A novel chronostratigraphic framework for the Aptian–Albian paleoclimate events

João M F Ramos; Jairo F Savian; Daniel R Franco; Milene F Figueiredo; Carolina G Leandro; Rodolfo Coccioni; Fabrizio Frontalini; Nicola Casadei; Hironao Matsumoto; Leonardo R Tedeschi; Luigi Jovane; Ricardo I F Trindade

doi:10.1038/s41598-026-35714-z

. 2026 Jan 20;16:5862. doi: 10.1038/s41598-026-35714-z

A novel chronostratigraphic framework for the Aptian–Albian paleoclimate events

João M F Ramos ^1,^2,^✉, Jairo F Savian ^1,³, Daniel R Franco ⁴, Milene F Figueiredo ⁵, Carolina G Leandro ³, Rodolfo Coccioni ⁶, Fabrizio Frontalini ⁷, Nicola Casadei ⁷, Hironao Matsumoto ^8,⁹, Leonardo R Tedeschi ¹⁰, Luigi Jovane ¹¹, Ricardo I F Trindade ¹²

PMCID: PMC12894972 PMID: 41559380

Abstract

The Aptian–Albian interval was marked by paleoclimatic changes encompassing volcanic events and monsoonal activity of which the precise timing were not well-defined. This study presents a ~ 20 Myr cyclostratigraphic framework and provides an extended astronomical timescale using magnetic susceptibility and anhysteretic remanent magnetization datasets from the Poggio le Guaine core (Umbria–Marche Basin, Italy). Using astronomical tuning, we infer new age constraints and timespans for: (1) the Aptian and Albian stages; (2) the magnetic polarity Chron M0r; (3) the Oceanic Anoxic Events; (4) bioevents and, (5) Cretaceous oceanic red beds. These results enhance our understanding of the interplay between paleoclimatic and biotic processes in the Early Cretaceous. This study establishes a new chronostratigraphic framework for the Aptian–Albian—the most comprehensive to date in terms of resolution, integration with other geochronological methods, and temporal coverage— offering new insights into bio-chemical–geomagnetic interactions and helping refine the Geological Time Scale.

Supplementary Information

The online version contains supplementary material available at 10.1038/s41598-026-35714-z.

Introduction

The Aptian–Albian interval (121.4 to 100.5 million years ago) was characterized by significant tectonic and paleoclimatic changes that profoundly influenced global sea levels and oceanic conditions^1–9. Continental breakup during this period drove the formation of new oceanic ridges and accelerated seafloor spreading, leading to shallower, potentially anoxic oceans and elevated global sea levels^10,11.

The Cretaceous experienced the largest long-term eustatic fluctuations of the entire Phanerozoic, largely controlled by the increasing ridge volume associated with the opening of the South Atlantic and Southern Ocean basins^11,12. This tectonic activity contributed to a sustained sea-level rise, primarily driven by changes in ocean basin volume and enhanced seafloor spreading linked to the development of these nascent ocean basins¹³. Superimposed on this long-term trend were pronounced short-term fluctuations, especially evident at the Aptian–Albian boundary, reflecting a complex interplay of processes and a shift in the dominant controls on sea-level change—from tectonic drivers to astronomically forced climate variations^11,14.

Concurrently, the interval witnessed major paleoclimatic shifts linked to volcanism, monsoonal activity, and changes in ocean circulation^1–9. These environmental changes induced rhythmic black shale deposition, abrupt oceanic redox fluctuations, alterations in precipitation and runoff that affected halocline and thermocline stratification, foraminiferal turnovers, and episodic hyperthermal and cooling events within an overall greenhouse climate¹⁵. However, the precise timing and duration of these events remain poorly constrained, fueling debate over their causes and global impacts^16–19. This uncertainty is further compounded by imprecise age models and data gaps in Cretaceous records, which limit detailed understanding of climate dynamics during this critical period.

During the Aptian, the proximity of continents caused unstable circulation in the Tethys Ocean, with frequent oceanic oscillations on a 10² kyr timescale⁹. This led to four oceanic anoxic events—OAEs (1a, 1b, 1c, and 1d) that denote intervals of episodic burial of organic-rich sediments (i.e., black shale layers) under oxygen-deficient bottom waters^7,17,20. The early Aptian OAE 1a, also known as the Selli event, was the most prominent (i.e., ca. 1-Myr-long-lasting event^16,17) event and was associated with Large Igneous Provinces (LIP) volcanism, as revealed by the osmium isotope record^8,21. The OAE 1b was characterized by a ca. 2.84 Myr prolonged period of organic carbon burial¹⁹ and encompassed multiple short-duration sub-events resulting from the combination of a warm climate (driven by volcanic CO₂ emissions), heavy precipitation, and intense weathering, all of which amplify orbital forcing²². The early-late Albian OAE 1c and late Albian OAE 1d were supra-regional marine anoxic events⁸ of limited magnitude, affecting the North Atlantic and proto-Atlantic Ocean.

Despite an overall warming trend, a global cooling event known as the “cold snap” occurred in the late Aptian^18,23. The end of this period is marked by a foraminiferal turnover²⁴ and one of the most pronounced breaks in the maximum magnitudes of short-term sea-level variation¹⁴, suggesting a major shift in the ocean–atmosphere system. These variations indicate a complex interplay of processes and suggest a shift in the dominant controls on global sea-level change—from tectonically driven mechanisms towards those modulated by astronomically forced climate variations¹¹.

We examine the Poggio le Guaine (PLG) core, a key sedimentary record for Cretaceous biostratigraphy²⁰ (Fig. 2). This core from the Umbria-Marche Basin (Tethys Ocean)^7,20 provides a comprehensive record of OAEs, organic-rich beds, and seven Cretaceous Oceanic Red Beds (CORBs) intervals (Fig. 2a). The PLG core records changes in ocean circulation from the period preceding the opening of the South Atlantic to the establishment of full connectivity between the South and North Atlantic via the equatorial seaway²⁵ (Fig. 1). This circulation shift, observed near the record’s top, marks a transition to more stable conditions with a well-defined thermocline and enhanced currents⁹. The bioevents, isotopic data, and Chron M0r²⁶ help develop a refined reference age model^3,8,27,28. The PLG core also records key biohorizons: the Globigerinelloides aptiensis zone at the base of the Aptian, Leupoldina cabri zone (OAE 1a), Microhedbergella miniglobularis zone (foraminiferal turnover), Microhedbergella renilaevis (top of the Aptian), Biticinella breggiensis zone (OAE 1c), and Thalmanninella globotruncanoides zone marking the Albian-Cenomanian boundary^7,20 (Fig. 2a). High-resolution paleoclimatic proxies from the PLG core enable precise positioning of these events within the Geological Time Scale (GTS)¹.

Fig. 1 — Paleogeographic reconstructions at 120 and 100 Ma. Location of the PLG (yellow circle) core in the Tethyan realm (using Gplates 2.3.0).

This study presents a new astronomically tuned age model using magnetic susceptibility (MS) and anhysteretic remanent magnetization (ARM) as proxies for the PLG core^7,20, which spans from the base of the Aptian stage to the base of the Cenomanian^18,20,26, (Fig. 2b,c). The astrochronology age model was anchored with the Chron M0r^26,32 and to the absolute dating of Vöhrum’s tuff^33,34, transferred via isotope correlation from the original location to the PLG core²². The astrochronological study presented here offers an opportunity to develop a high-resolution age model for a ~ 20 Myr-long interval of continuous sedimentary record—from OAE 1a to OAE 1d—using high-resolution magnetic proxies. We also determine the ages of CORBs and biostratigraphic events along this interval. Furthermore, by converting spatial (meters) to temporal domain (million years) using orbitally influenced proxies, we can convert the carbon (δ¹³C) stable isotope profile to the time domain and estimate the durations of volcanic events, as constrained by the ¹⁸⁷Os/¹⁸⁸Os ratios⁸, (Fig. 2d,e). Our results provide an astronomically tuned, high-definition stratigraphic reference record of global climate evolution for the Aptian-Albian interval.

Results and discussion

Integrated stratigraphy and biozones

White cherty limestone, cyclically intercalated with a few centimeter-thick black shale layers of the Maiolica Fm. occurs at the base of the core (Fig. 2a). The most prominent organic-rich interval, the Selli Level, is found between 89.24 and 91.19 m and represents the sedimentary expression of OAE 1a^18,26. Additional black shale layers between 88–88.2 and 79.31–79.37 m correspond to the Wezel and Fallot Levels, respectively, which have exclusively been reported from the Tethyan region¹⁸.

An organic-rich marker bed, the Jacob Level (67.44–67.52 m), marks the onset of OAE 1b (Fig. 2). This OAE is characterized by several prominent organic-rich intervals, including the sub-events designated as the Kilian (62.64–63.02 m), Monte Nerone (57.07–60.01 m), Urbino/Paquier (54.63–54.88 m), and Leenhardt (52.97–53.26 m) Levels²². The top of the Aptian interval, marked by the lowest occurrence (LO) of Microhedbergella renilaevis¹, is set at 63.40 m.

An approximately 3-m-thick interval (24.74–27.77 m), known as the Amadeus Segment^24,28, lies within the Albian interval and represents OAE 1c (Fig. 2). The regional sedimentary expression of OAE 1d³⁵ is at the top of the PLG core. Additionally, within the Scaglia Bianca Fm.³⁶, the Pialli/Breistroffer Level (0.95–4.91 m) is close to the boundary with the Cenomanian stage and exhibits a cyclic character influenced by orbital forcing²⁷.

The succession of primary planktonic foraminiferal bioevents allowed the subdivision of the section into the following zones and subzones (in stratigraphic order): Globigerinelloides aptiensis during the Barremian–Aptian transition, Leupoldina cabri, Globigerinelloides ferreolensis, Globigerinelloides algerianus, Hedbergella trocoidea, Paraticinella rohri, and Microhedbergella miniglobularis in the Albian (Fig. 2). The Aptian–Albian boundary is marked by the FO of Microhedbergella renilaevis¹. The Albian succession encompasses the Microhedbergella rischi, Ticinella primula; Ticinella praeticinensis, Pseudothalmanninella subticinensis (the last two forming the Biticinella breggiensis), Pseudothalmanninella ticinensis Parathalmanninella appenninica and ending with Thalmanninella globotruncanoides, which marks the top of the Albian²⁹.

The calcareous nannofossil zones³⁰ include Chiastozygus litterarius (NC6), Rhagodiscus angustus / Eprolithus floralis (NC7), Paraticinella columnata (subcircular) in the Aptian, Paraticinella columnata (circular, NC8A), Hayesites albiensis (NC8B), Tranolithus orionatus (NC8C); Axopodorhabdus albianus (NC9A), Eiffellithus monechiae (NC9B), and Eiffellithus turriseiffelii (NC10) in the Albian (Fig. 2).

Magnetic records and the Aptian–Albian astrochronology

MS measures the degree to which a sediment or rock is magnetized in the presence of a magnetic field, reflecting the concentration and grain size of magnetic minerals, such as magnetite. Low-field MS (χ, in m³/kg) provides an indirect estimate of the abundance of paramagnetic and ferromagnetic minerals and has been widely applied in cyclostratigraphic studies¹⁸. Variations in MS can serve as proxies for changes in sediment supply, weathering intensity, or bioproductivity. ARM, which is sensitive to particle size, concentration, and mineralogy, similarly captures environmental and climatic variations through time by recording subtle shifts in sedimentation dynamics.

