–80%, and sedimentation rate of b1 mm/ka ground sedimentation with relatively low sedimentation rates, while (McCoy and Sancetta, 1985; ~0.45 mm/ka by Gleason et al., 2002). the gray shales would be the results of turbiditic currents with high The most abundant mineralogical components of red clays from the sedimentation rates (Wagreich and Krenmayr, 2005). North Pacific Ocean are authigenic and detrital smectite, aeolian shale clasts of illite, kaolinite, and chlorite, and in variable propor- 9.3.2. Iron-bacterial mediation at the time of sedimentation tions, quartz, feldspar, zeolite, and authigenic metal oxides (Bryant This hypothesis involves the mediation of iron bacteria near the and Bennet, 1988). The average mineralogical composition of pre- water–sediment interface during sedimentation and is applicable to Miocene red clays in the northwest Pacific is illite 25%, smectite a number of marine limestones of different ages (see review by 55%, kaolinite and chlorite 9%, quartz 5%, feldspar 1.5%, palygorskite Mamet and Préat, 2006 and reference therein). Iron (II)-oxidation oc- 0.5%, and amorphous material 4% (Bryant and Bennet, 1988). curs in microaerobic or anoxic environments as a result of the activity The most important contribution to the allogenic fraction of mod- of microorganisms that oxidize Fe(II) to generate energy for growth ern red clays in the Pacific is aeolian dust, which is mainly derived (Mamet and Préat, 2006). The products of biologically oxidized iron from Asia, Australia, and Central America (Glasby, 1991). Besides are highly insoluble ferric oxide minerals (normally hematite), aeolian input, authigenic aluminosilicates and volcanic debris were which are sub-micrometric and occur in the matrix of limestones found in the chemically post-extraction, residual component in the (Mamet and Préat, 2006; Préat et al., 2011). Late Cenozoic red clays in the central Pacific(Ziegler et al., 2007). A number of elements are enriched in the Pacific red clays including 9.3.3. Iron oxidation in oligotrophic, highly oxic environment Mn, Co, Cu, Ni, and Fe, which occur as authigenic fractions such as The reddish limestones of Cretaceous Scaglia Rossa in Italy were Fe and Mn oxides (Glasby, 1991). Red clays form in regions of low suggested to be deposited over several decades in an oxic environ- sedimentation rates where rates of organic matter deposition are ment (Arthur and Fischer, 1977). The reddish Scaglia Rossa strata very low (Lyle, 1983). In such circumstances, where bottom water is are regularly bedded and are in conformable contact with the under- under well-oxidizing conditions, diffusion of oxygen into the sedi- lying whitish-colored strata. Within a single red limestone bed ments exceeds its consumption during the degradation of organic the red color of individual limestone beds is generally uniform, matter (Glasby, 1991). Iron is therefore present in the sediments but in other beds the color changes from white to pink and red 244 X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 from bottom to top (Hu et al., 2006a). These characters indicate transformed to hematite during late diagenesis. The red color of that the red color of the Scaglia Rossa most probably resulted from shales from the Chuangde section resulted from hematites of a differ- synsedimentary or early diagenetic processes. Furthermore, the red ent grain size, ranging from several tens of nanometers to several tens Scaglia Rossa limestones are enriched in hematite and depleted of of microns in size. redox-sensitive and nutrient-like trace elements, but the terrigenous Elemental geochemical data indicated that red limestones have a fraction is nearly the same as in white limestones (Hu et al., 2009a). similar composition of detrital input-sensitive elements (Al, Ti, K, Mg,

