# Lateritic Cenozoic paleoenvironmental and paleoclimatic conditions in the Central Amazon basin, Brazil, inferred from mineralogy, geochemistry and geochronology

C. Ansart, D. Guinoiseau, C. Quantin, D. Calmels, C. Gautheron, P. Agrinier, J. Bouchez, Z. Fekiacova, R. Pinna-Jamme, D. Demri, et al.

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C. Ansart, D. Guinoiseau, C. Quantin, D. Calmels, C. Gautheron, et al. Lateritic Cenozoic pa-leoenvironmental and paleoclimatic conditions in the Central Amazon basin, Brazil, inferred from mineralogy, geochemistry and geochronology. Palaeogeography, Palaeoclimatology, Palaeoecology, 2025, 662, pp.112674. 10.1016/j.palaeo.2024.112674. hal-05221618

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# Lateritic Cenozoic paleoenvironmental and paleoclimatic conditions in the Central Amazon basin, Brazil, inferred from mineralogy, geochemistry and geochronology

C. Ansart<sup>a</sup>, D. Guinoiseau<sup>a,*</sup>, C. Quantin<sup>a</sup>, D. Calmels<sup>a</sup>, C. Gautheron<sup>b</sup>, P. Agrinier<sup>c</sup>, J. Bouchez<sup>c</sup>, Z. Fekiacova<sup>d</sup>, R. Pinna-Jamme<sup>a</sup>, D. Demri<sup>a</sup>, E. Balan<sup>e</sup>, B. Heller<sup>f</sup>, Q. Bollaert<sup>e</sup>, Ö. Ataytür<sup>e</sup>, T. Allard<sup>e</sup>

## ARTICLE INFO

Editor: H Falcon-Lang

Keywords:

Stable isotope geochemistry

Tropical weathering

Amazon basin

Laterite

Kaolinite

Fe-oxide

## ABSTRACT

Lateritic profiles are natural archives recording the weathering history in the tropics over million-year time-scales. This study combines oxygen, hydrogen, and silicon isotope data with mineralogical, geochronological, and geochemical analyses of lateritic secondary mineral assemblages to constrain the environmental and climatic conditions prevailing in the central Amazon Basin during the Cenozoic. The studied lateritic profile was developed over the sedimentary succession of the Alter do Chão Formation in Manaus, Brazil. Three distinct weathering episodes were identified, each constrained by (U—Th)/He and EPR ages. Dated to the Oligocene, the earliest (Oligocene) and long-lasting (duration >10 Ma) weathering episode involves the formation of well-ordered kaolinites through *in situ* chemical weathering of parent minerals under well-drained conditions, indicative of a tropical climate with restricted seasonality. The second episode, occurring in the mid-Miocene (~16 ± 3 Ma), involves the formation of a ferruginous duricrust. Most probably resulting from lateral iron migration and precipitation at the oxidizing front of an oscillating water table, the duricrust developed in a context of rising sea levels. The third episode, during the Upper Miocene (~10 Ma), led to the replacement of the kaolinite initially precipitated at the top of the profile by lower-crystallinity kaolinite formed at faster kinetics. This episode coincides with the final phase of the Andean uplift and the onset of the transcontinental Amazon River, reflecting enhanced water drainage under a "monsoon-type" climate.

## 1. Introduction

The Amazonas basin has remained in equatorial position for at least the last 100 Myr and has thus endured a warm and humid climate over this whole period of time (Tardy et al., 1991). Together with the intra-cratonic position of the basin, the ensuing drainage conditions in a transport limited regime have enabled the development and the preservation of thick weathering covers, i.e. lateritic profiles (from the saprolite to the topsoil), rich in secondary minerals such as Fe- or Al-(oxyhydr)oxides and clays. Those lateritic profiles have formed over

millions of years and occupy 80 % of the Brazilian Amazon sedimentary basin (i.e., the Amazonas basin) nowadays (Kotschoubey, 1981; Bar-dossy and Aleva, 1990; Costa, 1991; Tardy and Roquin, 1998).

The Cenozoic represents a noticeable period of geodynamic activity in South America and particularly in its western part, due to the rise of the Andes since the Oligocene with the Altiplano uplift up to its present elevation from the mid-Miocene (Peru) to the Pliocene (Colombia) (Roddaz et al., 2009; Eude et al., 2015; Hurtado et al., 2018; Sundell et al., 2019; Bicudo et al., 2020, Mora et al., 2008, Hooghiemstra et al., 2006). This major event is expected to have led to significant

<footnote><sup>a</sup> Université Paris-Saclay, CNRS, GEOPS, 91405 Orsay, France</footnote>

<footnote><sup>b</sup> ISTERRE, Université Grenoble Alpes, USMB, CNRS, IRD, UGE, Grenoble, France</footnote>

<footnote><sup>c</sup> Institut de Physique du Globe de Paris, Université de Paris, CNRS, 75005-F Paris, France</footnote>

<footnote><sup>d</sup> Aix Marseille Univ, CNRS, IRD, INRAE, Coll France, CEREGE, Aix-en-Provence, France</footnote>

<footnote><sup>e</sup> IMPMC, UMR CNRS 7590, Sorbonne Université, IRD, MNHM, 75252 Paris Cedex 05, France</footnote>

<footnote><sup>f</sup> Muséum National d'Histoire Naturelle, 57 rue Cuvier, 75005 Paris, France</footnote>

<footnote>* Corresponding author.</footnote>

<footnote>E-mail address: damien.guinoiseau@universite-paris-saclay.fr (D. Guinoiseau).</footnote>
paleoenvironmental changes, such as the early to mid-Miocene marine incursion associated to the mega-wetland Pebas system formation in the western part of the basin (23–10 Ma) (Hovikoski et al., 2011; Hoorn et al., 2022), or the onset of the actual transcontinental Amazon river system at ca. 9 Ma (Figueiredo et al., 2009; Hoorn et al., 2010; van Soelen et al., 2017; Hurtado et al., 2018; Sacek et al., 2023). These major events have probably affected climate, at least locally, and thus have modified the environmental conditions prevailing in the weathering zone in the basin. In this context, the geochemical information recorded in lateritic secondary minerals (e.g. hematite, goethite, kaolinite, or gibbsite) is frequently used to assess the role of climate in laterite formation through time (Ramanaidou, 1996; Girard et al., 1997; Girard et al., 2002; Giral-Kacmarcik et al., 1998; Yapp, 1998; Balan et al., 2005; Horbe, 2011; Allard et al., 2018; Mathian et al., 2019, 2022; Gautheron et al., 2022; Heller et al., 2022). While gibbsite and goethite precipitate under high water activity, hematite and kaolinite tend to form in less humid and hotter conditions (Tardy and Nahon, 1985; Trolard and Tardy, 1989). For example, kaolinite typically crystallizes in equilibrium with the soil solution and is known to resist to any subsequent isotopic exchange once formed (Bird and Chivas, 1989; Sheppard and Gilg, 1996; Stern et al., 1997; Girard et al., 2000; Nahon, 2003). Therefore, the O and H isotope signatures of kaolinite (δ¹⁸O~K~ and δD~K~), both dependent on the temperature and composition of the crystallization fluid, are good proxies of paleotemperature and paleo-rainwater composition at the time of formation (Bird et al., 1993; Girard et al., 2000; Horbe, 2011). Additionally, the silicon isotope composition (δ³⁰Si) of kaolinite can be used to infer the kinetically-driven vs. equilibrium-driven nature of chemical weathering processes, which are themselves dependent on the rate of clay formation and water infiltration (Guinoiseau et al., 2021; Opfergelt et al., 2012; Steinhoefel et al., 2017).

The formation ages of secondary minerals constitute milestones of the temporal evolution of the lateritic profiles and represent key information when interpreting the paleoenvironmental archives hosted by these profiles. Over the last 30 years, techniques have been developed to determine the formation age of Mn-oxides s.l., Fe-oxides s.l., or kaolinite (Vasconcelos et al., 1994; Hénocque et al., 1998; Dammer et al., 1999; Pidgeon et al., 2004; Balan et al., 2005; Shuster et al., 2005a; Beauvais et al., 2008; Vasconcelos et al., 2013; Monteiro et al., 2014; Allard et al., 2018). Altogether, these dating constraints show that lateritic profiles mainly record discrete and intense weathering episodes, probably masking signs of any continuous and slower chemical weathering, and potentially erasing the initiation of weathering through a rejuvenation process (Balan et al., 2005; Retallack, 2010; Vasconcelos et al., 2015; Mathian et al., 2019; Ansart et al., 2022). Indeed, the chronological interpretation of a weathering cover can be complex due to the occurrence of several generations of the same mineral in the profile or in a given horizon, resulting from successive dissolution – reprecipitation processes in response to changes in weathering conditions such as drainage conditions, rainfall intensity, temperature, redox (Giral-Kacmarcik et al., 1998; Balan et al., 2007; Dublet et al., 2015; Riffel et al., 2016; Allard et al., 2018; Mathian et al., 2020a; Ansart et al., 2022). Coupling isotopic and geochronological data of secondary minerals is therefore a promising approach to enhance our understanding of the weathering history of lateritic profiles and more generally of the evolution of continental surfaces.

Collectively, several geochronological, mineralogical, geochemical, and isotopic investigations led in the Central Amazon basin have provided a first insight of the regional evolution of weathering over the Cenozoic (Fritsch et al., 2002; Balan et al., 2005; Fritsch et al., 2005; Balan et al., 2007; Allard et al., 2020; Guinoiseau et al., 2021; Gautheron et al., 2022; Horbe et al., 2022). In this region, lateritic profiles have been suggested to result from, at least two major weathering episodes (Balan et al., 2005; Gautheron et al., 2022). Geochronological data on kaolinite (EPR ages) and Fe-(oxyhydr)oxides ((U-Th)/He ages) collected at several locations reveal the occurrence of an old and long weathering episode of moderate intensity, divided into two phases

(middle Eocene and from middle Eocene to Early Miocene), forming well-crystallized kaolinites in the northern part of the region, and a ferruginous duricrust in the southern part (Balan et al., 2005; Allard et al., 2020; Gautheron et al., 2022). While the existence of a generation of less ordered kaolinites was described in early studies (Lucas et al., 1996), the subsequent geochronologic, crystal chemistry and geochemical investigations revealed that the most recent (Late Miocene) weathering episode was shorter and more intense (i.e. with intense water drainage), and rejuvenated the profile from top to bottom with the formation of poorly-ordered kaolinites (Balan et al., 2005; Guinoiseau et al., 2021). Despite this improved understanding of the timing and duration of the weathering episodes responsible for the formation of lateritic profiles in the Central Amazon basin, the interpretation of the geochemical signals hosted by secondary minerals in terms of paleoenvironmental conditions is challenging and accurate time constrains are still lacking to confirm the features of the Manaus profiles. For example, the δ¹⁸O~K~ signature collected along a weathering profile near Manaus has been interpreted as resulting from present-day seasonal fluctuations (Giral-Kacmarcik et al., 1998). However, recent geochronological studies challenge this interpretation, with kaolinites ages from the central Amazon Basin ranging from Quaternary to Paleocene (Balan et al., 2005; Allard et al., 2018, 2020; Mathian et al., 2020).

In order to improve our understanding of the paleoenvironmental conditions having prevailed during the evolution of lateritic profiles in the Central Amazon basin, and to better constrain the chronology of their formation, our study focuses on a profile located near the town of Presidente Figueiredo, Brazil. This profile is developed over the Alter do Chão Formation and its structure and parent rock are similar to those of previously studied profiles in the region, with the additional presence of a consolidated iron crust. The presence of this duricrust suggests that an additional – as compared to previously studied profiles – and probably discrete weathering episode enabled significant iron remobilization, highlighting a local environmental change at the time of the crust formation. This Fe-duricrust layer may significantly impact drainage and weathering conditions through the profile, which should be carefully considered when studying lateritic profile. Hence, by combining oxygen, hydrogen and silicon isotope compositions of secondary minerals, together with geochronological data on Fe-(oxyhydr)oxides and kaolinite, we discuss the environmental conditions (precipitation regime, temperature, geomorphic changes) that prevailed during the three distinct weathering episodes which participated to the formation of lateritic profiles in the Central Amazon basin over the Cenozoic.