In this study, log₁₀(MS) values range from − 0.48 to 1.71 m³/kg (Fig. 2b), while log₁₀(ARM) values vary from 1 × 10⁻⁹ to 1 × 10⁻⁵ Am²/kg (Fig. 2c). The highest MS and ARM values are observed in CORB 2, a lithofacies rich in clay, marl, and limestone in the Marne a Fucoidi Fm.²⁰, whereas the lowest values are found in the white lithofacies of the Maiolica Fm. Similar to CORB 2, CORBs 3, 4, and 5 exhibit elevated MS and ARM values (Fig. 2b,c).

Significant long- and short-wavelength quasiperiodic components are evident in both the MS and ARM datasets (Fig. 3). The 5π − MTM-Robust-AR(1) and eFFT spectra for both log₁₀(MS) and log₁₀(ARM) data (after detrending)—covering the stratigraphic interval from the top of the Albian (i.e., first occurrence (FO) Th. globotruncanoides²⁰) to the base of the Aptian (i.e., M0r²⁶) stages— distinguished three sets of low-frequency bands, common to both datasets and persistent throughout all studied intervals: (i) 1/1.75–1/2.63 m⁻¹ (central frequency of ~ 1/2.3 m⁻¹); (ii) 1/0.52–1/0.66 m⁻¹; (iii) 1/0.17–1/0.30 m⁻¹); and (iv) 1/0.10–1/0.16 m⁻¹) (Figs. 3, 4 and Supplementary Figs. S1 and S2). The observed wavelength ratios verified for the log₁₀(MS) spectra (0.44,1.27,1.71,4.47,7.81,9.4 = 1.0,2.9,3.9,10.2,17.8,21.4) are compatible to those expected from the Milankovitch spectral peak ratios³⁷ for the top of the Aptian stage (~ 113 Ma): 1/405,1/125,1/95,1/38.7,1/22.9,1/18.4 = 1.0,3.2,4.3,10.5,17.7,22.0.

Fig. 4 — Astronomical calibration of PLG. (a) Detrended log(MS), gray line, and 405-kyr component, red line. E-cycles¹⁷ are numbered in red and correlated using a dashed red line. (b) La2004 orbital solution⁴⁵ (eccentricity only, gray line) and 405-kyr cycle (red) used for tuning. (c) SAR curve based on 405-kyr tuning (green line). (d) Tuned logarithmic scale MS data (blue line) with 405-kyr filter output (red line). (e) Tuned ¹⁸⁷Os/¹⁸⁸Os ratios^3,9,27,31. The ages of the major unradiogenic shifts are highlighted. (f) PLG core magnetic polarity tuned²⁶. Gray rectangles indicate no data. (g) Tuned PLG core CORBs¹² (Table 1). (h) Planktonic foraminifera zones^20,29 in the geological time domain (Table 2). (i) Calcareous nannofossil zones^20,29,30 in the geological time domain (Table 2). M. = *Microhedbergella, min.* = *miniglobularis, P.* = *Paraticinella, H.* = *Hedbergella, G.* = *Globigerinelloides, ferr.* = *ferreolensis, L.* = *Leupoldina, B.* = *Biticinella, T.* = *Ticinella, Psth.* = *Pseudothalmanninella, Pth.* = *Parathalmanninella, Th.* = *Thalmanninella.*

Furthermore, by applying the correlation coefficient (COCO)/evolutionary COCO (eCOCO) methods³⁸, optimal mean sedimentation rates of ~ 0.56 cm/kyr for MS data and ~ 0.61 cm/kyr for ARM data were determined, with more than six astronomical parameters each, and the null hypothesis (= no orbital forcing) significance levels were less than 0.0005 (Fig. 3e,l). These are consistent with mean sediment accumulation rates (SARs) of 0.561 cm/kyr (log₁₀(MS)) and 0.571 cm/kyr (log₁₀(ARM)) at null hypothesis significance levels < 0.001 obtained by application of the average spectral misfit (ASM) method^39,40 (Fig. 3g,o and Supplementary Figs. S1 and S2). Such findings allow us to infer that that MS and ARM datasets record orbitally-paced variations at (i) 405-kyr long eccentricity; (ii) 95–125-kyr short eccentricity; (iii) obliquity; and (iv) precession.

The phase analysis of the PLG core was previously conducted by Ramos et al.²², which enables the estimation of the amplitude and phase of astronomical cycles during model fitting. We utilized their results prior to performing the astronomical tuning (see Ramos et al., 2024 for details). The phase of the 405-kyr sinusoidal component derived from the MS dataset is in phase with the long eccentricity signal (peak-to-peak, cave-to-cave). It is important to note that a phase offset between the classical La04 solution⁴¹ and more recent orbital models^42–44 can introduce an age discrepancy of up to 185 kyr—specifically between La04 and ZB20a—when applied in astronomical tuning (see Supplementary Figures S4 and S5).

We then construct a ~ 405-kyr tuned age model for the MS data based on the g₂–g₅ target curve from the La2004 astronomical solution³⁸ and the 405-kyr long-eccentricity sinusoidal curve from a Gaussian bandpass filtering (0.02–0.35 cycles/m) applied to the log₁₀(MS) dataset, assuming as “tiepoints” the age of ~ 113.27 Ma (related to the Kilian Level^33,34) and the M0r base (120.29 ± 0.09 Ma³²) (Fig. 4a–d). The number of internal cycles within the Selli Level (OAE 1a) was estimated using the ages at the base (91.19 m; 119.534 Ma) and top (89.24 m; 118.273 Ma) of this event as proposed by Li et al.¹⁶. The anchored astrochronology, which covers the entire ~ 96.02 m PLG core, identifies 47 long-eccentricity cycles from the top of the Barremian (base of M0r Chron) through the top of the Albian stage (E17 to E33), implying a timespan of ~ 19.96 Myr (100.53–120.49 Ma). This new age model provides new ages for important geological events throughout the Aptian–Albian interval (Tables 1, 2): 1) ~ 6.97 Myr for the Aptian stage (GTS2020 = 8.2 Myr¹) (from the base of Chron M0r and the LO of M. renilaevis), 2) ~ 12.82 Myr for the Albian (GTS2020 = 12.7 Myr) (from the LO of M. renilaevis to the LO of Th. Globotruncanoides), 3) ~ 430 kyr for the magnetic polarity Chron M0r; 4) 120.32 ± 0.203 Ma Ma for the Barremian–Aptian boundary and 5) the durations of the principal black shales and marker levels (e.g., the Amadeus Segment).

Table 1.

Estimated timespans of OAEs, M0r Chron, organic-rich marker beds and CORBs.

	Depth (m)	Age (Ma)	Timespan (kyr)
Events and Chron
OAE 1d	0.95–4.91	100.66–101.59	900
OAE 1c	14.50–33.57	103.66–107.56	3900
OAE 1b	52.09–67.52	111.40–114.10	2700
OAE 1a	89.24–91.19	118.37–119.50	1130
M0r Chron	92.70–95.10	119.89–120.32	430
Organic-rich marker
Pialli Level	0.95–4.91	100.66–101.59	930
Amadeus Segment	22.77–24.77	105.29–105.66	370
Leenhardt Level	52.97–53.26	111.40–111.44	40
Urbino Level	54.63–54.88	111.67–111.71	40
Monte Nerone Level	57.07–60.01	112.17–112.77	600
Kilian Level	62.64–63.02	113.21–113.27	60
113/Jacob Level	67.44–67.52	114.08–114.10	20
Fallot Level	79.31–79.37	116.63–116.64	10
Wezel Level	88.00–88.20	118.17–118.21	40
Selli Level	89.24–91.19	118.37–119.50	1130
CORB #
CORB 7	7.70–13.55	102.13–103.42	1290
CORB 6	27.30–30.43	106.18–106.84	660
CORB 5	34.40–35.89	107.72–107.99	270
CORB 4	49.18–50.46	110.56–110.77	210
CORB 3	57.07–60.01	112.17–112.77	600
CORB 2	63.02–67.28	113.28–114.05	770
CORB 1	68.31–82.24	114.25–117.11	2860

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Depths for the upper and lower boundaries of the OAEs, organic-rich marker beds and CORBs identified in the PLG core^7,8,17 and the ages and timespans estimated in this work. All the ages presented in the table have a maximum age uncertainty of ± 0.203 kyr.

Table 2.

Estimated timespans of biostratigraphic zones.

	Depth (m)	Age (Ma)*	Timespan (kyr)
Planktonic foraminiferal zones
Th. globotruncanoides**	0.00	100.63	-
Pth. appenninica	10.00	102.70	2130
Psth. ticinensis	16.82	104.16	1460
Psth. subticinensis	19.35	104.64	500
T. praeticinensis	37.55	108.27	3630
T. primula	50.03	110.70	2430
M. rischi	60.20	112.80	2100
M. renilaevis***	63.40	113.35	550
M. miniglobularis	64.50	113.56	210
P. rohri	71.47	114.91	1350
H. trocoidea	75.51	115.64	730
G. algerianus	85.76	117.77	2130
G. ferreolensis	87.44	118.08	310
L. cabri	91.35	119.60	1520
Calcareous nannofossil zones
NC10A (E. turriseiffelii)	15.09	103.80	3270
NC9B (E. monechiae)	18.05	104.40	600
NC9A (A. albianus)	39.37	108.63	4230
NC8C (T. orionatus)	41.37	109.05	420
NC8B (H. albiensis)	55.79	111.88	2830
NC8A (P. columnata (cir.))	65.78	113.78	1900
NC8A (P. columnata (sub.))	66.34	113.88	100
NC7 (E. floralis)	88.67	118.28	4400

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Depths for the upper boundaries of the planktonic foraminiferal and the calcareous nannofossil zones identified in the PLG core^7,29,30 and this work, and the ages and timespans estimated in this work. All the ages presented in the table have a maximum age uncertainty of ± 0.203 kyr.

⁽*⁾ base of the biostratigraphic zone,⁽**⁾ top of Albian,⁽***⁾ top of Aptian.

Astronomically dating Chron M0r, OAEs, organic-rich marker beds and CORBs

The chaotic behavior of the solar system^46,47 places a fundamental limit on the applicability of astronomical solutions for deep-time astrochronology. Even the most advanced models, such as ZB20x, are only reliable back to ~ 66 Ma⁴⁴, well short of the Aptian–Albian interval examined here. Consequently, the most robust anchor for early Cretaceous time calibration remains the stable 405-kyr eccentricity cycle, whose g₂–g₅ frequency term exhibits exceptional long-term stability⁴⁶. We therefore employ the 405-kyr component from the La2004 solution, ensuring both methodological continuity with previous studies of the PLG core and practical equivalence with more recent orbital solutions.

Because no orbital solution is validated for the 100–120 Ma interval, age estimates must incorporate uncertainties arising from phase offsets among available models. Comparisons of the 405-kyr components from La2004⁴¹, La2010a,b,c,d, and ZB20x show phase differences of up to ~ 185 kyr for 118–121 Ma, and the theoretical maximum offset between two fully anti-phased 405-kyr solutions is 202.5 kyr. We therefore assign a conservative uncertainty of ± 202.5 kyr to all astrochronologically derived ages. This uncertainty framework is supported by four high-precision absolute tie points for the base Aptian, in addition to well-established ages for the Kilian level and the top of the Albian, providing a robust structure for anchored astrochronology of the PLG core.