These factors indicate that the Scaglia Rossa red limestones were depos- Rb, Zr), higher content of Fe2O3, and depleted redox-sensitive elements ited under highly oxic, oligotrophic, and probably low-productivity (V, Cr, Ni, and U) and micronutrient elements Cu, Zn compared to white conditions (Hu et al., 2006a, 2009a, 2009b). Palaeomagnetic limestones, indicating a similar provenance source and a more oxic con- properties of the Upper Cretaceous Scaglia Rossa red pelagic limestones dition at the sediment–water interface. Red shales are comparable to support this interpretation (Channell et al., 1982). The boundaries of theAverageShale(Wedepohl, 1971) in terms of elements. Compared palaeomagnetic reversals deﬁned by hematite are a few tens of centi- to the gray shales, the red shales have higher Fe2O3, lower Mn and meters below those of magnetite. Therefore, they interpreted that lower Ba contents, with similar detrital input-sensitive elements indi- hematite grains were magnetized in the post-reversal ﬁeld, which dem- cating a stable terrigenous input. The color change between the gray onstrates that growth of hematite crystals occurred over ~105 years or white and red shales was caused most probably by the variation of during early diagenesis. redox bottom conditions in the deep ocean. Hematite platelets of the Upper Cretaceous pelagic red sedimenta- Detailed studies of the stratigraphic transitions from OAE1a–ORB1 ry rocks in Turkey are hexagonal in shape and grew epitaxially on the and OAE2–ORB9 lead to the hypothesis that deposition of the CORBs grains, strongly suggesting an early diagenetic origin (Eren and Kadir, was a possible consequence of the Cretaceous OAEs (Wang et al., 1999). A similar origin was suggested for CORBs in other areas (see 2011). Enhanced amounts of organic carbon and pyrite burial during Hu et al., 2005; Wagreich et al., 2009; Wang et al., 2009). the OAEs would have resulted in a large and abrupt fall in atmospher-

Field investigations confirmed that CORBs are not the products ic CO2 concentration, which probably induced significant global cli- either of superficial alteration or detrital origin derived from land (see matic cooling after the OAEs. Global cooling would have enhanced Hu et al., 2009a; Wagreich et al., 2009; Wang et al., 2009). The consis- formation of cold deep water, increasing its oxidizing capacity due tent interlamination of red and gray facies in some CORBs on a to the greater content of dissolved oxygen and would promote forma- meter-to-centimeter scale is incompatible with post-sedimentary tion of the ORBs. As we stated above, the transition from OAE 1a to weathering. As stated by Wang et al. (2009), CORB sedimentation was ORB1 during the Aptian took a relatively short time from 0.95 to in opposition to clastic deep-water sedimentation. The general rule is 1.2 Ma in western Tethys (Italy), and 1.3 Ma in Middle Tethys (Tur- that the greater the clastic input, the higher the sedimentation rate, key). The transition from OAE 2 to ORB9 took about 1.1 Ma in western and thus the decreased probability for CORB deposition (Wagreich Tethys (Italy and Austria). The stable carbon isotope signature of the and Krenmayr, 2005; Wagreich et al., 2009). Clearly some CORBs such OAE1a–ORB1 transition shares some similar characters with the as the Scaglia Rossa are not the products of iron-bacterial mediation be- OAE2–ORB9 transition. Generally, the δ13C values show a prominent cause no evidence was found. However, we are very careful to say that carbon positive excursion during the OAEs, they either stay in an iso- iron-bacterial mediation may occur in some CORBs, and more work topic plateau, or return to the pre-OAE values. We suggest that the oc- needs to be done before we can answer this question more surely. currence of the ORBs would contribute to the abrupt δ13C decrease following the OAEs, as in oxic conditions during the formation of 10. Conclusion the ORBs, organic carbon burial would be greatly decreasing. Data from the ODP Leg 171B (Holes 1049C and 1050C) indicate CORBs are reddish to pinkish to brownish sedimentary rocks de- that during the deposition of the Albian and Maastrichtian reddish in- posited in pelagic marine environments (Hu et al., 2005; Wang et tervals, low primary productivity and relatively high temperature al., 2009). Geochemically, the CORBs can be classified into three end would result in low organic carbon flux into the sediments, which members: Ca-CORBs, Al-CORBs, and Si-CORBs by CaO, Al2O3 and would reduce oxygen demand and produce an oxidizing early diage- SiO2 values. Ca-CORBs have CaCO3 contents of 48–95%, Al2O3 values netic condition. In such an oxic environment, iron oxides will form. of 1–10% and SiO2 value of 3%–30%. Al-CORBs have CaCO3 contents Low primary productivity is expected to be associated with low of 1–22%, Al2O3 values of 7–20%, SiO2 