## 2. Climatic and geological settings

The studied regolith profile is located 100 km north of Manaus (Fig. 1) near the city of Presidente Figueiredo in the Central Amazon basin (2°03'S – 59°55'W; 115 m a.s.l.). This intracratonic basin separates the underlying Precambrian Amazonian basement into the outcropping Guyana and Brazilian shields (Putzer, 1984). In the region around Manaus, the lateritic cover is developed on dissected plateaus from the Cretaceous – Early Paleogene sedimentary Alter do Chão Formation (Fig. 1), which is widespread in the eastern part of the Amazon basin (Daemon, 1975; de Fátima Rossetti et al., 2005; Mendes et al., 2012). This formation, originating from the erosion of the surrounding areas of the Amazonian craton, consists of a several hundred-meters thick, heterogeneous sedimentary material (Caputo, 1984) with sandy and clayey intercalated beds, typical of a fluvial/lacustrine system, with rare red-dish and Fe-rich layers (Daemon, 1975; Sioli, 1984; Hoorn et al., 2009; Mendes et al., 2015). Palynological records suggest that the Alter do Chão Formation was deposited under a relatively humid climate (Jaramillo et al., 2011). Kaolinites dated from 65 to 20 Ma, hosted in weathering profiles developed on this Cretaceous formation, suggest that post-depositional weathering occurred (Irion, 1984; Balan et al., 2005). Weathering profiles are typically covered by up to 10 m of a thick kaolinite-rich Ferralsol, the so-called *Belterra Clay*, whose origin remains
<figure><img src="image_4.png" /><figcaption>Fig. 1. Simplified geological map of the Central Amazon basin (after Mendes et al., 2015; data from geosgb.cprm.gov.br) and location of the PF2 profile (white star). Study sites from Balan et al. (2005), Guinoiseau et al. (2021), Horbe (2011), Horbe et al. (2022) and Gautheron et al. (2022) are shown (circles) for comparison.</figcaption></figure>

debated. While a sedimentary origin, either lacustrine or continental, was first proposed (Sombroek, 1966; Truckenbrodt et al., 1991), Lucas et al. (1984), Chauvel et al. (1987) and Balan et al. (2007) interpreted this formation as the result of *in situ* weathering. Although the Alter do Chão Formation has not been capped by other sedimentary formations in the region of Manaus, its southern part in the central Amazon region is covered by scarce Oligocene and Miocene sediments (e.g., Iranduba and Novo Remanso Formations, respectively) (Dino et al., 2012; Gautheron et al., 2022), outcropping along the Amazon river bank. The study of these sedimentary formations indicates an Oligocene age for the Iranduba sediment, which was later weathered during the Oligo-Miocene period (28–18 Ma). The weathered Iranduba sediment was then covered by the Novo Remanso Formation during the Miocene (8–18 Ma) as proposed by Gautheron et al. (2022). Finally, the late Neogene to Quaternary sediments was deposited in unconformity on top of all these older sedimentary series.

The present-day climate can be considered as tropical rainforest climate with average rainfall of 3000 mm/yr and temperature ~ 26 °C. Climate seasonality, characterized by a monsoonal period from December to June and a short dry season from July to September resulting from the migration of the Inter-Tropical Convergence Zone (ITCZ), has persisted, at least since the Miocene Climate Optimum (Kaandorp et al., 2005). Unlike the African Craton, the Amazon basin has remained in an equatorial position for the last 100 Ma (Parrish et al., 1982; Tardy et al., 1991), such that warm and humid conditions, favourable to lateritic profile development, have prevailed over this period of time. Paleoenvironmental reconstructions have proposed that the Eastern Amazon and the Solimões basins have long been evolving separately due to the Purus Arch (Caputo and Soares, 2016). Based on sedimentary and palynological investigations, a marine incursion and

period of South-North water drainage orientation of the main hydro-
system (60–33 Ma) has been followed by the development of the Pebas
lacustrine system in the Western part of the basin (23–10 Ma). Although
the successive tectonic phases related to the Andean orogeny have
directly influenced the Amazonian landscape through time (Roddaz
et al., 2009; Hoorn et al., 2010, 2017; Albert et al., 2018; Sundell et al.,
2019; Bicudo et al., 2020; Sacek et al., 2023; Boschman, 2021), the most
recent uplift phases in the Altiplano or the Eastern Cordillera *ca.* 10–15
Ma changed drastically the regional atmospheric circulation regime,
creating a hydrographic barrier (Sepulchre et al., 2009; Insel et al.,
2010; Sacek et al., 2023), and enabling the onset of the transcontinental
Amazon route (~10 Ma) (Campbell et al., 2006; Hoorn et al., 2010;
Hovikoski et al., 2011; Albert et al., 2018). It has also been suggested
that this orogeny was responsible for a shift in the regional hydrological
cycle toward the modern seasonal precipitation regime (Kaandorp et al.,
2005; Hoorn et al., 2010), but this effect is still debated as no significant
change in precipitation intensity has been shown nor modeled over the
Cenozoic (Insel et al., 2012; Jeffery et al., 2012).

**3. Material and methods**

3.1. *Profile description, sampling and sample preparation*

The studied lateritic PF2 profile, has been sampled ~100 km north of
Manaus, Brazil, along a roadcut (2°03'S – 59°55'W; 115 m a.s.l.) near
Presidente Figueiredo. This 16-m thick profile is developed from the
Alter do Chão Formation that is still outcropping at the base of the
profile (Fig. 2). The profile structure is roughly similar to the one
described 50 km south by Lucas et al. (1996) and Balan et al. (2005) and
includes from base to top a saprolite, a mottled zone, and a Ferralsol. A
<figure><img src="image_5.png" /><figcaption>Fig. 2. Sketch and picture of the PF2 lateritic profile sampled near Presidente Figueiredo (PF2 profile). White circles on the sketch represent sampling locations, with sample names on the right and sampling depth (in m) on the left.</figcaption></figure>

noticeable difference is the presence of a duricrust at 9 m depth and a nodular layer made of ferruginous gravels at the bottom of the Ferralsol.

The soil and lateritic profile display significant contrasts of colour and texture. At the bottom of the profile, from 16 m to 13.5 m, the alternation of white sandy and clayey layers with rare reddish layers (i. e., PF2_1435 2.5YR 7/2: pale red and PF2_1350 5YR 8/2: pinkish white) is identified as the saprolite. The saprolite is overlaid by the mottled zone (around 13 m to 9.5 m depth) characterized by tubular bleached, yellow, and reddish domains. This unit is topped by a ferruginous lateritic duricrust characterized by a homogenous texture with vacuolar porosity. A Ferralsol, evolving from a red (2.5YR 5/6) at its base to a reddish yellow colour (7.5YR 7/6) at its top, overlies the duricrust up to the top. At ~7.5 m depth, a nodule-rich layer composed of hard, millimetric to centimetric Fe-nodules without any particular structure, is found embedded in the Ferralsol material.

A total of 17 samples were collected from the profile. Bulk samples were stored in plastic bags, dried at room temperature and sieved at 2 mm. In order to perform EPR dating on kaolinite particles, bulk samples gently crushed and sieved at 250 µm were purified by repeated cycles of citrate-bicarbonate-dithionite (CBD) extractions (3 to 10, depending on the Fe content) to remove Fe-oxides, s.l. (Mehra and Jackson, 1958). CBD-treated samples were then rinsed with deionized water and the <2 µm fraction was collected as the supernatant after smooth centrifugation.

For O and H isotope analyses of Fe-(oxyhydr)oxides, hand-picked and finely crushed aliquots from the duricrust (PF2_870) and nodules from the nodular zone (PF2_750) were cleaned of any traces of kaolinites using an ultrasonic bath, until the cleaning water was clear. The remaining kaolinite was finally removed with 5 M NaOH treatments according to Yapp (1991).

## 3.2. Chemical analyses

Major and trace element concentrations of bulk samples were measured at the SARM (Service d'analyse des Roches et des Minéraux, Vandoeuvre-lès-Nancy, France), following the procedure of Carignan

et al. (2001). In brief, after lithium tetraborate alkali fusion, major elements concentrations were measured by optical ICP (ICP-OES, iCap6500 ThermoFisher) while trace and rare earth elements were measured by quadrupole ICP-MS (iCapQ ThermoFischer). Concentration uncertainties range from 2 to 25 % and from 5 to 20 % for major and trace elements, respectively, depending on the measured concentrations (Table S1). Further information about analysis can be found at www.sarm.cnrs.fr.

Element concentrations of fifteen clay-sized fractions samples were determined by a quadrupole ICP-MS (ICP-QMS 7900 Agilent) after HF-HNO₃ acid digestion at the PARI (High-Resolution Analytical Platform) platform (Institut de Physique du Globe de Paris, France). The BHVO-2 basalt reference material was analyzed together with the studied samples to ensure accuracy of the measurements (Table S2).

Elemental losses or gains along the weathering profile were quantified by the mass-transfer coefficient (Brimhall and Dietrich, 1987; Braun et al., 2012), as follows:

$$ \tau_{x,ref} = \frac{C_{x,w} \cdot C_{i,ref}}{C_{x,ref} \cdot C_{i,w}} - 1 \qquad (1) $$

where Cₓ,w is the concentration of the considered element X in the regolith, Cₓ,ref is the concentration of this element in a reference material (typically the unweathered material underlying the weathering profile) and Cᵢ,w and Cᵢ,ref are the concentrations of an immobile element in the regolith and the reference material, respectively. Here, we chose as reference material the Upper Continental Crust (Taylor and McLennan, 1995) because the underlying Alter do Chão material is made of heterogenous sedimentary material originating from various rock formations of the surrounding cratonic areas, and is thus thought to be best characterized by an igneous-type crustal average.

A negative value of τ indicates a loss of X from the regolith, relative to the reference material, while a positive value indicates a relative enrichment in the considered element. A value of 0 reveals that the element is immobile. Among commonly used immobile elements, we chose niobium (Nb), considered as the least mobile element during
intense chemical weathering (Cornu et al., 1999; Voicu and Bardoux, 2002; Ma et al., 2007; Babechuk et al., 2014; Jiang et al., 2018).

### 3.3. Mineralogical characterization

The mineralogical composition of bulk samples was determined by X-ray diffraction using a Panalytical X'Pert Pro diffractometer (PANO-PLY platform, Université Paris Saclay, France) equipped with a Ni-filter Cu Kα radiation with an operating voltage of 45 keV, a beam current of 40 mA, a step size of 0.017° 2θ with a total counting time of 4 h and recorded between 3° 2θ and 80° 2θ. Mineral composition of the fine fraction was performed by X-ray diffraction at the IMPMC (Institut de Minéralogie, de Physique des Matériaux et de Cosmochimie, Sorbonne Université, France) with a Panalytical Pro MPD equipped with an X'Celerator detector and Co-Kα radiations with an operating voltage of 45 keV. The data were collected in a 2θ range of 3–110° with a step size of 0.017° and an acquisition time of 280 ms/step. Diffractograms were analyzed with Highscore software.

The semi-quantification of mineral phases in the clay-sized fraction of fifteen samples was performed using JaHDe, a deconvolution software using the Julia language developed at GEOPS (Université Paris Saclay, France; details in Text S1).

Spatially resolved observations and chemical analyses were performed by scanning electron microscope coupled with Energy Dispersive X-ray Spectroscopy (SEM-EDS) on a Phenom X Pro Fondis desktop SEM at 15 keV for 30 s acquisition time per point (PANOPLY platform, Université Paris Saclay, France). Because of their low cohesion, samples were first embedded in epoxy resin and observations were carried out on polished sections.

### 3.4. Isotope analyses

Isotope compositions are expressed in ‰, using NBS-28 quartz standard for δ³⁰Si analysis and VSMOW standard for δ¹⁸O and δD analysis:

$$ \delta X (‰) = \left( \frac{R_{sample}}{R_{standard}} - 1 \right) \cdot 1000 \quad (2) $$

With X = ³⁰Si, X = ¹⁸O, or X = D, and R = ³⁰Si/²⁸Si, R = ¹⁸O/¹⁶O, or R = D/H, respectively.

#### 3.4.1. Oxygen and hydrogen isotope composition

The oxygen and hydrogen isotope composition of kaolinite was obtained on 16 clay-sized fractions, which are mainly composed of

kaolinite (Table 1). For Fe-(oxyhydr)oxides, two aliquots for each of the two samples (PF2_750 and PF2_870), were analyzed for oxygen isotope composition using the same method as for kaolinite. All of these analyses have been performed at the PARI platform (Institut de Physique du Globe de Paris, France).