Although the filtered 405-kyr La2004 curve and the MS record exhibit an excellent phase match—differing by only ~ 100 kyr—such agreement may still be coincidental. The incorporation of a ± 202.5 kyr uncertainty therefore provides an appropriately cautious and defensible representation of age within the constraints of current astronomical solutions.

The presence of the M0r Chron in the PLG core reinforces its suitability as a candidate reference section for the Aptian and Albian stages^7,18,20,32. This reverse chron is a proposed marker for the Barremian–Aptian boundary^1,48,49, and its age and duration are crucial factors for constraining past oceanic, tectonic, and geodynamic behavior³². Multiproxy orbital tuning for the PLG core provides an age of 120.32 ± 0.203 Ma for the base of Chron M0r (Barremian–Aptian boundary, Fig. 4, Table 1). This age is similar to the previously established ages (~ 120.2 Ma—by astrochronology¹⁸; 120.29 ± 0.09 Ma–by integration of U–Pb and ⁴⁰Ar/³⁹Ar ages^16,32. For the top of Chron M0r, we obtain an age of 119.89 ± 0.203 Ma , ~ 200 kyr older than the central age, 119.70 ± 0.12 Ma³². Regardless of the astronomical solution used (La2004⁴¹; La2010x^42,43; ZB20x⁴⁴) (see Supplementary Figures S4 and S5), the astrochronology presented here indicates a timespan of approximately 430 kyr for the M0r Chron (based on the La2004 solution⁴¹), which is shorter than the 540 ± 37 kyr duration proposed by Li et al.³².

Previous age estimates of the base of the Selli Level (OAE 1a) range from 124.55 to 119.30 Ma^18,50,51. Our astrochronology indicates an age for the base of Selli Level of 119.50 ± 0.203 Ma and a timespan of 1.13 Myr. This value nearly coincides with the timespan proposed of 1.11 ± 0.11 Myr⁵¹ (Table 1), and the U–Pb zircon ages of the tuffs constrain OAE 1a, indicating that it occurred between 118.273 ± 0.203 and 119.534 ± 0.203 Ma¹⁶. The correlation between the Selli event in the PLG core and the Cismon core using carbon (δ¹³C) stable isotope profiles and calcareous nannofossil zones^15,52 indicates that the brief hyperthermal event described at the onset of OAE 1a has an approximate duration of 90 kyr, suggesting that eccentricity orbital forcing may have acted as an amplifier of the climate-oceanographic disturbance caused by the 1.13 Myr volcanic long-term event that produced OAE 1a (Fig. 5 and Supplementary Fig. S3).

Fig. 5 — New Aptian–Albian GTS. Modified Aptian–Albian timescale, including ages of the planktonic foraminifera zones and calcareous nannofossil zones. CORB and OAE ages are also presented. Organic-rich beds markers are indicated by black rectangles. (h) anchored astrochronology generated in this study and (i) corresponding E-cycles¹⁷ in red. (j) arranged in the geological time domain. Extrapolated NI and TI indexes from Tethyan realm^15,36.

Despite the widespread acceptance of OEA 1b as a late Aptian–early Albian event^1,7, some controversies exist regarding OEA 1b divisions (or sub-events). Herrle et al.⁵³ have constrained OAE 1b in the Aptian–Albian boundary, an interval considerably restricted when compared to the definition by Coccioni et al.⁷ and the GTS2020¹. In this study, we set the base of the OAE 1b to the Jacob Level, and its top with the Leenhardt Level⁸ where magnetic minerals have recorded significant environmental change. The base of OAE 1b has been estimated at 114.10 ± 0.203 Ma and the top at 111.40 ± 0.203 Ma, with an inferred timespan of 2.7 Myr. Our study reveals that the sub-events within OAE 1b, that are, the Jacob, Kilian, Monte Nerone, Urbino, and Leenhardt Levels, have central ages of 114.09, 113.25, 112.49, 111.69, and 111.42 ± 0.203 Ma, with durations of ~ 20, ~ 70, ~ 560, ~ 50, and ~ 380 kyr, respectively (Table 1).

Initially postulated as an anoxic event at the boundary between the Albian and Cenomanian (corresponding to OAE 1d²¹), OAE 1c is now defined as several layers moderately enriched in organic carbon in the Biticinella breggiensis planktonic foraminiferal zone⁴⁵. OAE 1c has been interpreted as a detrital oceanic anoxic event characterized by high input of derived organic matter, resulting in the deposition of black shales. The Amadeus Segment^31,54, part of OAE 1c, though encompassing a long-term trend in NI¹⁵, has a low nannofossil fertility index and high temperatures³⁶ (Fig. 5). Periodic anoxic conditions during the Amadeus segment were associated with warmer and less fertile surface waters, where variations in temperature or salinity (i.e., reduction of surface water density) are expected to be linked to Milankovitch cycles³⁶. According to the results of this study, OAE 1c lasted 3.9 Myr (103.66–107.56 ± 0.203 Ma), within which the Amadeus segment is constrained at 105.66–105.29 ± 0.203 Ma.

Paleoenvironmental inferences across OAE 1d point to a warm climate (m-TI) and low fertility followed by a phase of lower temperature and higher fertility conditions within an overall warm climate that started prior to OAE 1d¹⁵ (Fig. 5). Other UMB records exhibit a decrease in the NI prior to OAE 1d, followed by an increasing trend initiating in the middle of the Pialli Level. Except for a ~ 3.2 Myr interval, the Aptian is a period of generally low fertility (oligotrophy) compared to the Albian¹⁵. Between OAE 1b and OAE 1d, there is a long phase (~ 8 Myr) of completely mesotrophic conditions, segmented by short periods of low fertility conditions. The maximum NI values were reached at the end of OAE 1d. There is no apparent relationship between the OAEs and fertility, as seen in the lower/middle Aptian, which exhibits long-term moderate nutrient conditions without any association with OAE (Fig. 5). Although OAE 1b is generally oligotrophic, black shale levels show high fertility¹⁵.

The Pialli/Breistroffer Level in the UMB correlates with the lowermost part of the positive carbon isotopic excursion characterizing the late Albian–early Cenomanian OAE 1d. This interval is marked by frequent, well-marked, organic-rich layers and is called the Niveau Breistroffer⁴⁵. For ~ 0.9 Myr (i.e., OAE 1d), anoxic conditions dominated in the Tethys Ocean. The black shale–poor facies deposited in the late Albian result from a shift in oceanic circulation from unstable to more stable conditions associated with a well-established thermocline and better-developed surface and bottom currents across interconnected basins³. Our study shows that, since the end of the deposition of the black shales related to OAE 1c (around ~ 103.7 Ma), the late Albian experienced a period of about 2.1 million years without significant deposition of black shale until the onset of OAE 1d at around ~ 101.6 Ma (Fig. 5).

CORBs and temperature during Aptian–Albian stages

CORBs are reddish to pinkish to brownish sedimentary rocks deposited in pelagic marine environments, comprising levels enriched in hematite and goethite^24,55,56. In the PLG core, CORB deposition began at approximately 117.1 ± 0.203 Myr (or 82.24 m), 1.3 Myr after the occurrence of the organic beds of OAE 1a, where approximately 14 m of reddish lithologies were deposited. During this period, a significant paleoenvironmental change occurred, as shown by a break in ARM data (proxy for fine-grained magnetic minerals) and a sudden change in sediment color. The oxidizing environment that resulted in CORB 1²⁰ lasted approximately 2.86 Myr (Fig. 5, Table 1). A marked change in the input of magnetic minerals matches well with the onset of OAE 1b, where both MS and ARM indicate the highest input of fine magnetic material. This input corresponds to an interval without any fertility anomaly and is not associated with any abrupt temperature change²⁷. Unlike the interval with high MS and ARM values present during the Monte Nerone Level, CORB 2 is unrelated to any significant δ¹³C_carb shift.

Within OAE 1b, two oxidant events (CORBs 2 and 3) are notable for their high MS and ARM values. Unlike CORB 1, CORBs 2 and 3 have short durations (0.77 and 0.60 Myr, respectively). The modulation of the 405-kyr component in these intervals suggests that the orbital forcing acts as an amplifier, even though it is not the main cause of the oxidant character of the oceans. Detailed mineralogical studies have indicated that hematite, goethite, and Mn²⁺-bearing calcite cause the red coloration of the CORBs in ODP 1049^55,56. In the PLG core, the amount of hematite and goethite present in the red beds is likely controlled by short eccentricity and precession cycles, which are modulated by long eccentricity. Long-term cycles (~ 1.2 and ~ 2.4 Myr) are described in Cretaceous geological records^46,47,57,58. Based on this timespan, we suggest that the large modulation cycle of obliquity, ~ 1 Myr (or a resonant state of the ~ 2 Myr eccentricity cycle^47,57), may control the differential input of fine magnetic minerals⁵⁷.

After the end of OAE 1c and before the major paleogeographic and palaeoceanographic changes at the end of the Albian, a new predominant oxidant period lasting 1.29 Myr begins (Table 1). Similar to CORB 6, CORB 7 is locally represented by predominantly orange-colored sediments and does not exhibit high values of MS and ARM. This interval preceding OAE 1d is characterized by a mesotrophic nutrient level and marks the period with the highest temperatures throughout the Albian. The end of CORB 7 represents a significant environmental change glimpsed through (i) the onset of the positive excursion of δ¹³C_carb and, consequently, the beginning of OAE 1d; (ii) the return to oligotrophic conditions; (iii) a new warming period, recorded by the positive excursion of δ¹⁸O_carb and m-TI SSTs (°C); and (iv) the deposition of the rare black shales of the Pialli Level^9,15,29.

Through the analysis of carbon isotope curves and calcareous nannofossil zones, we have correlated the measurements of the NI and TI from various sections and classic cores within the Tethyan context (i.e., Cismon core, Piobbico, and Mt. Petrano)^15,36 with the PLG core, allowing for precise positioning within the geological time scale (Fig. 5 and Supplementary Fig. S3). In general, the m-TI data indicate that the Albian was warmer than the Aptian^15,36. The m-TI suggests that temperatures remained consistently high until the end of OAE 1a, enduring for approximately 2.1 million years, with a rapid warming event occurring at the onset of OAE 1a. After OAE 1a, the m-TI data show a cooling trend that persisted through the late Aptian (~ 5 million years). Notably, no clear correlation emerges between temperatures and the OAEs. Despite a pronounced temperature drop immediately following OAE 1a, gradual warming occurred during OAE 1b, with thermal maxima associated with black shale deposition. A long-term pattern of elevated temperatures extends beyond OAE 1c, while temperature reductions are observed within OAE 1d.

In the Tethyan Ocean, the lower portion of the H. trocoidea planktonic foraminiferal zone marks the cold snap^18,23, a mid-Cretaceous global cooling event characterized by the breakdown of warm, equable climate conditions^23,58,59. Our study, based on correlations between the PLG core and Tethyan sections^15,18,23,36 (Supplementary Fig. S3), suggests that mild temperatures prevailed during most of the Aptian (~ 118.4 to 113.4 Ma). This “cold” period, characterized by the lowest surface water temperatures (as indicated by TI¹⁵), corresponds with elevated δ¹³C values, suggesting changes in oceanic circulation, carbon sequestration at the surface, and potential incorporation of carbon into marine organisms. Despite the lowest temperatures occurring within the cold snap interval^16,18, the interval may be extended to between 113.4 and 118.4 Ma, based on temperature data¹⁵ and the inferred paleoclimatic stability that allowed for the continuous deposition of CORB 1 over this time span (Fig. 5). Evidence of a milder climate spans from the NC6 nannofossil zone to the lower/middle NC8 zone^15,30 corresponding to the deposition of CORBs 1 and 2²⁰, with the coldest temperatures (~ 115.5 to 113.4 Ma) confined to the upper Aptian. The paleoenvironmental shift likely began earlier (G. algerianus zone, ~ 118.4 Ma), marked by sharp changes in δ¹⁸O values and TI¹⁵, as reflected by a sudden shift in sediment color^18,20.