values of 42–70%. Si-CORBs CaCO3 production and low Ba contents, as well as high contents of have CaCO3 contents of 0.2–1.4%, Al2O3 values of 1–4%, SiO2 values clay minerals and the terrigenous-related elements of Al, Si, K, and of 86–96%. Outer shelf to upper/middle bathyal CORBs are mainly Na, which represent high terrigenous input. red pelagic limestones (Ca-CORBs). Towards deeper bathyal depths, ORBs were found in nearly every period from Cambrian to Quaterna- red marls and marly shales become the dominant lithology/facies ry. We compiled the distribution of the known ORBs (5 localities of red type. Below the calcite compensation level, in abyssal depths, either shales including the Late Cenozoic red clays and 12 localities of red red claystones and carbonate-free shales (Al-CORBs) or red cherts limestones, to our knowledge), which indicated that the ORBs are not (Si-CORBs) are deposited. only found in the Cretaceous, but the whole Phanerozoic. Different hy- Chronostratigraphic data of numerous sections show that two potheses can be considered to explain the origin of the red pigmenta- groups are distinguished by their durations. Short-term CORBs gener- tion of limestones and shales including (1) detrital origin of the iron ally are less than 1 myr in duration and most long-term CORBs depo- derived from continental weathering; (2) iron-bacterial mediation at sition lasted longer than 4 myr. Many of the orbital-cycle control the time of sedimentation; and (3) iron oxidation in oligotrophic, highly CORBs approximate the duration of a climate frequency response to oxic environment. Abundant field and laboratory evidence support the the Milankovitch cycles. conclusion that CORBs were deposited under highly oxic, oligotrophic Hematite, goethite and Mn2+-bearing calcite are the minerals that conditions with probably a low sedimentation rate. impart the red color to CORBs. Hematite and goethite are responsible Supplementary data to this article can be found online at http:// for brown and orange colors from the Albian–Aptian sediments from dx.doi.org/10.1016/j.earscirev.2012.09.007. the ODP 1049C. Hematite clusters in several to tens of nanometers in the calcite structure is the main reason for the red coloring of lime- Acknowledgments stones, and manganese-bearing calcite may give additional red color to limestones in the Vispi Quarry. Goethite was suggested to be We thank Mr. Timothy Horscroft from Elsevier for his invitation originally formed together with hematite, but was subsequently and encouragement. Special thanks are given to our Cretaceous X. Hu et al. / Earth-Science Reviews 115 (2012) 217–248 245

IGCP colleagues (Willam W. Hay, Luba Jansa, Wolfgang Kuhnt, Channell, J., Freeman, R., Heller, F., Lowrie, W., 1982. Timing of diagenetic haematite growth in red pelagic limestones from Gubbio (Italy). Earth and Planetary Science Stephanie Neuhuber, Massimo Sarti, Petr Skupien, Michael Wagreich, Letters 58 (2), 189–201. Ines Wendler, Ismail Yilmaz, among others) for their discussion dur- Chen, W.B., 2008. Element Geochemical Record of Cretaceous Oceanic Red Beds: ing the IGCP workshops and meetings in the past years. We would Implications for Paleoceanic Environment. China University of Geosciences (Beijing), Beijing. (71 pp.). like to thank our Chinese colleagues (Xi Chen, Yongjian Huang, Chen, X., Wang, C., Kuhnt, W., Holbourn, A., Huang, Y., Ma, C., 2011. Lithofacies, Shaoyong Jiang, Guobiao Li, Xianghui Li, Yongxiang Li, Xiaoqiao microfacies and depositional environments of Upper Cretaceous oceanic red beds Wan, and Kuidong Zhao) and students (Wenbin Chen, Zhiyan Han, (Chuangde Formation) in southern Tibet. Sedimentary Geology 235, 100–110. Xiang Li) who were devoted to the study of CORBs. We thank Prof. Clarke, L.J., Jenkyns, H.C., 1999. New oxygen isotope evidence for long-term Cretaceous climatic change in the Southern Hemisphere. Geology 27 (8), 699–702. Frédéric Boulvain, Dr. Mingcai Hou, Dr. Wenzhong Li and Dr. Kuidong Coccioni, R., Premoli Silva, I., 2012. Upper Albian–Maastrichtian calcareous plankton Zhao for their kindness to supply photos for Fig. 25. This study was biostratigraphy and magnetostratigraphy of the Gubbio section (Italy). In: “ financially supported by the Chinese MOST 973 Project Coccioni, R., Menichetti, M. 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