Oxygen isotope ratios were acquired with the fluorination method (Taylor, 2004). Seven mg of sample were reacted with BrF₅ at 600 °C overnight to convert the oxygen within kaolinite into O₂. The oxygen isotope ratios were measured on a Thermofisher Delta V isotope-ratio mass spectrometer (IR-MS). Measurements were calibrated by analysing internal quartz standards (TRIM Ba, TRIM 2b, NBS 28). The mean measured δ¹⁸O values for these standards (12.28 ± 0.16 ‰ 2σ n = 7, 11.39 ± 0.16 ‰ 2σ n = 5 and 9.70 ± 0.09 ‰ 2σ n = 2, respectively) agrees well with certified values (δ¹⁸O = 12.4 ‰, δ¹⁸O = 11.5 ‰, and δ¹⁸O = 9.58 ‰, respectively).

For hydrogen isotope analyses, 25 mg of sample were heated at 250 °C for 15 min prior to gas extraction to eliminate impurities adsorbed on kaolinite and quartz tubes. Samples were then heated at 1100 °C to release water from the structure, according to the method described in Groot (2009). The extracted H₂O was reduced in H₂ at 800 °C in a uranium furnace. The gas was finally concentrated by a Toepler pump and collected before measurement. The hydrogen isotope composition of the extracted gas was measured on an IR-MS Thermofischer Delta Plus XP. Water standards (HAWAI and EPB) were analyzed to calculate the experimental fractionation occurring during H₂O reduction to recalibrate the raw measured data. These standards have been provided by the LSCE (Laboratoire des Sciences du Climat et de l'Environnement, Saclay, France), and are regularly calibrated against the SMOW. The mean measured δD values of HAWAI and EPB standards were 11.84 ± 2.9 ‰ (2σ, n = 6) and -62.40 ± 1.5 ‰ (2σ, n = 6), respectively. Effective sample δD was thus recalculated using the calibration line defined by the standards as follows: δDcalibration. = 1.1036·δDmeasured + 19.072.

The efficiency of each method, for both kaolinite and Fe-(oxyhydr) oxides (i.e., hematite), was checked by calculating the gas effective extraction yields for O₂ or H₂. Extractions with yields >100 % or <90 % were discarded, and the protocol was repeated.

#### 3.4.2. Silicon isotope composition

Silicon isotopes analyses were performed on the bulk fraction samples after NaOH alkali fusion. Silicon purification was performed following the protocol of Fernandez et al. (2019). About 20 µg Si were loaded onto a column filled with a pre-cleaned 1.8 mL AG-50 W-X12 resin bed. The eluate was directly collected and combined with a further

<caption>**Table 1**</caption>

Clay-sized fraction mineralogical compositions (obtained using the JaHDE software), O and H isotope composition of kaolinite (δ¹⁸O<sub>K</sub> and δD<sub>K</sub>), corresponding O and H isotope composition for water in equilibrium with kaolinite during its precipitation (δ¹⁸O<sub>Kw</sub> and δD<sub>Kw</sub>) and crystallization temperatures (T<sub>c</sub>) calculated using Eq. (1) (Clauer et al., 2015). With Fer. = Ferralsol, Dur. = duricrust; MZ = mottled zone; ADC = Alter do Chão saprolite, Kln = kaolinite, Q = quartz, Gbs = gibbsite, Flc = florencite, Ant = anatase. *ND = No data.

<table><thead><tr><th>Samples</th><th>Depth (cm)</th><th>%Kln</th><th>%Q</th><th>%Gbs</th><th>%Flc</th><th>%Ant</th><th>δ¹⁸O<sub>K</sub> (‰)</th><th>δD<sub>K</sub> (‰)</th><th>δ¹⁸O<sub>Kw</sub>. (‰)</th><th>δD<sub>Kw</sub> (‰)</th><th>T<sub>c</sub> (°C)</th></tr></thead><tbody><tr><td>PF2_50</td><td>50</td><td>92</td><td>0</td><td>4.1</td><td>0</td><td>3.9</td><td>19.2</td><td>69.8</td><td>-5.8</td><td>-36.8</td><td>21.4</td></tr><tr><td>PF2_230</td><td>230</td><td>93</td><td>0</td><td>3.9</td><td>0</td><td>3.6</td><td>18.6</td><td>70.4</td><td>-5.9</td><td>-37.7</td><td>23.6</td></tr><tr><td>PF2_460</td><td>460</td><td>94</td><td>0</td><td>3.8</td><td>0</td><td>2.3</td><td>18.7</td><td>66.3</td><td>-5.4</td><td>-34.0</td><td>25.6</td></tr><tr><td>PF2_630</td><td>630</td><td>94</td><td>0</td><td>3.5</td><td>0</td><td>2.4</td><td>19.0</td><td>70.1</td><td>-5.8</td><td>-37.2</td><td>22.1</td></tr><tr><td>PF2_750</td><td>750</td><td>94</td><td>0</td><td>3.4</td><td>0</td><td>2.7</td><td>19.1</td><td>63.1</td><td>-5.1</td><td>-30.8</td><td>25.7</td></tr><tr><td>PF2_810</td><td>810</td><td>95</td><td>0</td><td>2.3</td><td>0</td><td>2.6</td><td>18.9</td><td>70.2</td><td>-5.9</td><td>-37.4</td><td>22.6</td></tr><tr><td>PF2_870</td><td>870</td><td>94</td><td>4.5</td><td>1.5</td><td>0</td><td>0</td><td>19.8</td><td>55.4</td><td>-4.1</td><td>-23.3</td><td>26.7</td></tr><tr><td>PF2_845</td><td>990</td><td>94</td><td>1.2</td><td>2.3</td><td>0.2</td><td>2.7</td><td>19.5</td><td>57.2</td><td>-4.3</td><td>-25.1</td><td>27.0</td></tr><tr><td>PF2_935</td><td>1080</td><td>95</td><td>3.9</td><td>0.5</td><td>0.2</td><td>0.4</td><td>20.1</td><td>55.9</td><td>-4.1</td><td>-23.5</td><td>25.0</td></tr><tr><td>PF2_1205</td><td>1350</td><td>ND*</td><td>ND*</td><td>ND*</td><td>ND*</td><td>ND*</td><td>20.3</td><td>61.2</td><td>-4.7</td><td>-28.2</td><td>21.2</td></tr><tr><td>PF2_1345</td><td>1490</td><td>98</td><td>0.4</td><td>0</td><td>0.5</td><td>1.0</td><td>20.3</td><td>60.8</td><td>-4.7</td><td>-27.8</td><td>21.8</td></tr><tr><td>PF2_1355</td><td>1500</td><td>94</td><td>5.0</td><td>0.4</td><td>0.3</td><td>0</td><td>20.7</td><td>58.6</td><td>-4.4</td><td>-25.6</td><td>21.3</td></tr><tr><td>PF2_1380</td><td>1525</td><td>93</td><td>6.1</td><td>0.5</td><td>0.3</td><td>0</td><td>21.0</td><td>56.1</td><td>-4.1</td><td>-23.1</td><td>21.1</td></tr><tr><td>PF2_1415</td><td>1560</td><td>94</td><td>1.4</td><td>0.2</td><td>1.3</td><td>2.7</td><td>20.3</td><td>59.8</td><td>-4.6</td><td>-27.0</td><td>22.3</td></tr><tr><td>PF2_1435</td><td>1580</td><td>89</td><td>9.7</td><td>0.4</td><td>0.8</td><td>0</td><td>20.5</td><td>54.6</td><td>-3.9</td><td>-22.0</td><td>24.2</td></tr><tr><td>PF2_1455</td><td>1600</td><td>94</td><td>5.6</td><td>0.4</td><td>0.1</td><td>0</td><td>21.3</td><td>56.9</td><td>-4.1</td><td>-23.6</td><td>19.4</td></tr></tbody></table>
volume of 5 mL milli-Q water passed through the resin. Silicon isotope ratio measurements were performed on a ThermoFinnigan Neptune Plus MC-ICP-MS at the PARI platform (Institut de Physique du Globe de Paris, France). Typical running conditions are reported in Fernandez et al. (2019) and Guinoiseau et al. (2021).

Mass-dependent fractionation was checked in a δ³⁰Si - δ²⁹Si diagram (Fig. S1). Long-term accuracy and reproducibility of Si isotope measurements were checked through repeated analyses of the BHVO-2 and BCR-2 reference materials, yielding δ³⁰Si = -0.27 ± 0.08 ‰ (2σ, n = 31) and δ³⁰Si = -0.23 ± 0.06 ‰ (2σ, n = 22), respectively, in agreement with previously reported values (Savage et al., 2013).

### 3.5. Secondary mineral dating

#### 3.5.1. (U—Th)/He analyses of hematite

The (U—Th)/He dating method relies on the ⁴He production during alpha decay of the ²³⁸U, ²³⁵U, ²³²Th chains and of ¹⁴⁷Sm, trapped within the crystal structure of different mineral phases including goethite and hematite (Farley, 2002; Shuster et al., 2005; Ault et al., 2019). The (U—Th)/He ages were determined by measuring the ⁴He content extracted from a subsample and then compared with the content in U, Th, and Sm, considering the decay constants of ²³⁵U, ²³⁸U, ²³²Th, and ¹⁴⁷Sm and of their radioactive daughter products. A correction of 5 ± 5 ‰ on calculated ages was applied to consider the He loss from the mostly-hematitic crystal structure since its crystallization time, using values proposed by Bassal et al. (2022).

A fragment of the duricrust (sample PF2_870) and of nodules (PF2_750) were roughly crushed, washed from clay minerals in an ultrasonic bath with milliQ-water for less than 30 min to avoid He loss (Vasconcelos et al., 2013), and cleaned with ethanol. The obtained material was observed with a binocular to select appropriate subsamples for (U—Th)/He dating (i.e., microscopically homogeneous and metallic shiny subsamples). A total of 6 subsamples of ca. 60–100 µg were selected and loaded into a pure Nb tube (purity 99.9 %). Helium, U, Th and Sm content were analyzed at the PANOPYL platform (GEOPS, Université Paris Saclay, France), following the procedure of Gautheron et al. (2021) adapted for Fe-(oxyhydr)oxides. Each capsule was heated with a Yb diode laser under vacuum for 30 min (Allard et al., 2018). The temperature was kept under 1000 °C to avoid actinide volatilization as observed on hematite (Danish et al., 2013; Hofmann et al., 2020). The extracted gas was mixed with a known amount of ³He and purified from H₂O, CO₂, H₂, and Ar gases. The remaining gas contents were determined with a Pfeiffer Prisma Quadrupole mass spectrometer. Encapsulated subsamples were collected and dissolved to measure the U, Th, and Sm contents, following the procedure of Heller et al. (2022), using an ICP-MS (Agilent) at the PARI platform (Institut de Physique du Globe de Paris, France). No alpha ejection correction was done due to the large size of the samples. A goethite internal standard with known age and U standard concentration (Verhaert et al., 2022) was analyzed to check for data accuracy. Analytical uncertainty is lower than 2 % (1σ) for U, Th, Sm, and He concentration and 5 % for (U—Th)/He age for which secular equilibrium in the U—Th series is assumed.

#### 3.5.2. Electron paramagnetic resonance dating of kaolinites

In absence of signal saturation, the age of kaolinite formation is calculated as the ratio of the paleodose over dose-rate. In the present study, the dating of kaolinites using EPR spectroscopy has been improved with respect to our previous works and we provide in the following part comprehensive information about an up to date methodology that is summarized in Allard et al. (submitted).

##### 3.5.2.1. Electron paramagnetic resonance spectroscopy. The crystal chemistry characterization of kaolinite populations within the profile was assessed using the electron paramagnetic resonance (EPR) spectrum of structural Fe³⁺ as a probe, which has been shown to be very sensitive

to any kind of crystalline disorder (Gaite et al., 1993; Balan et al., 1999, 2007). Sixteen spectra were recorded at ambient temperature on a Bruker EMX plus spectrometer at X-band hyper frequency (ca. 9.8 GHz) and in the 0–600 mT range. Analytical conditions were 0.5 mT, 100 kHz, and 40 mW for modulation amplitude, frequency, and microwave power, respectively. Spectra were normalized for sample mass, recording gain and a constant filling factor of the resonance cavity.