The most pronounced break in the maximum magnitudes of short-term sea-level variation (at the end of Trend II¹⁴) is associated with the termination of the Aptian cold snap. During this transition, the dominant high-frequency, high-amplitude, orbitally forced glacio-eustasy that characterized the late Aptian was replaced by aquifer-eustasy. Consequently, the elevated magnitude values observed during the Aptian gave way to significantly lower values in the Albian¹⁴. This shift coincides with the final rupture of Gondwana and the establishment of full marine connectivity between the South and North Atlantic Oceans¹³ suggesting that tectonic reorganization played a key role in altering eustatic drivers.

Over intermediate timescales of 2–10 Myr, dynamic topography exerts a significant influence on sea-level change and cannot be disregarded. Notably, the late Aptian increase in Fe and Ti concentrations—proxies for enhanced weathering and aquifer discharge—indicates that during the Albian, aquifer-eustasy modest eustatic fluctuations may have overtaken astronomically driven climate mechanisms as the primary control on global sea-level variations²².

Aptian–Albian bioevents

The bioevents of the PLG core provide a detailed biostratigraphic scheme based on planktonic foraminifera and calcareous nannofossils^7,20,29,30. Our study also provides new ages for planktonic foraminiferal and calcareous nannofossil bioevents (Fig. 4 and Table 2). Through this study, using as reference the PLG core osmium peaks and carbon and oxygen isotope excursions (all of them with supra-regional to global character), it is possible to refine the biozonations and assess the degree of diachroneity⁵⁹ and the patterns of species dispersal between the Tethyan Ocean and other oceanic regions.

The ages and durations of the bioevents in the PLG core (Table 2) indicate that the top of the Aptian and Albian have ages of 113.35 ± 0.203 (LO of M. renilaevis) and 100.63 ± 0.203 Ma (LO of Th. globotruncanoides), respectively. These values are very close to the ages proposed in the GTS2020¹. However, our astrochronology enables refinement of the ages within these stages, and there are large discrepancies between the ages obtained in our work and those in the GTS2020. For example, the LO of Ticinella praeticinensis occurs 0.87 Myr earlier in our astrochronology than the GTS2020 (108.27 ± 0.203 Ma in the PLG and 107.4 Ma in the GTS2020). The top of this biozone occurs at 104.64 ± 0.203 Ma in the PLG and 107.0 Ma in the GTS2020, resulting in a difference of 2.76 Myr. In other words, the T. praeticinensis planktonic foraminiferal zone lasts 0.4 Myr in the GTS2020 age model and 3.63 Myr in our age model.

The Leupoldina cabri planktonic foraminiferal zone, correlated in the PLG core with the unradiogenic shifts associated with the Selli and Wezel Levels, has its base at 119.60 ± 0.203 Ma and spans 1.52 Myr. The Globigerinelloides algerianus planktonic foraminiferal zone (117.77–115.64 ± 0.203 Ma), associated with the ISEA reverse chron (M"-1r")⁶⁰, has an important ¹⁸⁷Os/¹⁸⁸Os_i peak in its intermediate portion. These two events could potentially serve as tiepoints for correlations between basins. The short time interval of the Microhedbergella miniglobularis zone (timespan of 0.21 Myr) and the base of the M. renilaevis zone (113.35 ± 0.203 Ma) is well-anchored by two ¹⁸⁷Os/¹⁸⁸Os_i peaks, serving as an excellent marker for the top of the Aptian². The sub-events of the OAE 1b Monte Nerone and Urbino/Paquier Levels are included in the Microhedbergella rischi zone (110.70–112.80 ± 0.203 Ma)^7,20.

Assuming that the latter has a global character, the diachroneity between the species of this zone can be evaluated in cyclostratigraphic studies using this black shale level as a tiepoint. Despite the long-lasting period related to the T. praeticinensis (104.64–108.27 ± 0.203 Ma) and B. breggiensis (104.16–108.27 ± 0.203 Ma) zones, significant internal markers within these zones can be used as temporal anchors. In addition to CORB 5 (107.56–107.99 ± 0.203 Ma, basal portion of these zones) and CORB 6 (106.18–106.84 ± 0.203 Ma), the peak of ¹⁸⁷Os/¹⁸⁸Os_i (107.11 Ma) between these two CORBs can serve as an excellent temporal marker to evaluate both the diachroneity of species and issues involving the local vs. global nature of the CORBs^24,55,56. Furthermore, the second ¹⁸⁷Os/¹⁸⁸Os_i peak (104.17 Ma) is practically coincident with the LO of B. breggiensis.

The boundary between the NC6 and NC7 Eprolithus floralis calcareous nannofossil zones³⁰ occurs at 118.28 Ma, between the Selli and Wezel events. This boundary is also marked by high ¹⁸⁷Os/¹⁸⁸Os_i. The top of the NC7 zone (113.88 ± 0.203 Ma, Table 2), which marks the base of the NC8A zone (P. columnata³⁰), is located at the PLG core near the base of OAE 1b²². High values of MS and ARM (reflecting modulation in the anchored astrochronology obtained) coincide with this boundary, highlighting it as an important marker for future correlations between basins. The boundary between the NC8A and NC8B (H. albiensis³⁰) zones, still within the context of OAE 1b, is stratigraphically placed just below the Urbino/Paquier Level^7,20,22 in an interval characterized by the absence of unradiogenic Os shifts^8,61. The NC8C zone (T. orionatus, 108.63 to 109.05 Ma) is set in the upper part of the T. primula zone, near the boundary with T. praeticinensis. In the lower part of the NC9A zone (A. albianus³⁰, 104.40 to 108.63 ± 0.203 Ma), a significant unradiogenic ¹⁸⁷Os/¹⁸⁸Os_i shift occurs at 107.11 Ma. The NC9B zone (E. monechiae³⁰) can be associated with OAE 1c. Within the context of OAE 1c, a second unradiogenic Os shift (104.17 Ma, Fig. 4e) marks the transition between the NC9A and NC9B zones. Starting at 103.80 ± 0.203 Ma (Table 2), the NC10 zone (E. turriseiffelii³⁰) begins and is characterized by the absence of unradiogenic ¹⁸⁷Os/¹⁸⁸Os_i shifts and the presence of OAE 1d in its uppermost portion.

Ages, timespans, and consequences of unradiogenic shifts

The most significant unradiogenic shift begins at 119.48 ± 0.203 Ma and spans the entire OAE 1a interval (Fig. 2). A second shift, lasting 0.34 Myr (118.31–118.97 ± 0.203 Ma), is longer than the Wezel event (0.04 Myr). The third unradiogenic shift, associated with the Fallot Level, occurs between 116.89 and 116.42 ± 0.203 Ma. Similar to the second shift, this shift lasts longer than the deposition of the black shale (0.47 Ma vs. 0.01 Ma; Table 1).

OAE 1b, a mix of volcanic and monsoon-driven OAEs⁸, was triggered by atmospheric circulation reorganization⁴⁵. This event is linked to Hg anomalies¹⁹, suggesting volcanism during this period. The first shift of OAE 1b (113.56 ± 0.203 Ma) coincides with a major planktonic foraminiferal turnover^7,56,62, lasting ~ 0.02 Myr and showing a peak in Hg/TOC, reinforcing the link between volcanic activity and ocean acidification³. Additionally, the Kilian, Monte Nerone (central interval), and Leenhardt subevents are tied to unradiogenic shifts and Hg anomalies¹⁹.

OAE 1c features two unradiogenic shifts. The first, beginning at 107.87 Ma with a peak at 107.11 Ma and pre-shift values at 105.47 Ma, is linked to the onset of the anoxic event. The Amadeus Segment lacks Os data, preventing determination of any unradiogenic shifts. A second shift starts at 104.32 Ma, peaking at 104.17 Ma (Pseudothalmanninella subticinensis planktonic foraminiferal zone). Though within OAE 1c, no black shale is associated with this shift. The monsoon-driven nature of OAE 1d⁸ is evident at the top of the studied interval, where the anoxic event lacks any Os shift.

The 405-kyr long-eccentricity cycles derived from MS and ARM data (0.44 cycles/cm or 2.3 m) provide high-resolution cyclostratigraphy. Using astronomical tuning of the long-eccentricity low-pass filter from the La2004 orbital solution⁴¹, we: (i) propose a anchored astrochronology based on 49 long-eccentricity cycles, enabling a ~ 20 Myr age model from the top of the Barremian to the top of the Albian; (ii) refine the timing and duration of OAEs, black shales, CORBs, and Os unradiogenic shifts; and (iii) estimate bioevent durations using our age model for the PLG core. This study establishes a new chronostratigraphic framework for the Aptian–Albian, enhancing understanding of the biological, chemical, and geomagnetic events and updating the current GTS.

Age model for Aptian–Albian events

The Aptian-Albian interval is characterized by abrupt paleoclimatic shifts, driven by extensive volcanic activity^3,8,17,63 and orbital forcings that intensified the monsoonal system^22,31. Some OAEs and black shales are linked to peaks in volcanism, while others display mixed features, including OAEs and organic-rich layers associated with monsoonal activity^8,22. Our study places key events within the geological timescale, offering a clearer cause-and-effect relationship (Fig. 6).

Fig. 6 — Aptian–Albian events. Aptian–Albian events, including sedimentological indications of the Aptian icehouse interlude, biotic crisis, volcanic episodes and monsoonal activity, responsible for the implantation of the OAEs and the deposition of the black shale levels.

The longest volcanic period (1.13 Myr) in the Aptian-Albian begins with a brief thermal maximum19. The nannoconid crisis¹⁸, occurring ~ 100 kyr before OAE 1a, is likely tied to volcanism starting at 119.50 ± 0.203 Ma. The delay between these events may reflect Os residence time (~ 50 kyr¹⁸) and the environmental response to early volcanic activity. This volcanic phase persists until ~ 118 Ma, when the Aptian icehouse interlude^{15,18,23,53,57}. A subsequent volcanism peak (116.89 to 116.42 Ma) leads to the deposition of Fallot black shale. This mild climate period, with possible ice at high latitudes, shows a long-term carbon curve shift. Two hyperthermals occur at the end of the period: the first, linked to Jacob/113 black shales with no volcanism; the second, related to a unradiogenic ¹⁸⁷Os/¹⁸⁸Osi shift and volcanism^8,22. This volcanism peak marks a significant carbon isotope shift, ends glendonites and dropstones, and concludes the Aptian icehouse interlude.