The assignment of EPR spectra to known species was verified using their effective and characteristic spectroscopic factor g values derived from the resonance condition:

$$h\nu = g \times \beta \times H_0 \quad (3)$$

where: h is the Planck constant; v is the resonance frequency; g is the Landé factor, which is a tensor factor; β is the Bohr magneton, and H₀ is the external magnetic field. For the radiation-induced defects (RIDs) spectra, the principal values of the g tensor (i.e., g// and g⊥) were compared with those of a reference kaolinite (GB1, Allard et al., 1994) spectra themselves initially calibrated using a 2,2-diphenyl-1-picrylhydrazyl (DPPH) standard (gDPPH = 2.0036) with an accuracy of 0.001.

Concentrations of RIDs in kaolinite samples were estimated from a simplified procedure (e.g., Balan et al., 2005) based upon the measurement of the amplitude of the perpendicular component, and expressed in arbitrary units that are proportional to a number of spins per unit of sample mass, with an error of ±10 ‰.

The concentration of structural Fe³⁺ in kaolinite can be estimated (in µg/g) from a double integration of part of the EPR spectrum and using a relevant kaolinite standard, with an error of 35 ‰, as detailed in Balan et al. (2000). These authors also showed that concentrations of structural trivalent iron are around 3000 µg/g, both in the Ferralsol and in the saprolite in a lateritic profile from the Manaus region.

##### 3.5.2.2. He⁺ irradiation. The clay-sized fraction samples were irradiated with 1.5 MeV He⁺ ions on the ARAMIS linear accelerator (IJCLab, Orsay, France) to simulate ionizing radiation effects (Allard and Muller, 1998). Samples were deposited and dried on stainless steel plates with a real thickness of 4 µm (i.e., excluding porosity), which is the range of high-energy He⁺ ions in kaolinite as calculated with the SRIM code (Ziegler et al., 2010). After irradiation, the samples were heated for 2 h at 250 °C to anneal the unstable defects, according to the protocol of Allard et al. (1994) confirmed by Ataytür et al. (2024). Measurement of RIDs concentrations allowed us to plot the dosimetry curves in the 3.10¹¹–10¹³ ions/cm² fluence range corresponding to doses within 70–2342 kGy, which is known to be appropriate for kaolinite dating in lateritic covers (Balan et al., 2005; Allard et al., 2018; Mathian et al., 2019, 2020, 2022).

##### 3.5.2.3. Dosimetry and dose rate determination. The kaolinite dating requires determination of the paleodose, i.e. the dose naturally received by the kaolinite, and dose rate independently. The paleodose is measured from experimental He⁺ dosimetry. Experimental data were fitted with two functions: a common exponential function (e.g., Allard et al., 1994) and an exponential + linear function (Duval, 2012) according to the following equations:

$$[C] = [Cs](1 - \exp(-\lambda(D + P))) \quad (4)$$

$$[C] = [Cs](1 - \exp(-\lambda(D + P))) + LF(D + P) \quad (5)$$

where [C] is the present concentration of RIDs, [Cs] is the saturation value, λ an efficiency exponential factor, D the artificial dose, LF is a linear factor and P the paleodose, which is the parameter required for dating. When the residues were close between the functions (3) and (4), a statistical Fisher Snedecor test was applied (see result section) to select the best formalism for the paleodose calculation (Allard et al., submitted).

The dose rate is calculated from the uranium (U), thorium (Th) and
potassium (K) contents measured by ICP-AES and ICP-MS techniques in bulk samples together with tabulated elementary values provided by Aitken (1985) or Guérin et al. (2011). Several corrections of the dose rate are required to consider various natural factors that influence the energy deposition of the radiations in kaolinites. These corrections are related to the U and Th microscopic distribution, which is relevant for alpha particles, the degree of aperture of the decay chains (assimilated as the radon (Rn) loss or excess), the effect of water in porosity (acting as a radiation absorber), the role on the alpha radiation dose of minerals that do not host radioelements such as quartz and the possible macroscopic effect of contrasting heterogeneities on the gamma dose rate.

In a closed and infinite system where radioelements are distributed in the clayey matrix, alpha particles are responsible for 90 % of the dose derived from the decay chains of U and Th. Therefore, U and Th mapping at microscopic scale may result in an important correction to the dose rate. In the case of large U or Th-bearing minerals (> 20 µm), the alpha particles whose ranges in silicates are 20 µm do not impact the bulk sample but only a thin rim around the U, Th- bearing mineral and thus are neglected. The correction does not affect beta and gamma rays whose radiation ranges are between ca. 2 mm and 30 cm, respectively. To perform U mapping, irradiations with thermal neutrons were carried out on several thin sections of bulk samples to induce fission tracks (e.g., Price and Walker, 1963). This was performed on the neutron reactor at IPEN, São Paulo, Brazil. Fission tracks were produced in a Kapton film placed in contact with the thin section (Sélo, 1983) and using a cumulative neutron flux of 6.9 × 10¹⁵ neutrons/cm². The fission tracks formed in the Kapton foils were etched with a hypochlorite solution (at least 5 %) at 100 °C for 8 min. Subsequently, the tracks were counted using an optical microscope with a x 500 or 1000 magnification. Homogeneity of the neutron flux at the scale of the sample was guaranteed by the IPEN team. The proportion of U in the clayey matrix and in the U-bearing minerals was then determined. (e.g., Allard et al., 2020; submitted).

The spatial distribution of Th, which had not been measured in previous studies, was assessed by X-ray fluorescence on the Nanoscopium beamline at the SOLEIL synchrotron facility, Saint-Aubin, France (Somogyi et al., 2013). Indurated thin sections on pure 5 × 5 mm silica glass slides were prepared. The incident energy of 20 keV was used to collect the X-ray fluorescence of many elements with atomic mass above Al. After large-scale imaging (ca. 650 × 250 µm) with a 1-µm resolution, a focus on 50 × 50 µm areas with a 150 nm beam size resolution was performed on all samples. Data were processed with the PyMca software to plot the X-ray fluorescence spectra with the fitted elementary contributions, allowing us to map Th and relevant elements.

The estimate of the disequilibrium in the U decay chain was derived from the gamma spectrometry analyses of raw samples. In our protocol this is assessed from the Rn loss or accumulation, as this element is a rare gas with no or poor interaction with matter and which can migrate in the weathering cover. A migration of Rn can have a substantial effect on the dose rate because its daughters contribute to around 60 % of the total dose rate of the whole U decay chain in a closed system. Bulk samples were analyzed by high-resolution, low-background gamma-ray spectrometry using a germanium detector at low temperature at the Institut de Paléontologie Humaine, Paris, for measurement of U and Th decay chains isotopes. Samples were packed in plastic boxes and stored for one month to equilibrate radon. To measure ²³⁸U activity, the gamma radiation arising from ²³⁴Pa (its daughter) at 63.3 keV was used. The ²³²Th activity was measured using those of ²²⁸Ra (²²⁸Ac at 911.1 keV) and ²²⁸Th (²⁰⁸Tl at 583.2 keV represents 35 % of disintegration of ²¹²Bi), respectively. In order to estimate the possible Rn loss or gain, the ²²⁶Ra was derived from the ray at 186.1 keV, and radon (²²²Rn) was determined using the peaks of ²¹⁴Bi at 609 keV and of ²¹⁴Pb at 352 keV (Yokoyama and Nguyen, 1980).

The water content, which is a secondary order correction, was taken as 20 %, from the year-averaged value provided by the study of Bruno et al. (2006) on laterites.

Quartz grains are usually inherited minerals in lateritic profiles. As they most often do not contain U or Th, they are associated to two competing effects on the dose rate. The first one is due to their diluting effect of the concentrations of U and Th. Indeed, concentrations of these elements must be normalized to the clayey matrix content and are thus increased when compared to the concentration in the bulk material. The second effect only affects alpha particles due to the absorption of the radiations, which lowers the effective dose rate in the clayey matrix. A simple geometrical model was used to assess the effects of quartz on the dose rate, which is based upon the average size of quartz grains and the range of alpha particles in silicates, their concentration and the concentration of kaolinite (Allard et al., 2020; submitted). The size of quartz grains was assessed using thin sections, by assimilating them to spherical particles for which the average radius can be estimated through imaging. The potential effect of different levels around the sample and mostly inherited from sedimentary structures was estimated from the volume contribution of each different level to the gamma dose rate in a 30 cm radius sphere around the sample.

# 4. Results

## 4.1. Mineralogy and chemistry of the bulk and clay-sized fractions

The mineralogical and chemical composition of the bulk samples is displayed in Fig. 3A and Table S1, respectively. Sandy (i.e., PF2_1455, PF2_1435, PF2_1380, PF2_1355, PF2_1350 and PF2_1205 samples) and clayey layers (i.e., PF2_1415 and PF2_1345 samples) of the Alter do Chão Formation (herein considered as the saprolite) are mainly composed of quartz and of large booklets of kaolinite (ca. 60–100 µm), with accessory occurrence (< 2 wt%) of mica or illite, anatase, zircon, and florencite, as well as hematite and goethite traces in the reddish layers (Fig. 3A). The overlying mottled zone (13–9.5 m) is characterized by bleached (Fe-poor) and reddish (Fe-rich/hematite-rich) domains where kaolinite dominates over quartz (80 wt% vs. 20 wt%). The weakly indurated Fe-rich lateritic duricrust (9.5–9 m) is hematitic with residual embedded quartz and kaolinite (Fig. 4C). The top 9 m Ferralsol unit is made of a kaolinite matrix with minor gibbsite, goethite, and hematite (Figs. 3, 4), and contains a level of centimetric to millimetric hematite-rich nodules at 7.5 m depth (Fig. 4B).

In the clay-sized fractions (major and trace element concentrations in Table S2), kaolinite represents more than 90 wt% of the material, while quartz, absent in the Ferralsol, never exceeds 10 wt%, even for clay-sized fractions separated from sand-rich layers of the saprolite (Table 1). Gibbsite is detected but never exceeds 5 wt% (Figs. 3, 4 and Table 1). Phosphate minerals (i.e., florencite) are found in the Alter do Chão saprolite and explain the enrichment of associated elements (i.e., La, Ce, Ba, Pb). All redox-sensitive elements (e.g., As, Fe, Mn), which are assumed to be carried by hematite or individualized as Mn-oxides are depleted, showing the efficiency of the CBD leaching protocol.

Enrichment/depletion of elements throughout the profile is reported as mass transfer coefficients in Fig. S2 and Fig. S3 for bulk and clay-sized fractions, respectively. Soluble elements (e.g., Na, Ca, Mg) are fully depleted in the early stage of weathering, i.e., in the saprolite ($\tau_{X,Nb} < -0.9$), while elements carried by phosphate minerals (i.e., La, Ce, Ba, Pb) are enriched in the saprolite and lost in the upper parts of the profile for bulk samples (Fig. S2B and S2E) and the clay-sized fractions (Fig. S3B and S3E). The $\tau_{X,Nb}$ values of elements that are classically considered immobile, such as Ti, Hf, Ta, Sc, are relatively constant over the whole profile, meaning that these elements roughly follow the same behavior as Nb (Fig. S2D). By contrast redox-sensitive elements (e.g., Cr, As, V) are strongly enriched in the duricrust and the nodular horizon (Table S1; Fig. S2C). Although the Ferralsol mainly contains kaolinites, it is strongly depleted in Si ($\tau_{Si,Nb} < -0.8$), most likely due to the depletion of Si-bearing minerals such as quartz and kaolinites, compared to the parent material.
<figure><img src="image_6.png" /><figcaption>**Fig. 3.** XRD spectra of (A) bulk PF2 samples and (B) clay-sized fraction of PF2 samples. With Ilt = illite; Kln = kaolinite; Gbs = gibbsite; Gth = goethite; Qz = quartz; Flc = florencite; Hem = hematite; Ant = Anatase. The PF2_1350 sample has not been purified due to the low amount of available material.</figcaption></figure>

## 4.2. Structural Fe³⁺ in kaolinite

The EPR spectra of purified kaolinites display signals at low magnetic field that are characteristic of structural Fe³⁺ (Fig. 5), i.e., diluted trivalent iron substituting for Al³⁺ in the octahedral layer (Gaite et al., 1993; Balan et al., 1999). Two different signals named Fe³⁺(I) and Fe³⁺(II) are recognized (Fig. 5), which correspond to disordered and ordered environments in the kaolinite structures, respectively. In the studied profile, the structural Fe³⁺ signal strongly varies with depth, in full agreement with what was previously observed in the Manaus region by Balan et al. (2005, 2007): kaolinite at depth in the saprolite is well crystallized, with

a dominant Fe³⁺(II) contribution, whereas kaolinite from the Ferralsol is highly disordered, with a dominant Fe³⁺(I) contribution (Fig. 5A). Some kaolinites in the profile exhibit spectra of structural Fe³⁺ that are intermediate between those observed in the topsoil and the saprolite. Consequently, kaolinite samples from the profile may be interpreted as a mixture between these two endmembers of kaolinite (Balan et al., 2007). The spectra of intermediate samples were linearly decomposed with a least-square fitting procedure, providing the proportion of each of the two endmembers, here chosen to be best represented by samples PF2_50 (top Ferralsol) and PF2_1455 (deep saprolite) (Fig. 6D). Corresponding fits were generally satisfactory (Table S3 and Fig. S4), although 3
<figure><img src="image_7.png" /><figcaption>**Fig. 4.** SEM-BSE images of kaolinite (Kln) in the different units of the profile. (A) The Ferralsol (e.g., sample PF2_460) is mainly composed of a matrix of kaolinite with few Fe-(oxyhydr)oxides. (B) In the nodular zone (sample PF2_750), the centimetric nodules are made of hematite (Hem) surrounded by a soft Fe-rich kaolinitic matrix, similar to that of the Ferralsol. (C) Corroded quartz (Qz) and large kaolinite booklets embedded in the hematitic duricrust (sample PF2_870). (D) Large kaolinite booklets in the Alter do Chão saprolite (e.g., sample PF2_1380).</figcaption></figure>

samples from the duricrust (PF2_870 sample) and clay-rich layers from the saprolite (PF2_1345 and PF2_1415 samples) showed a slight misfit indicating that at least one of the chosen endmembers may not be fully appropriate to reconstruct the sample spectrum: at least another (third) unknown population could be invoked. These three samples put aside, the results of the spectra decomposition show that the proportion of disordered kaolinite is close to 100 % in the top 4.5 m of the profile, then continuously decreasing down to 0 % in the saprolite at around 11 m. Kaolinite from the duricrust (PF2_870) is relatively well ordered, similarly to kaolinite from the saprolite (Fig. 5, 6D). This is consistent with XRD spectra (Fig. 3B) and SEM observations (Fig. 4C) that show pattern characteristic of high, long-range order (due to low amount of stacking defects) for kaolinites from the duricrust and saprolite. In contrast, kaolinites found in the surrounding material within the nodular zone are similar to those from the Ferralsol. In addition, the kaolinites trapped within the nodules analyzed after careful removal of the surrounding clay exhibit strong crystalline disorder, even slightly stronger than the kaolinites from the surrounding material. These observations indicate that kaolinites present in nodules and in the duricrust differ in their origin.

The structural Fe³⁺ concentrations in the PF2 profile (Fig. S5) vary with depth and exhibit a strong contrast between the Ferralsol and the saprolite (around 2500–3000 µg/g and around 1000 µg/g, respectively). In detail, kaolinites from the duricrust are the most concentrated in structural Fe³⁺ in the profile, and two samples (PF2_810 and PF2_935) located between the duricrust and the saprolite have intermediate structural Fe³⁺ content. The strong contrast between the Ferralsol and the saprolite in term of crystalline disorder is thus also observed for the

concentration of structural Fe³⁺, which is indicative of different conditions of formation. As stated in the discussion, the local disorder and the structural Fe³⁺ content in kaolinite allow us to single out different populations of kaolinite in the profile and thus different conditions of formation in the Ferralsol and saprolite.

### 4.3. Isotope signatures

#### 4.3.1. Oxygen and hydrogen isotopes

The oxygen and hydrogen isotope ratios measured in the clay-sized fraction, which is thought to represent the kaolinite O and H isotope composition (δ¹⁸O~K~ and δD~K~) range from 18.6 ‰ to 21.3 ‰ and –70.4 ‰ to –54.6 ‰, respectively (Table 1, Fig. 6A and B). The δ¹⁸O~K~ and δD~K~ signatures differ between the saprolite and the Ferralsol. In the mottled zone, δD values are similar to those found in the saprolite while δ¹⁸O values are significantly lighter. The δ¹⁸O values obtained here are consistent with those reported by Giral-Kacmarcík et al. (1998) in a lateritic profile also developed on Alter do Chão formation and sampled close to our study area (~50 km south). This observation implies that even if minor gibbsite is present in the Ferralsol (Fig. 3B), our δ¹⁸O data for the profile PF2 represent the kaolinite composition.

To confirm that the measured δ¹⁸O~K~ and δD~K~ values are only associated with kaolinite, we use the semi-quantitative constraint on the proportion of individual mineral phases from XRD spectra. Kaolinite represents between 89 and 98 % of the sample masses. The effect of florencite on δ¹⁸O and δD values can be discarded based on its low abundance in all samples (from 0 % in Ferralsol to 1.3 % in PF2_1415). The amount of quartz (from 0 to 9 %) is below the threshold where it can
<figure><img src="image_8.png" /><figcaption>**Fig. 5.** (A) Total X-band EPR spectra of clay-sized fractions from the PF2 profile. Note the strong contrast in structural Fe³⁺ signals and amplitude of RIDs between saprolite and Ferralsol. (B) Comparison of signals of structural Fe³⁺ of kaolinites between PF2_750 bulk (black), PF2_750 nodule (red) and PF2_50 bulk (aqua) samples EPR spectra. Note that PF2_750 nodule and bulk PF2_50 EPR spectra are similar, but different from that of PF2_750 bulk (see arrows). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

significantly influence the δ¹⁸O of the bulk sample (up to 27 ‰ according to Clauer et al., 2015). Based on the 4 ‰ maximum gibbsite contribution throughout the profile and on the 9 ‰ difference of δ¹⁸O reported for gibbsite and kaolinite formation at 20 °C (Clauer et al., 2015; Sheppard and Gilg, 1996), the kaolinite δ¹⁸O values corrected from the gibbsite component would not differ by more than ±0.3–0.4 ‰, which is below analytical error.

Hematite from sample PF2_870 and from nodules of samples PF2_750 exhibit distinct δ¹⁸O (δ¹⁸O<sub>Fe</sub>) signatures of –0.6 ‰ and –0.8 ‰ (averaging at –0.7 ‰), and –2.1 ‰ and –2.1 ‰ (averaging at –2.1 ‰) respectively.

The isotopic composition of water from which kaolinites precipitated (δ¹⁸O<sub>Kw</sub> and δD<sub>Kw</sub>) was calculated (Fig. 7) based on fractionation factors from Sheppard and Gilg (1996), assuming (1) isotopic equilibrium between water and kaolinite during kaolinite crystallization, (2) crystallization temperature T<sub>c</sub> representing the average temperature prevailing during kaolinite formation and (3) a similar global meteoric water line in the past than today, i.e., δD<sub>w</sub> = 8.0 · δ¹⁸O<sub>Kw</sub> + 10. The calculated values of δ¹⁸O<sub>Kw</sub> and δD<sub>Kw</sub> vary from –5.9 ‰ to –3.9 ‰ and from –37.7 ‰ to –22.0 ‰, respectively. For hematite, the δ¹⁸O value of parent water from which minerals were formed (δ¹⁸O<sub>FeW</sub>) has been calculated using the oxygen fractionation equations of Yapp (1990) and a mean temperature of 25 °C. The calculated δ¹⁸O<sub>FeW</sub> are –6.7 ‰ for PF2_870 and –8.2 ‰ for nodules from PF2_750. Considering estimated analytical uncertainties on δ¹⁸O values ≤2 ‰ (i.e., including measurement bias, temperature estimation uncertainties and isotopic composition of contaminant), δ¹⁸O<sub>FeW</sub> are in line with δ¹⁸O<sub>Kw</sub>.

In comparison, the present-day rainwater composition in the region of Manaus, over ten years, range between δ¹⁸O = –1.8 ‰ and δD = –8 ‰; δ¹⁸O = –1.4 ‰ and δD = –6 ‰ and δ¹⁸O = –4.9 ‰ and δD = –29 ‰

for the annual, the wet and the dry season averages, respectively (Honório et al., 2010), although time series highlight significant inter-annual variability (Table S4).

4.3.2. Silicon isotopes

Values of δ³⁰Si are almost constant across the sand-rich and kaolinite-rich layers of the saprolite, the mottled zone, and duricrust (Fig. 6C, δ³⁰Si = –0.09 ± 0.09 ‰). This value is close to that of the upper continental crust (δ³⁰Si = –0.25 ± 0.16 ‰; Savage et al., 2013). From a depth of 10 m to the surface, δ³⁰Si progressively shifts toward lighter values, reaching –2.00 ‰ at the top of the profile. To compare the results obtained for the bulk samples with previously reported δ³⁰Si values for the clay-sized fraction by Guinoiseau et al. (2021) in the same region, we estimated the clay δ³⁰Si values of the PF2 profile. In bulk samples and corresponding fine fractions, Si is carried entirely by quartz and kaolinite, while Al is shared between gibbsite and kaolinite. Due to the variability of kaolinite crystalline order across the profile, the semi-quantification of quartz vs. kaolinite proportion using X-ray diffraction data only is quite imprecise. We thus combined XRD patterns and element concentration to determine the clay δ³⁰Si values of kaolinite. First, the distribution of bulk Al and Si between kaolinite, quartz (only for Si) and gibbsite (only for Al in PF2_50, PF2_870, and PF2_845 samples) was determined using Al and Si concentrations (Table S1). Then, the δ³⁰Si of the kaolinite end-member in bulk samples was calculated assuming a fixed δ³⁰Si of the quartz endmember at –0.20 ‰ based on the measured δ³⁰Si values in sand-rich layers (e.g. PF2_1435 and PF2_1205 samples) and on the average δ³⁰Si values reported for quartz in previous studies (e.g., Poitrasson, 2017; Steinhoefel et al., 2017). Finally, using the kaolinite, gibbsite, and quartz proportions determined from XRD spectra of the fine fraction (Table 1), and assuming that the
<figure><img src="image_9.png" /><figcaption>**Fig. 6.** Clay-sized fractions (A) oxygen and (B) hydrogen isotopes, (C) bulk silicon isotope compositions and (D) calculated proportion of disordered kaolinites to total kaolinite content by linear decomposition of EPR spectra, within the weathering profile of Presidente Figueiredo. (E) Values of δ¹⁸O and (F) δ³⁰Si vs. proportion of disordered kaolinite. The δ¹⁸O values obtained by *Giral-Kacmarcik et al. (1998)* are reported in panel (A) (grey circles), with samples depth accommodated for comparison according to their description of the distinct layers in their studied lateritic profile.</figcaption></figure>

δ³⁰Si of the kaolinite and quartz end members are similar for bulk samples and fine fractions, we determined the clay δ³⁰Si values. The difference between the measured bulk and calculated clay δ³⁰Si is minimal (absolute value of Δ³⁰SiₜBulk-clay-sized fraction lower than 0.1 ‰, Fig. 8) and shows that the bulk δ³⁰Si values measured along the PF2 profile can be faithfully used as representative of those of the clay-sized fraction. The δ³⁰Si profiles in the present study and in *Guinoiseau et al. (2021)* agree well (Fig. 8). The lightest isotope composition observed near the surface of the "Site III" profile of *Guinoiseau et al. (2021)* in comparison with our data is likely due to a non-negligible effect of biological overturning (Lucas et al., 1993; Sutton et al., 2018).

## 4.4. Geochronology

### 4.4.1. (U–Th)/He of hematite

The 5% corrected (U–Th)/He ages of five hematite subsamples from the nodular zone (PF2_750) and of six hematite subsamples from the duricrust (PF2_870) are all of Miocene age and are homogeneous within a given sample (Fig. 9, Table S7). Nodules are younger (average age of 6.3 ± 0.9 Ma) than the hematite from the duricrust (average age of 15.0

± 1.6 Ma). Chemical analysis performed on these oxides show a larger Th/U ratio in nodules (Th/U ≈ 145) than in the duricrust (Th/U ≈ 43), which can be explained by distinct episodes of Fe-rich layers formation under different redox conditions (Riffel et al., 2016).