Foraminiferal turnover⁶² suggests the greenhouse shift was a major change in the ocean–atmosphere system. In OAE 1b, a new hyperthermal causes a significant carbon curve shift and deposition of the Urbino Level. The lack of a unradiogenic ¹⁸⁷Os/¹⁸⁸Os_i anomaly and paleoclimatic proxies link this shale to a monsoonal event²². OAE 1b ends with a volcanic event leading to the Leenhardt Level (Fig. 6). Sporadic volcanism, driven by orbital forcings and high fertility, deposits black shales during OAE 1c. The Amadeus Segment, the main level, lacks a unradiogenic ¹⁸⁷Os/¹⁸⁸Osi shift and may relate to orbital monsoonal activity³¹. OAE 1d, characterized by no volcanic activity, is marked as an oligotrophic period¹⁵.

The correlation between unradiogenic osmium shifts, volcanism, and climate perturbations strengthens the link between astronomical forcing, paleogeographic evolution, and oceanic redox fluctuations during the Aptian-Albian. The updated chronostratigraphic model refines the existing GTS, providing a foundation for inter-basin correlations and enhancing our understanding of the complex interactions between tectonic, climatic, and biotic processes in the Early Cretaceous.

Materials and methods

Geological setting and sampling strategy

The PLG core (lat. 43°32′42.72"N; long. 12°32′40.92"E) was recovered at Poggio le Guaine (northern Apennines, central Italy), a site known to contain one of the most continuous, comprehensive, and well-preserved Aptian–Albian sedimentary sequence. The PLG drill hole cored the uppermost Barremian–lowermost Cenomanian succession of the Umbria-Marche Basin, deposited in the southern margin of the central-western Tethys Ocean^7,20. These pelagic sediments formed following the lithification of the nannofossil-planktonic foraminiferal ooze deposited well above the calcite compensation depth at middle to lower bathyal depths (1000–1500 m) and ~ 20°N paleolatitude^7,18,20 and provides an uninterrupted archive of fossil-rich pelagic strata from 100.5 to 121.4 Ma.

The uppermost Barremian–lowermost Cenomanian succession investigated here is represented by the thin-medium white to gray limestones interbedded with the black shales of the Maiolica Formation and by the overlying distinctive varicolored interlude with more shale of the lower part of the Marne a Fucoidi Formation. The latter consists of thinly interbedded pale reddish to dark reddish, pale olive to dark reddish-brown, and pale olive to grayish olive marlstones and calcareous marlstones together with dark gray to black carbon-rich shales, commonly with a low carbonate content, and yellowish-gray to light gray more or less argillaceous limestones.

Some distinctive organic-rich black shale and calcareous mudstone marker beds occur within the Aptian interval, some of which have been correlated with black shale horizons elsewhere with varying degrees of success and identified as the regional sedimentary expression from the OAE 1a and OAE 1d^7,8,18,20,26. From bottom to top, they are: 1) the Selli Level, known as the organic-rich expression of OAE 1a; 2) the Wezel Level horizon; 3) the Fallot Level; 4) the 113/Jacob Level; 5) the Kilian Level; 6) the Monte Nerone black shales cluster Level; 7) the Urbino/Paquier Level; 8) the Leenhardt Level; 9) the Amadeus segment black shale cluster Level and; 10) the Pialli/Breistroffer black shale cluster Level.

Discrete ~ 8 cm³ cubic samples were then retrieved from the center of the split working halves for paleomagnetic analysis. A total of 3,092 cubic samples were collected along the studied portion of the PLG core (from 96.02 to 0.36 m; average sampling resolution of ~ 3 cm). A total of 355 paleomagnetic cubic samples were also used to measure the stable isotopes (δ¹³C) with a ~ 10-cm resolution.

Rock magnetism

Rock magnetic parameters have largely been used to identify orbital cycles in sedimentary strata⁵⁴ because they are fast, low-cost, and non-destructive. In addition, they allow analysing large sample populations. Rock magnetic measurements indicate that magnetite serves as the primary magnetic carrier²⁶. Low-field MS (χ, in m³/kg) is an indirect measurement of the concentration of paramagnetic and ferromagnetic minerals in geological samples. It has been extensively used as a proxy for terrestrial detrital input in mixed carbonate–clay successions⁴⁶. Nevertheless, variations in the MS occur due to several other factors, such as magnetic mineral concentration, composition, grain size, and shape⁵⁷.

The ARM has been considered in cyclostratigraphy investigations as an alternative proxy to MS⁵⁷. Notably, the ARM obtained at 100 mT (hereafter referred to as ARM) is useful in providing information on fine‐grained (< 20 μm) low-coercivity ferromagnetic minerals⁵⁷. Our rock magnetic cyclostratigraphy analyses comprise both MS and ARM datasets to better assess and compare the spectral content and depositional time range. The MS measurements were performed at the Laboratório de Paleomagnetismo of Universidade de São Paulo (USPMag). They were made on an MFK1-FA Multi-Function Kappabridge at an operating frequency of 976 Hz in a field of 200 A/m. The acquisition of the ARM dataset was performed by the following protocol: remanence measurements were all made in a SQUID magnetometer, model 755 (2G-Enterprises), housed in a magnetically shielded room with internal field < 500 nT at USPMag. The samples underwent a stepwise alternating field (AF) demagnetization over 17 steps: 0–4 mT (step = 2 mT), 4–10 mT (step = 3 mT), 10–40 mT (step = 5 mT), 40–100 mT (step = 10 mT). After AF demagnetization, the sample was submitted to a stepwise ARM acquisition along the same AF demagnetization steps until 100 mT with a direct current bias field of 0.05 mT.

Planktonic foraminifera and calcareous nannofossils

The sample set consists of bulk rocks. At least 30 g of rock were processed for each sample using different methodologies according to the lithology and hardness of the sediment. Samples from softer lithologies were soaked in hydrogen peroxide and desogen. Where required, samples were additionally treated with the surfactant benzalkonium chloride. Samples from hard lithologies were mechanically disaggregated into small fragments (3–8 mm) and treated using the cold acetolysis technique by sieving through a 40-μm mesh and drying at 50 °C. The cold acetolysis method enabled the extraction of easily identifiable foraminifera, even from indurated limestones. This technique offered the possibility of accurate taxonomic determination and detailed analyses of planktonic foraminiferal assemblages, allowing more precise placement of primary and secondary bioevents and zonal boundaries.

Planktonic foraminifera from the washed residues were studied under a stereomicroscope to characterize assemblages and identify biostratigraphic marker species. Taxonomic concepts for genera and species of reference studies were followed (refs.^7,53,62). Calcareous nannofossil assemblages were semi-quantitatively investigated using a Zeiss AxioCam imaging polarizing light microscope at 1250 × magnification. The sediment was processed to obtain homogeneous smear slides following the standard preparation technique: a small rock chip was powdered in a mortar with distilled water buffered with ammonium. A few drops of the solution were smeared on a cover glass, dried on a hot plate, and mounted on a glass slide with two drops of Norland optical adhesive. A total of 45 smear slides were analyzed; in each slide, nannofossil species abundances were logged as follows: A = abundant, > 1 specimen/field of view. C = Common, 1 specimen in 50/fields of view. F = Few, 1 specimen in 100/fields of view. R = Rare, 1 specimen in 200/fields of view. For this analysis, 300 fields of view were scanned on each slide. Calcareous nannofossil zones are updated from Lamm et al.³⁰ and Coccioni and Frontalini²⁹.

Cyclostratigraphy

Cyclostratigraphic analyses for MS and ARM datasets were conducted with Acycle (version 2.4.1⁶⁴) and Astrochron (version 1.2^39,40). All MS and ARM datasets were log-transformed to harmonize fluctuations⁶⁵ before the resampling process every 2 cm. After interpolation, the datasets were detrended using a second-order polynomial curve. Spectral analysis was conducted using the prolate multitaper spectral estimator, comparing it against a robust red noise null model, with confidence levels set at the mean, 90%, 95%, 99%, and 99.9%. The MS and ARM series carry a long-term trend (Million-year scale band¹⁸) with high amplitude, leading to power leakage from the low-frequency components into the 405-kyr cycle; therefore, both series were high-passed (cutoff of 0.02 cycles/m). In the SAR analysis, MSB signals were removed to suppress the very high amplitude for the low frequencies.

Gaussian bandpass filters (0.02 and 0.35 cycles/m) were applied to isolate the interpreted long-eccentricity component sinusoidal curve from the datasets tuned prior to the tuning process according to the g₂–g₅ target curve from the La2004 astronomical solution⁴¹. Additionally, we performed evolutionary Fast Fourier Transform (eFFT) analyses⁵⁷ to observe the continuity of the spectral frequencies along the MS and ARM series. The sediment accumulation rates were evaluated throughout the PLG using COCO/eCOCO³⁸ and ASM^39,40. We conducted the COCO/eCOCO analysis by using 2,000 Monte Carlo simulations, applying sedimentation rates ranging from 0.26 to 10 cm/kyr referenced to the La2004 solution with a median age of 115 Ma (maximum frequency of 0.06/kyr), employing the Pearson correlation method.

Additionally, we utilized the time–frequency implementation of the ASM from the Astrochron package, which allows for quantitative testing of the orbital influence on sedimentation. We used the vector of the candidate astronomical cycles observed in the MS data spectrum (0.44, 1.27, 1.71, 4.47, 7.81, 9.4), the vector of the theoretical astronomical frequencies (1/405, 1/125, 1/95, 1/38.7, 1/22.9, 1/18.4), a Nyquist frequency of 25 cycles/m, and sedimentation rates ranging from 0.26 to 4.0 cm/ka (100 sedimentation rates). We conducted 100,000 Monte Carlo simulations for significance testing. These results enabled us to reject the null hypothesis that the observed spectrum was derived from a random signal with a high degree of confidence.

Supplementary Information

Below is the link to the electronic supplementary material.

Supplementary Material 1^{(129.1KB, xlsx)}

Supplementary Material 2^{(1.3MB, pdf)}

Acknowledgements

This study was financially supported by the Specialization and Postgraduate Program of Petrobras. J.M.F. RAMOS thanks Petrobras for Ph.D. process. The paper is a part of the projects Processamento e interpretação de dados magnetoestratigráficos do Cretáceo das Bacias Brasileiras and Magneto-cicloestratigrafia do Cretáceo na Margem Equatorial Brasileira (MEQ), both financed by Petróleo Brasileiro S.A.—Petrobras (FAURGS 8368 and 8892). J.F.S. also thanks CNPq for grants #304022/2018–7 and #311231/2021–7. D.R.F. thanks the Foundation Carlos Chagas Filho Research Support of the State of Rio de Janeiro (FAPERJ–grant #E-26/200.931/2022) and CNPq (grant #314462/2020-1). The authors sincerely thank Scientific Reports Assistant Editor Simran Gavhane and Editorial Board Member Annalisa Ferretti for their suggestions and reviews. We also thank Dr. Mike Simmons and the two anonymous reviewers for their valuable suggestions and comments during the peer-review process.

Author contributions

J.M.F.R.: conceptualization, methodology, validation, cyclostratigraphic analyses, astronomical tuning, and writing. J.F.S.: conceptualization, methodology, validation, paleomagnetic analysis, and investigation, writing, project administration, and funding acquisition. D.R.F.: methodology, cyclostratigraphic and astronomical tuning analyses, writing. M.F.: geochemical analysis and writing. C.G.L.: cyclostratigraphic and astronomical tuning analyses, writing. F.F.: conceptualization, biostratigraphical analysis, and writing. R.C.: conceptualization, biostratigraphical analysis, and writing. H. M.: geochemical analysis and writing. N.C.: conceptualization, biostratigraphical analysis, and writing. L.R.T.: geochemical analysis and writing. L.J.: sedimentary analysis and writing. R.I.F.T.: conceptualization, methodology, validation, paleomagnetic analysis, investigation, writing, project administration, and funding acquisition.