### 4.4.2. Kaolinite dating

#### 4.4.2.1. Determination of paleodoses: He⁺-irradiation experiments.
The dosimetry experiments are critical steps of the EPR dating approach allowing for the determination of the paleodose, which is the dose experienced by the clay sample in its environment since its formation. The growth curves representing the RID concentration as a function of dose were built using the amplitudes of the corresponding EPR spectra (Fig. S6) for the twelve samples selected in the He⁺ ion beam experiment. For all of them, the amplitude of RIDs (dominantly so-called A-centers, Clozel et al., 1994) is growing with increasing dose (expressed here as fluence in ions/cm²). The scale of the derivative intensity axis (arbitrary units) also provides a rough idea of the relative concentrations of defects, which is most often lower in the Ferralsol and higher in the
<figure><img src="image_10.png" /><figcaption>**Fig. 7.** Calculated δ¹⁸O values of mineral-forming water along the studied profile of Presidente Figueiredo.</figcaption></figure>

<figure><img src="image_11.png" /><figcaption>**Fig. 8.** Evolution of δ³⁰Si values of bulk samples (measured, grey dots) and of the fine fractions (calculated, black squares) in the lateritic profile of Presidente Figueiredo (PF) compared to those measured in the fine fraction of the “Site III” from Guinoiseau et al. (2021) (grey diamonds). EPR-calculated δ³⁰Si (yellow circles) have been obtained using the proportion of well-crystallized vs. poorly crystallized kaolinite populations (Figs. 5, 6) and assuming a δ³⁰Si value of -2 ‰ for poorly crystallized kaolinites (same as that of sample PF_50, the top profile) and of 0.03 ‰ for well-crystallized kaolinites (average of δ³⁰Si values calculated for the fine fraction of the saprolite). Note that the actual thickness of the different lateritic units described for the profile of “Site III” and Presidente Figueiredo differ. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

saprolite.

Subsequently, the concentration of RIDs as a function of dose (converted to kGys from fluences, see Allard et al., 1994) is plotted in Fig. 10

for the twelve samples. The fitting of these growth curves using exponential (exp.; Eq. (3)) or exponential + linear (exp. + lin.; Eq. (4)) functions were experimented for each set of points.

The fits were of good quality, as inferred from their residues (Table S5). Consequently, in order to select the formalism to be retained for paleodose determination, a Fisher Snedecor test was applied on the residue on one hand, and on the square of the difference between the model and the experimental value on the other hand, these parameters being normalized to the degree of freedom (the number of samples minus the number of parameters). With both statistical parameters, this analysis clearly shows that adding a linear component to the fit is not significantly better for samples of the upper part of the profile (from PF2.50 to PF2.845). Consequently, for these samples, the simpler exponential function (Eq. (4)) was selected to fit the experimental growth curve of RIDs (Fig. 10). The kaolinite PF2.750 from the nodules is an exception because the dosimetry function is linear. Such a linear growth curve has never been encountered so far in kaolinite samples from lateritic profiles or sediments (Balan et al., 2005; Allard et al., 2018, 2020; Mathian et al., 2019, 2020, 2022) and is thus exceptional. It implies that the corresponding paleodose derived from the linear function stands for a maximum value, as the initial exponential growth intercept with the dose axis would inevitably occur at a lower value.

In contrast, the samples from the lower part of the profile (PF2.935 to PF2.1455) show that adding a linear component to the exponential function yields significantly better fits (with 90 to 95 % confidence). Consequently, for these samples rich in well-ordered kaolinite, the paleodose was determined from the fit with the exp. + lin. Function (Eq. (4)). It is mentioned that the sample PF2.1380 has a very low linear slope and thus can be assimilated, in the investigated range of dose, to an exponential growth curve yielding the same result. In addition, because the dosimetry behavior apparently distinguishes the well crystallized (exp. + lin. function) from the disordered kaolinites (exp. function), the paleodose of the duricrust sample PF2.870 has tentatively been determined by the exp. + lin. Function, since similar well crystallized kaolinites have been observed in the saprolite and within the duricrust.

The fitting parameters are shown in Table S5, and the paleodoses are reported in Table 2. For four samples that have not been irradiated (PF2.230, 450, 1355 and 1435) a model growth curve from a close sample in terms of local disorder was considered to calculate the paleodose (Table 2).

Paleodoses vary within a wide range between 47 and 1282 kGy. The lowest paleodoses correspond to the two kaolinite-rich levels which present in addition uncommon chemical compositions and mineralogy compared to the saprolite (PF2_1345 and PF2_1415, composed of 98–94 % of kaolinite).

While the linear contribution may be related to trap creation during irradiation (Levy, 1989), the steep increase of the curve around the natural RID concentration (PF2_1345 or 1415) remains unclear. Such a particular behavior was also observed in previous studies of kaolinite from lateritic profiles (see e.g. Mathian et al., 2020) or of quartz from sediments (Voinchet et al., 2015, 2020). In addition, the exponential growth suggests the classical decoration of a limited and fixed amount of precursors in the kaolinite structure.

#### 4.4.2.2. Dose rate determination.
The corrections performed on the values of doses rates derived from U, Th, K concentrations are detailed hereafter. As mentioned in the method section, the corrections are related to the spatial distributions of U and Th (for the effects of alpha particles), the disequilibrium of the decay chains, the water content and the role of quartz grains (concerning alpha particles), and the effect of macroscopic heterogeneities on the gamma dose rate. The resulting dose rates are reported in the Table 2 and Table S5.

The spatial distribution of uranium at microscopic scale was assessed using the induced fission-track method on several samples, namely PF2.50, _630, _870, _935, _1345, _1380, allowing for the determination
<figure><img src="image_12.png" /><figcaption>**Fig. 9.** Photographs of the studied nodules (PF2_750) and duricrust (PF2_870) samples, with their 5 % corrected ages, Th/U mass ratios and (U—Th)/He ages distribution.</figcaption></figure>

of the concentration of U in the clayey matrix and in bearing minerals.

These distributions of U were taken as models for the other samples that are close in terms of composition or location within the profile. Fig. 11 shows examples of thin sections of PF2_50 and PF2_1380 samples and pictures of induced fission tracks (both on different regions). The PF2_50 samples contained 26 % of U in U-bearing minerals that appear as urchin-like clusters while U was fully scattered in the clayey matrix in the PF2_1380 sample, which is the only one for which the dosimetry behaves with a clear exponential growth.

In addition, because Th is by far more concentrated than U in all the samples (Table 2), it strongly influences the dose rate correction. Its spatial distribution, analyzed using X-ray fluorescence at the SOLEIL synchrotron facility (Saint Aubin, France), is presented herein for two contrasting samples, PF2_50 from the topsoil and PF2_1415 from the saprolite (Fig. 12). In the saprolite sample, thorium was present in small Ce-bearing phases interpreted as florencites according to the presence of such minerals in the XRD patterns of the raw samples. These florencites are most often smaller than one micron and relatively close to each other when compared to the 20 µm propagation range of alpha particles in silicates.

In the topsoil (PF2_50), small zircons were also found as Th-bearing minerals, together with some Ce phases. In both cases the Th-bearing minerals are much smaller than the mean alpha range, so they irradiate the clayey matrix, and thus the whole alpha dose rate from thorium was considered in the total budget. It appears that this analysis verifies the assumption on Th distribution proposed in Balan et al. (2005) for the Manaus profiles, who considered a homogeneous distribution of Th in the clayey matrix.

For the correction of the degree of aperture of the radioactive system, nine samples were analyzed that showed a slight accumulation of Rn with respect to U, except for PF2_810 and PF2_870 for which a loss of Rn was measured (Table 2, see details in Table S6). Corresponding effects of disequilibria on the actual dose rate were assumed to hold for the whole life of kaolinites.

The potential correction of gamma dose rate was studied on samples from the saprolite where inherited sedimentary levels are smaller than the 30 cm radius sphere of gamma energy deposition on the sample. It was verified in all cases that despite differences in U, Th, K concentrations in these levels, there is no significant effect on the total dose rate and the age. This is due to the minor contribution of gamma rays (a few %) in the total dose rate. Indeed, the alpha dose rate is by far dominant because of the large distribution of radioelements in the matrix, especially for Th which is relatively abundant.

4.4.2.3. *Ages of kaolinites.* The resulting ages vary between 1.1 ± 0.3 and 56.7 ± 17.0 Ma (Table 2). The oldest age, obtained in PF2_1435, is derived from a modeled paleodose, taking the PF2_1455 dosimetry parameters as a reference, and thus may not be fully reliable. Putting aside this sample, the ages of all kaolinites in the profile are plotted in Fig. 13 as a function of depth, and compared to the ages of hematite from the duricrust and the nodules determined by the (U—Th)/He method. Even considering the error bars, the overall trend shows a decrease in age from around 30 Ma at depth in the saprolite to around 10 Ma in the uppermost horizon. This trend is in line with the results from the Manaus profiles studied by Balan et al. (2005) and Guinoiseau et al. (2021) except that the topsoil was younger in these former studies (6–8 Ma), a
<figure><img src="image_13.png" /><figcaption>**Fig. 10.** He⁺ dosimetry curves retained for the determination of paleodose, which represents the intercept of the curve with dose axis.</figcaption></figure>
<caption>**Table 2**</caption>

EPR data and resulting EPR kaolinite ages. U, Th and K values are taken from the bulk samples composition. * and ** are paleodoses calculated from PF2_630 and PF2_1455 dosimetry parameters, respectively; *** maximum paleodose and age.

<table><thead><tr><th>Sample</th><th></th><th>Depth (m)</th><th>U (ppm)</th><th>Th (ppm)</th><th>K (%)</th><th>Paleodose (kGy)</th><th>Dose rate (mGy/ka)</th><th>Age (Ma)</th></tr></thead><tbody><tr><td>PF2_50</td><td></td><td>50</td><td>3.03</td><td>36.78</td><td>0</td><td>329</td><td>30,976</td><td>10.6 ± 3.2</td></tr><tr><td>PF2_230</td><td></td><td>230</td><td>2.85</td><td>35.02</td><td>0</td><td>360*</td><td>28,831</td><td>11.8 ± 3.5</td></tr><tr><td>PF2_460</td><td></td><td>460</td><td>2.91</td><td>34.3</td><td>0</td><td>459*</td><td>27,952</td><td>15.4 ± 4.6</td></tr><tr><td>PF2_630</td><td></td><td>630</td><td>3.37</td><td>36.6</td><td>0</td><td>495</td><td>30,751</td><td>15.0 ± 4.5</td></tr><tr><td>PF2_750 nod</td><td></td><td>750</td><td>2.06</td><td>54.69</td><td>0</td><td>1282***</td><td>52,448</td><td>22.9 ± 6.9</td></tr><tr><td>PF2_810</td><td>Fer.</td><td>810</td><td>2.46</td><td>29.01</td><td>0</td><td>536</td><td>23,139</td><td>21.8 ± 6.5</td></tr><tr><td>PF2_870</td><td>Dur.</td><td>870</td><td>0.44</td><td>12.06</td><td>0</td><td>174</td><td>7822</td><td>22.2 ± 6.7</td></tr><tr><td>PF2_845</td><td></td><td>990</td><td>3.18</td><td>28.81</td><td>0</td><td>758</td><td>25,911</td><td>27.0 ± 8.1</td></tr><tr><td>PF2_935</td><td>MZ</td><td>1080</td><td>2.33</td><td>17.49</td><td>0.0332</td><td>263</td><td>14,532</td><td>14.6 ± 4.4</td></tr><tr><td>PF2_1205</td><td></td><td>1350</td><td>0.78</td><td>9.14</td><td>0.0282</td><td>256</td><td>9795</td><td>26.2 ± 7.9</td></tr><tr><td>PF2_1345</td><td></td><td>1490</td><td>5.18</td><td>37.95</td><td>0.1300</td><td>132</td><td>32,763</td><td>3.6 ± 1.1</td></tr><tr><td>PF2_1355</td><td></td><td>1500</td><td>1.27</td><td>8.319</td><td>0.0473</td><td>255**</td><td>8768</td><td>25.4 ± 7.6</td></tr><tr><td>PF2_1380</td><td></td><td>1525</td><td>1.14</td><td>8.54</td><td>0.0523</td><td>271</td><td>9043</td><td>30.0 ± 9.0</td></tr><tr><td>PF2_1415</td><td></td><td>1560</td><td>5.3</td><td>47.93</td><td>0.1360</td><td>47</td><td>39,635</td><td>1.1 ± 0.3</td></tr><tr><td>PF2_1435</td><td></td><td>1580</td><td>0.89</td><td>13.93</td><td>0.0299</td><td>838**</td><td>13,771</td><td>56.7 ± 17.0</td></tr><tr><td>PF2_1455</td><td>ADC</td><td>1600</td><td>0.81</td><td>10.24</td><td>0.0423</td><td>232</td><td>9286</td><td>23.0 ± 6.9</td></tr></tbody></table>

<figure><img src="image_14.png" /><figcaption>**Fig. 11.** Examples of thin sections (a, c) and fission tracks (b, d) for the PF2_50 (a, b) and PF2_1380 (c, d) samples (thin sections and fission tracks pictures do not represent the same locations in the thin section). The PF2_50 sample contains U-bearing minerals (see the blue arrow in b) whereas the PF2_1380 sample only exhibits U in the clayey matrix. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

discrepancy that may be due to the more complete protocol of dating that was applied in the present study. It supports an older Oligocene population of kaolinite that was rejuvenated by a middle Miocene major process of weathering, as previously proposed. Nevertheless, when considering the uppermost horizon around 10 Ma, and knowing that EPR analysis encompasses a great number of clay-sized particles, it cannot be excluded that the kaolinite sample may also be a mixture of generations and that younger ones may be present in a small proportion.