Data availability

The MS and ARM datasets used in this study are available in the Zenodo repository https:doi.org/10.5281/zenodo.6383426 .

Code availability

All mathematical treatments were performed with the freeware Acycle designed by ref.⁶⁴ and Astrochron R Package ref.^39,40.

Declarations

Competing interests

The authors declare no competing interests.

Footnotes

Publisher’s note

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References

1.Gale, A. S., Mutterlose, J. & Batenburg, S. In Geologic Time Scale 2020, (eds. Gradstein, F. M., Ogg, J. G., Schmitz, M. D., Ogg, G. M.) 1023–1068.
2.Judd, E. J. et al. A 485-million-year history of Earth’s surface temperature. Science385(1316), 1–9 (2024). [DOI] [PubMed] [Google Scholar]
3.Wignall, P. B. Large igneous provinces and mass extinctions. Earth Sci. Rev.53, 1–33 (2001). [Google Scholar]
4.Large igneous province record through time and implications for secular environmental changes and geological time-scale boundaries. Ernst, R.E., Bond, D.P. et al. From: Ernst, R.E., Dickson, A.J. and Bekker, A. (eds) Large Igneous Provinces: a Driver of Global Environmental and Biotic Changes. AGU, Geophysical Monographs255: 3–26 (2021).
5.Radiometric constraints on the timing, tempo, and effects of large igneous province emplacement. Kasbohm, J., Schoene, B. and Burgess, S. From: Ernst, R.E., Dickson, A.J. and Bekker, A. (eds) Large Igneous Provinces: A Driver of Global Environmental and Biotic Changes. AGU, Geophysical Monographs255: 27–82 (2021).
6.Jiang, Q., Jourdan, F., Olierook, H. K. & Merle, R. E. An appraisal of the ages of Phanerozoic large igneous provinces. Earth Sci. Rev.237, 104314 (2023). [Google Scholar]
7.Coccioni, R. et al. The neglected history of Oceanic Anoxic Event 1b: insights and new data from the Poggio le Guaine section (Umbria–Marche Basin). Stratigraphy11, 245–282 (2014). [Google Scholar]
8.Matsumoto, H. et al. Mid-Cretaceous marine Os isotope evidence for heterogeneous cause of oceanic anoxic events. Nat. Commun.13, 239 (2022). [DOI] [PMC free article] [PubMed] [Google Scholar]
9.Giorgioni, M. et al. Paleoceanographic changes during the Albian-Cenomanian in the Tethys and North Atlantic and the onset of the Cretaceous chalk. Glob. Planet. Change126, 46–61 (2015). [Google Scholar]
10.Schlanger, S. O. & Jenkyns, H. C. Cretaceous oceanic anoxic events: causes and consequences. Geol. en. Mijnb.55, 179–184 (1976). [Google Scholar]
11.Simmons, M. D., et al. Phanerozoic eustasy. In Geologic Time Scale 2020, (eds. Gradstein, F. M., Ogg, J. G., Schmitz, M. D., Ogg, G. M.) 357–387 (2020).
12.Conrad, C. P. The solid Earth’s influence on sea-level. Geol. Soc. Am. Bull.125(7–8), 1027–1052. 10.1130/B30764.1 (2013). [Google Scholar]
13.Dummann, W., et al. The impact of Early Cretaceous gateway evolution on ocean circulation and organic carbon burial in the emerging South Atlantic and Southern Ocean basins. Earth Planet. Sci. Lett.530, 115890. 10.1016/j.epsl.2019.115890 (2020).
14.Ray, D. C., et al. The magnitude and cause of short-term eustatic Cretaceous sea-level change: A synthesis. Earth-Sci. Rev.197, 102901 (2019). 10.1016/j.earscirev
15.Bottini, C. & Erba, E. Mid-Cretaceous paleoenvironmental changes in the western Tethys. Clim. Past14, 1147–1163 (2018). [Google Scholar]
16.Li, Y. et al. Radioisotopic chronology of Ocean Anoxic Event 1a: Framework for analysis of driving mechanisms. Sci. Adv.10, eadn8365 (2024). [DOI] [PMC free article] [PubMed]
17.Matsumoto, H. et al. Multidisciplinary evidence for synchroneity between Ontong Java Nui volcanism and early Aptian oceanic anoxic event 1a. Sci. Adv.11 (9), 10.1126/sciadv.adt0204. [DOI] [PMC free article] [PubMed]
18.Leandro, C. G. et al. Astronomical tuning of the Aptian stage and its implications for age recalibrations and paleoclimatic events. Nat. Commun.13 (2022). [DOI] [PMC free article] [PubMed]
19.Sabatino, N. et al. Mercury anomalies in upper Aptian-lower Albian sediments from the Tethys realm. Palaeogeogr. Palaeoclimatol. Palaeoecol.495, 163–170 (2018). [Google Scholar]
20.Coccioni, R. et al. Umbria-Marche Basin, Central Italy: A Reference Section for the Aptian-Albian Interval at Low Latitudes. Sci. Drill.13, 42–46 (2012). [Google Scholar]
21.Arthur, M. A., Brumsack, H. J., Jenkyns, H. C., & Schlanger, S. O. In. Cretaceous Resources, Events and Rhythms, 75–119. (Springer, 1990). Erba, E. et al. Environmental consequences of Ontong Java Plateau and Kerguelen Plateau volcanism. In The origin, evolution, and environmental impact of oceanic large igneous provinces. Geol. Soc. Am. Spec. Pap.511.
22.Ramos, J. M. F. et al. Anoxic Event 1b and its implications for the cause of mid-Cretaceous events: a multiproxy record. Paleoceanogr. Paleoclimatol.39(11), 1–30 (2024). [Google Scholar]
23.McAnena, A. et al. Atlantic cooling associated with a marine biotic crisis during the mid-cretaceous period. Nat. Geosci.6(7), 558–561 (2013). [Google Scholar]
24.Overview of Cretaceous Oceanic Red Beds (CORBs): a Window on Global Oceanic and Climate Change. Wang, C. et al. From: SEPM Society for Sedimentary Geology 91, Cretaceous Oceanic Red Beds: Stratigraphy, Composition, Origins, and Paleoceanographic and Paleoclimatic Significance: 13–33 (2009).
25.Dummann, W. et al. The early opening of the Equatorial Atlantic gateway and the evolution of Cretaceous peak warming. Geology51(5), 476–480 (2023). [Google Scholar]
26.Savian, J. et al. The Barremian-Aptian boundary in the Poggio le Guaine core (central Italy): evidence for magnetic polarity Chron M0r and oceanic anoxic event 1a. Geol. Soc. Am. Spec. Pap.524, 57–78 (2016). [Google Scholar]
27.Sabatino, N. et al. High-resolution chemostratigraphy of the late Aptian–early Albian oceanic anoxic event (OAE 1b) from the Poggio le Guaine section (Umbria–Marche Basin, central Italy). Palaeogeogr. Palaeoclimatol. Palaeoecol.426, 319–333 (2015). [Google Scholar]
28.Coccioni, R. & Galeotti, S. The mid-Cenomanian Event: prelude to OAE 2. Palaeogeogr. Palaeoclimatol. Palaeoecol.190, 427–440 (2003). [Google Scholar]
29.Coccioni, R., & Frontalini, F. Cretaceous planktonic foraminiferal biostratigraphy of the Umbria–Marche Basin (Central Italy) In: Hart, M.B., Batenburg, S.J., Huber, B.T., Price, G.D., Thibault, N., Wagreich, M. and Walaszczyk, I. (eds) Cretaceous Project 200 Volume 2: Regional Studies. Geological Society, London, Special Publications, 545(1), SP545–2023, 10.1144/SP545-2023-87 (2024).
30.Lamm, F. et al. Calcareous nannofossils from the Poggio le Guaine core (Umbria-Marche Basin, central Italy): Biostratigraphy and discussions on the bioevents of the Aptian–Albian interval. Cretac. Res.145, 1–19 (2023). [Google Scholar]
31.Tateo, F. et al. Orbital control on pelagic clay sedimentology: The case of the late Albian “Amadeus Segment” (central Italy). Bull. Soc. Geol. Fr.171(2), 217–228 (2020). [Google Scholar]
32.Li, Y. et al. Revised onset age of magnetochron M0r: Chronostratigraphic and geologic implications. Geology51(6), 565–570 (2023). [Google Scholar]
33.Selby, D., Mutterlose, J. & Condon, D. J. U-Pb and Re-Os Geochronology of the Aptian/Albian and Cenomanian/Turonian stage boundaries: implications for timescale calibration, osmium isotope seawater composition and Re-Os systematics in organic-rich sediments. Chem. Geol.265, 394–409 (2009). [Google Scholar]
34.Bornemann, A. et al. A first high-resolution carbon isotope stratigraphy from the Boreal (NW Germany) for the Berriasian to Coniacian interval—implications for the timing of the Aptian-Albian boundary. Front. Earth Sci.11, 1173319 (2023). [Google Scholar]
35.Coccioni, R. & Premoli Silva, I. Revised Upper Albian-Maastrichtian planktonic foraminiferal biostratigraphy and magneto-stratigraphy of the classical Tethyan Gubbio section (Italy). Newsl. Stratigr.48(1), 47–90 (2015). [Google Scholar]
36.Tiraboschi, D. et al. Origin of rhythmic Albian black shales (Piobbico core, central Italy): Calcareous nannofossil quantitative and statistical analyses and paleoceanographic reconstructions: Paleoceanography24 (2009).
37.Waltham, D. Milankovitch period uncertainties and their impact on cyclostratigraphy. J. Sediment. Res.85, 990–998 (2015). [Google Scholar]
38.Li, M. et al. Tracking variable sedimentation rates and astronomical forcing in Phanerozoic paleoclimate proxy series with evolutionary correlation coefficients and hypothesis testing. Earth Planet. Sci. Lett.501, 165–179 (2018). [Google Scholar]
39.Meyers, S. R. astrochron: An R Package for Astrochronology. https://cran.r-project.org/package=astrochron (2014).
40.Meyers, S. R. & Sageman, B. Quantification of deep-time orbital forcing by average spectral Misfit. Am. J. Sci.307, 773–792 (2007). [Google Scholar]
41.Laskar., J. et al. A long-term numerical solution for the insolation quantities of the Earth. Astron. Astrophys.428, 261–285 (2004).
42.Laskar, J. et al. La2010: A new orbital solution for the long-term motion of the Earth. Astron. Astrophys.532, A89. 10.1051/0004-6361/201116836 (2011). [Google Scholar]
43.Laskar, J. et al. Strong chaos induced by close encounters with ceres and vesta. Astron. Astrophys.532, L4. 10.1051/0004-6361/201117504 (2011). [Google Scholar]
44.Zeebe, R. E. & Lourens, L. J. Geologically constrained astronomical solutions for the Cenozoic era. Earth Planetary Sci. Lett.592, 117595.10.1016/j.epsl.2022.117595 (2022).
45.The Mid-Cretaceous Organic-Rich Sediments from the Vocontian Zone of the French Southeast Basin. Bréhéret, JG. (1994). In: Mascle, A. (eds) Hydrocarbon and Petroleum Geology of France. Special Publication of the European Association of Petroleum Geoscientists, vol 4. Springer, Berlin, Heidelberg.
46.Laskar, J. Astrochronology. In The Geologic Time Scale 2020, (eds. Grandstein, F. M., Ogg, J. G., Schmitz, M. D. & Ogg, G.) 139–158 (Elsevier, 2020).
47.Ma, C., Meyers, S. R. & Sageman, B. B. Theory of chaotic orbital variations confirmed by Cretaceous geological evidence. Nature542(7642), 468–470 (2017). [DOI] [PubMed] [Google Scholar]
48.Helsley, C. E. & Steiner, M. B. Evidence for long intervals of normal polarity during the Cretaceous period. Earth Planet. Sci. Lett.5, 325–332 (1968). [Google Scholar]
49.Zhang, Y. et al. Magnetostratigraphy of U-Pb–dated boreholes in Svalbard, Norway, implies that magnetochron M0r (a proposed Barremian-Aptian boundary marker) begins at 121.2 ± 0.4 Ma. Geology49 (2021).
50.Huang, C., Hinnov, L. A., Fischer, A. G., Grippo, A. & Herbert, T. Astronomical tuning of the Aptian stage from Italian reference sections. Geology238, 899–903 (2010). [Google Scholar]
51.Malinverno, A., Erba, E. & Herbert, T. D. Orbital tuning as an inverse problem: Chronology of the early Aptian oceanic anoxic event 1a (Selli Level) in the Cismon APTICORE. Paleoceanogr. Paleocl.25, PA2203 (2010).
52.Bottini, C. et al. Climate variability and ocean fertility during the Aptian Stage. Clim. Past11, 383–402 (2015). [Google Scholar]
53.Herrle, J. O. et al. High-resolution carbon isotope records of the Aptian to lower Albian from SE France and the Mazagan Plateau (DSDP Site 545): A stratigraphic tool for paleoceanographic and paleobiologic reconstruction. Earth Planet. Sci. Lett.218, 149–161 (2004). [Google Scholar]
54.Coccioni, R. & Galeotti, S. Orbitally induced cycles in benthonic foraminiferal morphogroups and trophic structure distribution patterns from the Late Albian “Amadeus Segment” (Central Italy). J. Micropalaeontol.12, 227–239 (1993). [Google Scholar]
55.Li, X. et al. Quantitative analysis of iron oxide concentrations within Aptian-Albian cyclic oceanic red beds in ODP Hole 1049C, North Atlantic. Sediment. Geol.235, 91–99 (2011). [Google Scholar]
56.Hu, X. et al. Cretaceous oceanic red beds (CORBs): Different time scales and models of origin. Earth Sci. Rev.115, 217–248 (2012). [Google Scholar]
57.Kodama, K. P. & Hinnov, L. A. Rock Magnetic Cyclostratigraphy Vol. 5 (John Wiley & Sons, 2014).
58.Mutterlose, J., Bornemann, A. & Herrle, J. O. The Aptian–Albian cold snap: Evidence for mid Cretaceous icehouse interludes. Neues Jb. Geol. Paläontol. Abh.232, 217–225 (2009). [Google Scholar]
59.Lam, A. R. et al. Diachroneity rules the mid-latitudes: a test case using Late Neogene Planktic Foraminifera across the Western Pacific. Geosciences12, 1–43 (2022). [Google Scholar]
60.Ramos, J. M. F., Savian, J. F., Franco, D. R., Figueiredo, M. F., Leandro, C. G., & Frontalini, F. Orbital tuning of short reversed geomagnetic polarity intervals in the Cretaceous Normal Polarity Superchron. Geophys. Res. Lett.51, e2024GL110530. 10.1029/2024GL110530. (2024).
61.Matsumoto, H. et al. Marine OS isotopic evidence for multiple volcanic episodes during Cretaceous Oceanic Anoxic Event. Sci. Rep.10, 12601 (2020). [DOI] [PMC free article] [PubMed] [Google Scholar]
62.Huber, B. T. & Leckie, R. M. Planktic foraminiferal species turnover across deep–sea Aptian/Albian boundary sections. J. Foraminiferal Res.41, 53–95 (2011). [Google Scholar]
63.Matsumoto, H. et al. Long-term Aptian marine osmium isotopic record of Ontong Java Nui activity. Geology49, 1148–1152 (2021). [Google Scholar]
64.Li, M., Hinnov, L. & Kump, L. Acycle: Time-series analysis software for paleoclimate research and education. Comput. Geosci.127, 12–22 (2019). [Google Scholar]
65.Herrle, J. O. et al. Mid-cretaceous High Arctic stratigraphy, climate, and oceanic anoxic events. Geology43(5), 403–406 (2015). [Google Scholar]