As for crystalline disorder or concentration of structural trivalent iron (Fig. 6D, S5), the PF2_935 sample is intermediate between the Ferralsol and the bottom saprolite, which adds consistency to our data interpretation as a mixture of two kaolinite endmembers. However, in the saprolite, two samples (PF2_1345, PF2_1415) are much younger

than the surrounding material and correspond to Quaternary and Pliocene periods. The origin of these anomalies in the saprolite is discussed in section 5.2.1.

In addition, consistency with isotopic (Si, O, H) data would imply that these two clayey levels evolved recently (1–4 Ma) as closed systems that did not result in isotopic fractionation, which can be achieved if the whole material influenced the water chemistry and was recrystallized *in situ*. In this assumption, it is the structure of the saprolitic horizon that imposes the conditions of kaolinite formation, more than the climatic parameters. Alternatively, low paleodose and high dose rates could be related to a recent accumulation of Th, consistent with occurrence of florencite in the profile. However, such an anomalous accumulation of Th is not revealed by the mass transfer coefficients analysis ($\tau_{Th,Nb}$),
<figure><img src="image_15.png" /><figcaption>**Fig. 12.** X-ray fluorescence mapping for PF2–50 (a, b) and PF2–1415 (d, e) and related spectra (c, f). Th-bearing minerals in PF2–50 are interpreted as zircons and occasionally as florencites. Those in PF2–1415 are dominantly interpreted as florencites.</figcaption></figure>

which rather shows Th enrichment in the whole saprolite and not specifically in the two clayey levels (Fig. S3). Consequently, the interpretation of the dating results for the two clayey levels PF2_1345 and PF2_1415 in the saprolite remains open.

The average age proposed for the kaolinites trapped in the Fe duricrust, assuming that they behave like saprolite kaolinites, is older (ca. 22 Ma) than the embedding hematite (16 Ma), which is meaningful because once the duricrust is formed, it is not expected that secondary clays can form within this compact horizon except if porosity remains open and connected. In addition, the kaolinites from the nodules are much older (ca. 23 Ma) than the hematite of the nodules (6 Ma). These kaolinites may belong to the general trend of rejuvenation as their age is a maximum age, owing to their specific linear dosimetry behavior.

<figure><img src="image_16.png" /><figcaption>**Fig. 13.** Ages of secondary minerals in the profile. With Q: Quaternary; P: Pliocene. Ages of kaolinites are derived from experimentally determined paleodose (black points) or paleodose modeled using dosimetry parameters of similar samples (grey points). The age of well-ordered kaolinite from duricrust is calculated with a paleodose determined by an exp. +lin. Fit (Eq. (4)) similar to well-ordered kaolinite from the saprolite while for disordered kaolinite (Ferralsol) an exp. fit was used. The age of nodule kaolinites (sample PF2_750) is a maximum age, as its dosimetry curve is linear. Red squares stand for ages of hematite of duricrust and nodules. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

# 5. Discussion

## 5.1. Origin of the lateritic cover

The origin of the thick Ferralsol (also called “Belterra Clay”) covering the Amazonas sedimentary basin as observed at the top of the Presidente Figueiredo profile, has been strongly debated (e.g., Sombroek, 1966; Lucas et al., 1984; Chauvel et al., 1987; Truckenbrodt et al., 1991; Horbe and da Costa, 2005; Balan et al., 2007), and remains an open question. We acknowledge that significant work is required to pinpoint the origin of this kaolinite-rich Ferralsol and to understand its occurrence and persistence over long periods of time, but geochemical, mineralogical and geochronological data provided in this study can shed light on the its origin and genetic link to the Alter do Chão Formation. A first feature was observed regarding the mineralogy of the lateritic profile, because it was formerly predicted a progressive desilication in topsoil with increasing content of gibbsite over kaolinite, as expected from thermodynamic considerations (Tardy and Roquin, 1992). However, in the Manaus region, kaolinite remains predominant in the Ferralsol, a preservation that would require Si saturation with respect to kaolinite in the pore waters from which this layer was formed. Such an evolution was interpreted by Lucas et al. (1993) as arising from biological inputs via litter degradation which would lead to Si inputs to the percolating fluid, which in turn could explain the lower δ³⁰Si values (Poitrasson, 2017) in the near-surface Ferralsol as observed in the PF2 profile (Fig. 6C). This
process derived from recycling of Si by plants is also consistent with the rejuvenation of kaolinite to the expanse of an older generation, as observed in our study.

Sombroek (1966) first proposed a sedimentary origin for the "Belterra clay", with prior formation of kaolinite in the Andes region, and later transport and deposition in lacustrine systems of the Amazon region after the Andean uplift. Considering from our study that topsoil kaolinites were formed at around 10 Ma, this scenario implies that the Alter do Chão sediment and the "Belterra clay" have different sources (surrounding Amazonian craton for the former and from Andean Cordillera for the latter). Such distinct sources could be the result of the switch in the direction of the paleo-Amazon river flow in the late Miocene (10–9 Ma) (Figueiredo et al., 2010; Hoorn et al., 2010; Jaramillo et al., 2017; van Soelen et al., 2017), but several observations argue against this idea. First, Irion et al. (1995), concluded that a sedimentary origin of the "Belterra clay" is not likely because it should result in deposition structures (as in the saprolite over the Alter do Chão formation), together with thickness and mineral compositions independent from the parent material, which they did not observe in more than 20 profiles from the Central Amazonia. Second, Balan et al. (2005) interpreted the Ferralsol formation in a lateritic profile near Manaus, 60 km south of the present studied area, as a progressive rejuvenation of kaolinite toward the surface, being the consequence of a complex in situ geochemical reorganization due to weathering. The authors concluded that the kaolinites of this Ferralsol, which display similar mineralogical features and EPR ages to those of kaolinite from the present study have been precipitated at an early stage of the Ferralsol formation and preserved. In addition, the invariant Ti/Nb ratios measured within the studied profile (saprolite and Ferralsol), and in profiles developed regionally over the Alter do Chão formation in Balan et al. (2005) and Gautheron et al. (2022) indicate the absence of variation between the parental material and the Ferralsol.

Two additional genetic relationships can be questioned to address the present issue, the link between the Ferralsol and the iron-rich levels (duricrust/nodules) and the link between the nodules and the duricrust. Even though Fe-duricrusts are not observed in all lateritic profiles of the region, other theories are inherently related to a link between the duricrust and the Ferralsol and intend to understand how such thick soil layers could be formed above a duricrust. Indeed, a common idea of lateritic profile formation indicates that a duricrust is formed at the top due to seasonal variation of the water table leading to fluctuating redox conditions, and with the downward progression of the weathering front (Tardy, 1993). However, in the present profile, the duricrust is found more than 8 m below the surface and is overlaid by the thick Ferralsol. Horbe and da Costa (2005) noted that the REE distribution from the duricrust to the Ferralsol (including nodules), is relatively invariant, which supports the idea of a common source for all these horizons. These authors interpreted the Ferralsol as the weathering product of a former duricrust, with kaolinite remaining in the profile while Fe-(oxyhydr) oxides having been completely leached. However, a thickness of up to 20 m of Ferralsol as encountered in the sedimentary Amazonas basin (Truckenbrodt et al., 1991) would imply the weathering of several tens of meters of duricrust considering its low amount of kaolinite (Fig. 4C), and without any geochemical reworking of Fe-(oxyhydr)oxides as suggested by our (U—Th)/He dating data. A reworking of the Ferralsol was also attributed to the termites activity, but the 20 m thickness that can reach this soil and its widespread occurrence over large areas are hardly explained by such process (Truckenbrodt et al., 1991 and references therein). Additionally, the degree of disorder and concentration of structural iron in kaolinite from the duricrust and the Ferralsol (Fig. 6D) is strongly different, which indicates distinct conditions of formation and not direct inheritance.

Nodules found in the Ferralsol have also been interpreted in several studies (e.g., Journaux, 1975; Irion et al., 1995; Horbe and da Costa, 2005) as evidence of the presence of an erosion surface of a duricrust under arid climate, and supposedly followed by the "Belterra clay"

deposition. This interpretation is not supported by our data because the nodules are embedded in, and not below, the Ferralsol material, strengthening the scenario of an *in situ* formation of nodules within the soil. Additionally, ages obtained in the present study for hematite in nodules differ from the ones of the lateritic duricrust by up to 10 Ma and reveal that nodules cannot simply be an erosion product of the duricrust, but rather have a distinct formation history as also attested by distinct Th/U ratios (Fig. 9, Table S7) (Monteiro et al., 2014). Furthermore, kaolinites trapped in the duricrust and the nodules also show contrasting crystal chemistry, which can only be acquired through dissolution/reprecipitation processes and not direct inheritance between the ferruginous materials. The discussion of Lucas et al. (1989) support this idea by reporting a significant difference in the dissolution kinetics of hematite and gibbsite nodules embedded at the base of the Ferralsol. Their observation implies that gibbsite nodules are formed at the expense of pre-existing ferruginous layers while the weathering front propagates downward. This means that the deposition of the Ferralsol layer on top of a pre-existing duricrust is unlikely in this area, or that further geochemical reorganization has completely erased any trace of the allochthonous origin of the Ferralsol.

The fact that the Ferralsol is derived from the weathering of the Alter do Chão formation, just as the saprolite below the duricrust, while also being of detrital nature, could be reconciled if the Ferralsol had been deposited on top of the duricrust after its formation, through sediment transport, triggered e.g. by mass wasting processes from the surrounding relief. Although this scenario is challenged by the lack of any known significant tectonic activity in the area, we cannot strictly rule out this possibility. Regardless, the following discussion on the paleoenvironmental and paleoclimatic evolution of the Amazonas basin does not rely on the knowledge of the exact formation mechanisms of the Ferralsol layer.

Overall, our data tend to support the *in situ* formation of the Ferralsol by the progressive evolution of the EPR spectra of the fine fraction showing a progressive upward decrease in kaolinite crystalline order and $\delta^{30}$Si signature (Fig. 6F, 8), a marker of chemical weathering (Opfergelt et al., 2012; Baronas et al., 2020; Guoiseau et al., 2021), suggesting a genetic connection between the Alter do Chão Formation saprolite, the duricrust and the Ferralsol.

## 5.2. Weathering episodes recorded at Presidente Figueiredo

EPR ages of kaolinites from the PF2 profile are in line with previously studied profiles around Manaus (Balan et al., 2005) confirming the record of at least two major weathering events around the Oligocene and the end of Miocene (Fig. 13). Specific weathering conditions are revealed in the two clayey levels in the saprolite (PF2_1415 and PF2_1345). Episodes of Fe mobilization are also dated ca. mid-Miocene and Late Miocene (Fig. 13), and suggest changes in paleoenvironmental conditions at these periods of time.