Associated Data

This section collects any data citations, data availability statements, or supplementary materials included in this article.

Supplementary Materials

Supplementary Material 1^{(129.1KB, xlsx)}

Supplementary Material 2^{(1.3MB, pdf)}

Data Availability Statement

The MS and ARM datasets used in this study are available in the Zenodo repository https:doi.org/10.5281/zenodo.6383426 .

All mathematical treatments were performed with the freeware Acycle designed by ref.⁶⁴ and Astrochron R Package ref.^39,40.

[CR1] 1.Gale, A. S., Mutterlose, J. & Batenburg, S. In Geologic Time Scale 2020, (eds. Gradstein, F. M., Ogg, J. G., Schmitz, M. D., Ogg, G. M.) 1023–1068.

[CR2] 2.Judd, E. J. et al. A 485-million-year history of Earth’s surface temperature. Science385(1316), 1–9 (2024). [DOI] [PubMed] [Google Scholar]

[CR3] 3.Wignall, P. B. Large igneous provinces and mass extinctions. Earth Sci. Rev.53, 1–33 (2001). [Google Scholar]

[CR4] 4.Large igneous province record through time and implications for secular environmental changes and geological time-scale boundaries. Ernst, R.E., Bond, D.P. et al. From: Ernst, R.E., Dickson, A.J. and Bekker, A. (eds) Large Igneous Provinces: a Driver of Global Environmental and Biotic Changes. AGU, Geophysical Monographs255: 3–26 (2021).

[CR5] 5.Radiometric constraints on the timing, tempo, and effects of large igneous province emplacement. Kasbohm, J., Schoene, B. and Burgess, S. From: Ernst, R.E., Dickson, A.J. and Bekker, A. (eds) Large Igneous Provinces: A Driver of Global Environmental and Biotic Changes. AGU, Geophysical Monographs255: 27–82 (2021).

[CR6] 6.Jiang, Q., Jourdan, F., Olierook, H. K. & Merle, R. E. An appraisal of the ages of Phanerozoic large igneous provinces. Earth Sci. Rev.237, 104314 (2023). [Google Scholar]

[CR7] 7.Coccioni, R. et al. The neglected history of Oceanic Anoxic Event 1b: insights and new data from the Poggio le Guaine section (Umbria–Marche Basin). Stratigraphy11, 245–282 (2014). [Google Scholar]

[CR8] 8.Matsumoto, H. et al. Mid-Cretaceous marine Os isotope evidence for heterogeneous cause of oceanic anoxic events. Nat. Commun.13, 239 (2022). [DOI] [PMC free article] [PubMed] [Google Scholar]

[CR9] 9.Giorgioni, M. et al. Paleoceanographic changes during the Albian-Cenomanian in the Tethys and North Atlantic and the onset of the Cretaceous chalk. Glob. Planet. Change126, 46–61 (2015). [Google Scholar]

[CR10] 10.Schlanger, S. O. & Jenkyns, H. C. Cretaceous oceanic anoxic events: causes and consequences. Geol. en. Mijnb.55, 179–184 (1976). [Google Scholar]

[CR11] 11.Simmons, M. D., et al. Phanerozoic eustasy. In Geologic Time Scale 2020, (eds. Gradstein, F. M., Ogg, J. G., Schmitz, M. D., Ogg, G. M.) 357–387 (2020).

[CR12] 12.Conrad, C. P. The solid Earth’s influence on sea-level. Geol. Soc. Am. Bull.125(7–8), 1027–1052. 10.1130/B30764.1 (2013). [Google Scholar]

[CR13] 13.Dummann, W., et al. The impact of Early Cretaceous gateway evolution on ocean circulation and organic carbon burial in the emerging South Atlantic and Southern Ocean basins. Earth Planet. Sci. Lett.530, 115890. 10.1016/j.epsl.2019.115890 (2020).

[CR14] 14.Ray, D. C., et al. The magnitude and cause of short-term eustatic Cretaceous sea-level change: A synthesis. Earth-Sci. Rev.197, 102901 (2019). 10.1016/j.earscirev

[CR15] 15.Bottini, C. & Erba, E. Mid-Cretaceous paleoenvironmental changes in the western Tethys. Clim. Past14, 1147–1163 (2018). [Google Scholar]

[CR16] 16.Li, Y. et al. Radioisotopic chronology of Ocean Anoxic Event 1a: Framework for analysis of driving mechanisms. Sci. Adv.10, eadn8365 (2024). [DOI] [PMC free article] [PubMed]

[CR17] 17.Matsumoto, H. et al. Multidisciplinary evidence for synchroneity between Ontong Java Nui volcanism and early Aptian oceanic anoxic event 1a. Sci. Adv.11 (9), 10.1126/sciadv.adt0204. [DOI] [PMC free article] [PubMed]

[CR18] 18.Leandro, C. G. et al. Astronomical tuning of the Aptian stage and its implications for age recalibrations and paleoclimatic events. Nat. Commun.13 (2022). [DOI] [PMC free article] [PubMed]

[CR19] 19.Sabatino, N. et al. Mercury anomalies in upper Aptian-lower Albian sediments from the Tethys realm. Palaeogeogr. Palaeoclimatol. Palaeoecol.495, 163–170 (2018). [Google Scholar]

[CR20] 20.Coccioni, R. et al. Umbria-Marche Basin, Central Italy: A Reference Section for the Aptian-Albian Interval at Low Latitudes. Sci. Drill.13, 42–46 (2012). [Google Scholar]

[CR21] 21.Arthur, M. A., Brumsack, H. J., Jenkyns, H. C., & Schlanger, S. O. In. Cretaceous Resources, Events and Rhythms, 75–119. (Springer, 1990). Erba, E. et al. Environmental consequences of Ontong Java Plateau and Kerguelen Plateau volcanism. In The origin, evolution, and environmental impact of oceanic large igneous provinces. Geol. Soc. Am. Spec. Pap.511.

[CR22] 22.Ramos, J. M. F. et al. Anoxic Event 1b and its implications for the cause of mid-Cretaceous events: a multiproxy record. Paleoceanogr. Paleoclimatol.39(11), 1–30 (2024). [Google Scholar]

[CR23] 23.McAnena, A. et al. Atlantic cooling associated with a marine biotic crisis during the mid-cretaceous period. Nat. Geosci.6(7), 558–561 (2013). [Google Scholar]

[CR24] 24.Overview of Cretaceous Oceanic Red Beds (CORBs): a Window on Global Oceanic and Climate Change. Wang, C. et al. From: SEPM Society for Sedimentary Geology 91, Cretaceous Oceanic Red Beds: Stratigraphy, Composition, Origins, and Paleoceanographic and Paleoclimatic Significance: 13–33 (2009).

[CR25] 25.Dummann, W. et al. The early opening of the Equatorial Atlantic gateway and the evolution of Cretaceous peak warming. Geology51(5), 476–480 (2023). [Google Scholar]

[CR26] 26.Savian, J. et al. The Barremian-Aptian boundary in the Poggio le Guaine core (central Italy): evidence for magnetic polarity Chron M0r and oceanic anoxic event 1a. Geol. Soc. Am. Spec. Pap.524, 57–78 (2016). [Google Scholar]

[CR27] 27.Sabatino, N. et al. High-resolution chemostratigraphy of the late Aptian–early Albian oceanic anoxic event (OAE 1b) from the Poggio le Guaine section (Umbria–Marche Basin, central Italy). Palaeogeogr. Palaeoclimatol. Palaeoecol.426, 319–333 (2015). [Google Scholar]

[CR28] 28.Coccioni, R. & Galeotti, S. The mid-Cenomanian Event: prelude to OAE 2. Palaeogeogr. Palaeoclimatol. Palaeoecol.190, 427–440 (2003). [Google Scholar]

[CR29] 29.Coccioni, R., & Frontalini, F. Cretaceous planktonic foraminiferal biostratigraphy of the Umbria–Marche Basin (Central Italy) In: Hart, M.B., Batenburg, S.J., Huber, B.T., Price, G.D., Thibault, N., Wagreich, M. and Walaszczyk, I. (eds) Cretaceous Project 200 Volume 2: Regional Studies. Geological Society, London, Special Publications, 545(1), SP545–2023, 10.1144/SP545-2023-87 (2024).

[CR30] 30.Lamm, F. et al. Calcareous nannofossils from the Poggio le Guaine core (Umbria-Marche Basin, central Italy): Biostratigraphy and discussions on the bioevents of the Aptian–Albian interval. Cretac. Res.145, 1–19 (2023). [Google Scholar]

[CR31] 31.Tateo, F. et al. Orbital control on pelagic clay sedimentology: The case of the late Albian “Amadeus Segment” (central Italy). Bull. Soc. Geol. Fr.171(2), 217–228 (2020). [Google Scholar]

[CR32] 32.Li, Y. et al. Revised onset age of magnetochron M0r: Chronostratigraphic and geologic implications. Geology51(6), 565–570 (2023). [Google Scholar]

[CR33] 33.Selby, D., Mutterlose, J. & Condon, D. J. U-Pb and Re-Os Geochronology of the Aptian/Albian and Cenomanian/Turonian stage boundaries: implications for timescale calibration, osmium isotope seawater composition and Re-Os systematics in organic-rich sediments. Chem. Geol.265, 394–409 (2009). [Google Scholar]

[CR34] 34.Bornemann, A. et al. A first high-resolution carbon isotope stratigraphy from the Boreal (NW Germany) for the Berriasian to Coniacian interval—implications for the timing of the Aptian-Albian boundary. Front. Earth Sci.11, 1173319 (2023). [Google Scholar]

[CR35] 35.Coccioni, R. & Premoli Silva, I. Revised Upper Albian-Maastrichtian planktonic foraminiferal biostratigraphy and magneto-stratigraphy of the classical Tethyan Gubbio section (Italy). Newsl. Stratigr.48(1), 47–90 (2015). [Google Scholar]

[CR36] 36.Tiraboschi, D. et al. Origin of rhythmic Albian black shales (Piobbico core, central Italy): Calcareous nannofossil quantitative and statistical analyses and paleoceanographic reconstructions: Paleoceanography24 (2009).

[CR37] 37.Waltham, D. Milankovitch period uncertainties and their impact on cyclostratigraphy. J. Sediment. Res.85, 990–998 (2015). [Google Scholar]

[CR38] 38.Li, M. et al. Tracking variable sedimentation rates and astronomical forcing in Phanerozoic paleoclimate proxy series with evolutionary correlation coefficients and hypothesis testing. Earth Planet. Sci. Lett.501, 165–179 (2018). [Google Scholar]

[CR39] 39.Meyers, S. R. astrochron: An R Package for Astrochronology. https://cran.r-project.org/package=astrochron (2014).

[CR40] 40.Meyers, S. R. & Sageman, B. Quantification of deep-time orbital forcing by average spectral Misfit. Am. J. Sci.307, 773–792 (2007). [Google Scholar]

[CR41] 41.Laskar., J. et al. A long-term numerical solution for the insolation quantities of the Earth. Astron. Astrophys.428, 261–285 (2004).

[CR42] 42.Laskar, J. et al. La2010: A new orbital solution for the long-term motion of the Earth. Astron. Astrophys.532, A89. 10.1051/0004-6361/201116836 (2011). [Google Scholar]

[CR43] 43.Laskar, J. et al. Strong chaos induced by close encounters with ceres and vesta. Astron. Astrophys.532, L4. 10.1051/0004-6361/201117504 (2011). [Google Scholar]

[CR44] 44.Zeebe, R. E. & Lourens, L. J. Geologically constrained astronomical solutions for the Cenozoic era. Earth Planetary Sci. Lett.592, 117595.10.1016/j.epsl.2022.117595 (2022).

[CR45] 45.The Mid-Cretaceous Organic-Rich Sediments from the Vocontian Zone of the French Southeast Basin. Bréhéret, JG. (1994). In: Mascle, A. (eds) Hydrocarbon and Petroleum Geology of France. Special Publication of the European Association of Petroleum Geoscientists, vol 4. Springer, Berlin, Heidelberg.

[CR46] 46.Laskar, J. Astrochronology. In The Geologic Time Scale 2020, (eds. Grandstein, F. M., Ogg, J. G., Schmitz, M. D. & Ogg, G.) 139–158 (Elsevier, 2020).

[CR47] 47.Ma, C., Meyers, S. R. & Sageman, B. B. Theory of chaotic orbital variations confirmed by Cretaceous geological evidence. Nature542(7642), 468–470 (2017). [DOI] [PubMed] [Google Scholar]

[CR48] 48.Helsley, C. E. & Steiner, M. B. Evidence for long intervals of normal polarity during the Cretaceous period. Earth Planet. Sci. Lett.5, 325–332 (1968). [Google Scholar]

[CR49] 49.Zhang, Y. et al. Magnetostratigraphy of U-Pb–dated boreholes in Svalbard, Norway, implies that magnetochron M0r (a proposed Barremian-Aptian boundary marker) begins at 121.2 ± 0.4 Ma. Geology49 (2021).

[CR50] 50.Huang, C., Hinnov, L. A., Fischer, A. G., Grippo, A. & Herbert, T. Astronomical tuning of the Aptian stage from Italian reference sections. Geology238, 899–903 (2010). [Google Scholar]

[CR51] 51.Malinverno, A., Erba, E. & Herbert, T. D. Orbital tuning as an inverse problem: Chronology of the early Aptian oceanic anoxic event 1a (Selli Level) in the Cismon APTICORE. Paleoceanogr. Paleocl.25, PA2203 (2010).

[CR52] 52.Bottini, C. et al. Climate variability and ocean fertility during the Aptian Stage. Clim. Past11, 383–402 (2015). [Google Scholar]

[CR53] 53.Herrle, J. O. et al. High-resolution carbon isotope records of the Aptian to lower Albian from SE France and the Mazagan Plateau (DSDP Site 545): A stratigraphic tool for paleoceanographic and paleobiologic reconstruction. Earth Planet. Sci. Lett.218, 149–161 (2004). [Google Scholar]

[CR54] 54.Coccioni, R. & Galeotti, S. Orbitally induced cycles in benthonic foraminiferal morphogroups and trophic structure distribution patterns from the Late Albian “Amadeus Segment” (Central Italy). J. Micropalaeontol.12, 227–239 (1993). [Google Scholar]

[CR55] 55.Li, X. et al. Quantitative analysis of iron oxide concentrations within Aptian-Albian cyclic oceanic red beds in ODP Hole 1049C, North Atlantic. Sediment. Geol.235, 91–99 (2011). [Google Scholar]

[CR56] 56.Hu, X. et al. Cretaceous oceanic red beds (CORBs): Different time scales and models of origin. Earth Sci. Rev.115, 217–248 (2012). [Google Scholar]

[CR57] 57.Kodama, K. P. & Hinnov, L. A. Rock Magnetic Cyclostratigraphy Vol. 5 (John Wiley & Sons, 2014).

[CR58] 58.Mutterlose, J., Bornemann, A. & Herrle, J. O. The Aptian–Albian cold snap: Evidence for mid Cretaceous icehouse interludes. Neues Jb. Geol. Paläontol. Abh.232, 217–225 (2009). [Google Scholar]

[CR59] 59.Lam, A. R. et al. Diachroneity rules the mid-latitudes: a test case using Late Neogene Planktic Foraminifera across the Western Pacific. Geosciences12, 1–43 (2022). [Google Scholar]

[CR60] 60.Ramos, J. M. F., Savian, J. F., Franco, D. R., Figueiredo, M. F., Leandro, C. G., & Frontalini, F. Orbital tuning of short reversed geomagnetic polarity intervals in the Cretaceous Normal Polarity Superchron. Geophys. Res. Lett.51, e2024GL110530. 10.1029/2024GL110530. (2024).

[CR61] 61.Matsumoto, H. et al. Marine OS isotopic evidence for multiple volcanic episodes during Cretaceous Oceanic Anoxic Event. Sci. Rep.10, 12601 (2020). [DOI] [PMC free article] [PubMed] [Google Scholar]

[CR62] 62.Huber, B. T. & Leckie, R. M. Planktic foraminiferal species turnover across deep–sea Aptian/Albian boundary sections. J. Foraminiferal Res.41, 53–95 (2011). [Google Scholar]

[CR63] 63.Matsumoto, H. et al. Long-term Aptian marine osmium isotopic record of Ontong Java Nui activity. Geology49, 1148–1152 (2021). [Google Scholar]

[CR64] 64.Li, M., Hinnov, L. & Kump, L. Acycle: Time-series analysis software for paleoclimate research and education. Comput. Geosci.127, 12–22 (2019). [Google Scholar]

[CR65] 65.Herrle, J. O. et al. Mid-cretaceous High Arctic stratigraphy, climate, and oceanic anoxic events. Geology43(5), 403–406 (2015). [Google Scholar]

PERMALINK

A novel chronostratigraphic framework for the Aptian–Albian paleoclimate events

João M F Ramos

Jairo F Savian

Daniel R Franco

Milene F Figueiredo

Carolina G Leandro

Rodolfo Coccioni

Fabrizio Frontalini

Nicola Casadei

Hironao Matsumoto

Leonardo R Tedeschi

Luigi Jovane

Ricardo I F Trindade

Abstract

Supplementary Information

Introduction

Fig. 1.

Fig. 2.

Results and discussion

Integrated stratigraphy and biozones

Magnetic records and the Aptian–Albian astrochronology

Fig. 3.

Fig. 4.

Table 1.

Table 2.

Astronomically dating Chron M0r, OAEs, organic-rich marker beds and CORBs

Fig. 5.

CORBs and temperature during Aptian–Albian stages

Aptian–Albian bioevents

Ages, timespans, and consequences of unradiogenic shifts

Age model for Aptian–Albian events

Fig. 6.

Materials and methods

Geological setting and sampling strategy

Rock magnetism

Planktonic foraminifera and calcareous nannofossils

Cyclostratigraphy

Supplementary Information

Acknowledgements

Author contributions

Data availability

Code availability

Declarations

Competing interests

Footnotes

References

Associated Data

Supplementary Materials

Data Availability Statement

ACTIONS

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RESOURCES

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