### 5.2.1. From the Alter do Chão sediment to a saprolite

The first weathering episode recorded in the studied lateritic profile is attributed to the formation during the Oligocene of sandy-clayey saprolite and pure kaolinitic intercalated beds (Balan et al., 2005), from the weathering of the Alter do Chão Formation (Fig. 13). Indeed, kaolinized feldspars and large kaolinite booklets inconsistent with any transport (Fig. 3D) show that the saprolite results from an *in-situ* weathering of fluvio-lacustrine sediments deposited during the Late Cretaceous-Early Paleogene period (Irion, 1984; Horbe et al., 2006; Mendes et al., 2012).

In bulk samples, the evolution of the Al/Si weight ratio, a tracer of grain size (Bouchez et al., 2011), and the mass-transfer coefficients of elements mainly carried by florencite in the Alter do Chão Formation such as La, Ce or Pb, show that all samples in the whole profile are derived from sandy-clayey horizons, with pure kaolinitic beds falling off the distinct trend (Fig. 14). Considering the whole Alter do Chão
<figure><img src="image_17.png" /><figcaption>**Fig. 14.** Mass-transfer coefficient of La vs. the Al/Si ratio in the samples of the lateritic profile at Presidente Figueiredo (PF2), revealing the geochemical continuity between the sandy-rich saprolite samples and the lateritic cover. The two outliers from the saprolite are from kaolinite-rich beds (samples PF2_1415 and PF2_1345).</figcaption></figure>

saprolite, conditions for kaolinite formation appear rather invariant based on their homogeneity in terms of crystalline order, δ¹⁸O, δD and δ³⁰Si signatures (Fig. 6). In particular, high degree of order indicates relatively long kinetics of formation, according to what is observed in synthesis experiments (Fialips et al., 1999). As shown by Guinoiseau et al. (2021), large and well-crystallized kaolinite grains with Si isotope composition close to that of the UCC (Fig. 6) likely results from a long weathering episode, i.e., ~15 Myr dated from Oligocene to Miocene, and of moderate intensity. During this episode, kaolinite had time to crystallize as large booklets and to record isotope fractionation of thermodynamic (rather than kinetic) origin (Oelze et al., 2014). The homogeneous δ¹⁸O (21 ± 1 ‰) and δD (−58 ± 3 ‰) values of kaolinites hosted in Alter do Chão Formation suggest that the recorded climate was relatively stable during their pluri-millennial period of formation.

However, the two kaolinite-rich clayey levels (i.e., PF2_1345 and PF2_1415) do not fit within this reconstruction due to their younger formation age (<4 Ma). Even if their crystal chemistry (low crystalline disorder and low concentration of structural Fe) is similar to those found in the surrounding saprolite material, these two levels are also anomalous in terms of ages, mineralogy (i.e., almost pure kaolinite), geochemistry (trace elements including REE), dosimetry (steep growing curves, Fig. S6) and crystal chemistry (EPR spectrum of structural Fe³⁺). The possibility of a strong and recent accumulation of U and/or Th to explain the elevated dose rate in these clayey levels can be discarded because the τ<sub>Th,Nb</sub> are not anomalous compared to saprolite layers (Fig. S3). An alternative and more likely scenario is a water channel circulation or favored retention of water in the microporosity of these kaolinite-rich levels, with low permeability and long residence time, in comparison to the other saprolite samples that belong to more sandy levels. This would lead to the dissolution of pre-existing kaolinites, saturation of weathering fluids and subsequent slow precipitation of younger kaolinites with the same low structural Fe-content, low degree of kaolinite disorder and an unfractionated δ³⁰Si, δ¹⁸O, and δD signatures under a close system, compared with the other saprolite samples. The local environment, and not only climatic parameters, thus appears as an important factor that defines the identity of kaolinite.

### 5.2.2. Lateritic profile evolution through duricrust formation and kaolinite rejuvenation

The second weathering episode affecting the Presidente Figueiredo profile has already been described in the region (Balan et al., 2005) and is characterized by the replacement of the first generation of kaolinites by a generation of smaller, poorly-crystallized, and isotopically lighter (for Si, O and H isotopes) kaolinites (Fig. 6E, F). EPR dating of this

weathering episode at 10 Ma (at the top of the Presidente Figueiredo profile) is consistent with previous data available in the region (Balan et al., 2005), suggesting the regional extent of this rejuvenation process. In addition, the lightest δ³⁰Si values (≈ −2.0 ‰) occurring in the Ferralsol can be explained by kinetic isotope fractionation obtained during fast kaolinite precipitation rather than from a progressive replacement process over tens of million years (Guinoiseau et al., 2021). High degree of disorder is also consistent with fast kinetics of crystallization. In addition, reactive transport modelling of this weathering episode coupled with kaolinite EPR ages determined in lateritic covers of the region suggests a short episode (< 2 Myr), characterized by faster drainage in unsaturated material compared to the first episode (Balan et al., 2005; Guinoiseau et al., 2021). This observation also agrees with the micro-aggregated structure of Ferralsols. However, while the evolution of δ³⁰Si signatures in the lateritic profile studied by Guinoiseau et al. (2021) and of the δD – δ¹⁸O in the Presidente Figueiredo profile likely result from mixing between two populations of kaolinite (Fig. 6E, 8), intermediate δ³⁰Si signatures recorded here at middle depths (7–10 m) are not consistent with two endmember mixing (Fig. 6F, 8). Indeed, expected δ³⁰Si values of kaolinite have been recalculated (yellow dots in Fig. 8) by assuming binary mixing between a well-ordered and a disordered kaolinite, and using the apparent δ³⁰Si signatures of each endmember (−2.00 ‰ value for poorly crystallized kaolinite and, 0.03 ‰ for well-crystallized kaolinite, that is the average δ³⁰Si value of the saprolite clay-sized fraction). The relative contribution of each kaolinite population was determined by linear decomposition of the EPR spectra. The measured δ³⁰Si signatures in the close vicinity of the duricrust, i.e., in the Ferralsol and in the mottled zone, are heavier than those expected (Fig. 8). We contend that this slightly heavier Si isotope composition is related to the formation of the iron duricrust (dated at maximum ca. 18 Ma) before the occurrence of the second episode of kaolinite formation (dated at ~10 Ma). The duricrust may act as a partial hydrological barrier that could slow down the downward progression of the weathering front, thus lowering the replacement of ordered kaolinites by disordered ones and eventually weakening the kinetic effect of Si isotope fractionation. Unlike Si isotope signatures, uniform δ¹⁸O and δD values of kaolinites formed during this second weathering episode show that these isotope systems are not directly impacted by the dynamics of kaolinite formation. This is because O and H isotope ratios mostly depend on temperature and the signature of the water mass from which kaolinites precipitate, rather than on the kinetics of kaolinite precipitation (Savin and Hsieh, 1998).

The formation of the duricrust in the porosity of the clayey matrix, implies a change in physical and chemical conditions prevailing along the lateritic profile. The formation age of the duricrust, determined at least at 18 ± 2 Ma by (U–Th)/He, is younger than the chemical weathering of the Alter do Chão saprolite during the Oligocene, but older than the formation of the Ferralsol (Late Miocene). This chronology is supported by preserved well crystallized kaolinites embedded in the dated Fe-duricrust, dated at 22.2 ± 6.7 Ma, and characterized with large booklets morphology, similar to the one found in the saprolite. The formation of iron duricrusts needs tropical climate with strong seasonality enabling the dissolution and precipitation of Fe-rich layers. Two hypothesis of formation are regularly put forward in the literature: controlled by water table fluctuations or controlled by lateritization (e. g., Herbillion and Nahon, 1988; Widdowson, 2007). The former involves the fluctuation of a water table level in tectonically quiescent regions to allow duricrust progressive formation. In details, lateral groundwater transfers of dissolved Fe²⁺ during wet periods (reducing conditions) are followed by the precipitation of Fe-(oxyhydr)oxides mostly driven by the water table fall during the dry season (Braun et al., 1990; Ollier and Galloway, 1990; Beauvais and Colin, 1993; Gilkes et al., 2003). The lateritization-based process is considered in the literature as the ultimate stage of lateritization, with a progressive enrichment of insoluble Fe-(oxyhydr)oxides following saprolite chemical weathering (Widdowson, 2007). A dominant vertical water flow is causing a progressive
desilication of the profile, which ended with the cap of the weathering profile by an insoluble Fe-duricrust. To avoid the production and the systematic destruction of this duricrust, a tectonic uplift, even limited is needed. In the present study context, the groundwater-based model is the most likely because it can occur in tectonically inactive regions, explaining why such a duricrust has been formed and preserved at 9 m depth below a siliciclastic regolith. Significant Fe-like elements enrichments in this duricrust level in comparison with the saprolite also support this idea (Fig. S2C). This scenario also stands up for the episodic occurrence of duricrusts in the region. Thus, the duricrust formation should be seen locally as a relic of a former water table level and associated with an external input of elements through lateral advection, without any genetic link with the sediment constituting the PF2 profile. Therefore, the duricrust is evidencing a third weathering episode and a clear disconnection between weathering episodes affecting silicate and Fe-(oxyhydr)oxides in this lateritic profile.

The nodular zone above the duricrust is the second iron-rich level of the Presidente Figueiredo profile. We have concluded in section 5.1 that these nodules have likely been formed *in situ* and are not genetically connected to the duricrust. The hematite precipitation age (~6 ± 1 Ma) is slightly younger than kaolinite formation in the surrounding Ferralsol (~10 Ma). As advanced for the duricrust, the presence of these nodules can likely be the result of a paleowater table, and associated advection of Fe lateral fluxes filling the porosity of the Ferralsol. The lower induration of the nodule-rich level can result presumably from a shorter and less significant event than for the duricrust formation. Because of their rare occurrence in the region except in the profiles close to Manaus (Lucas et al., 1996; profile III from Balan et al., 2005), and of their reduced expression in the Presidente Figueiredo profile (only 30 cm thick), we believe that these nodules reflect local and transient conditions of formation rather than a regional change of climate.

Accordingly, a decoupling between the duricrust formation and the kaolinite formation events occurred at Presidente Figuereido, both temporally and in terms of physicochemical conditions, challenging the historical view of lateritic profiles as active weathering covers formed from progressive, vertical, and contemporaneous weathering fronts (Nahon and Tardy, 1992; Tardy, 1993; Lucas et al., 1996; Nahon, 2003). The clear shift in δ¹⁸O~Kw~ and δD~Kw~ values of parent water between the two episodes of kaolinite formation also highlights a drastic change of the water mass signatures infiltrating into the lateritic profile, which can be interpreted in terms of major climatic and environmental modifications over the Cenozoic in the studied area.

## 5.3. Constraints on paleoenvironmental conditions during the Cenozoic in the Central Amazon basin

### 5.3.1. Climate control on the δ¹⁸O and δD signatures of secondary minerals

The δD~K~ and δ¹⁸O~K~ values in the PF2 profile (Fig. 15) plot on a single line parallel to the Kaolinite Weathering Line (KWL) defined by Sheppard et al. (1969). This indicates that kaolinites in the PF2 profile have been formed at a roughly constant temperature, in equilibrium with paleo-meteoric waters under supergene conditions, and that they have preserved the original isotopic signature of paleo-rainwaters (Savin and Epstein, 1970). Interestingly, although δD~K~ and δ¹⁸O~K~ values are homogenous within each of the two defined populations of kaolinites, clear differences can be observed between old and young kaolinites (Fig. 15). The homogenous δD~K~ – δ¹⁸O~K~ signatures for each kaolinite population of Presidente Figueiredo profile suggests that over their respective long periods of formation (i.e., Myr-scale for both populations), kaolinites integrate long-period variability (i.e., over kyr or Myr) in rainwater isotope composition and thus in climatic trends. As kaolinite is expected to form during humid period with Al-Si-rich fluid circulation (Fritz and Tardy, 1973; Del Nero and Fritz, 1990), we calculated the composition of paleo-rainwater in δD and δ¹⁸O (see section 4.3.1) using the δ¹⁸O~K~ – δD~K~ signature (Table 1) and the δ¹⁸O~Fe~ values. We compared these calculated δ¹⁸O~Kw~, δ¹⁸O~FeW~ and δD~Kw~ values (Fig. 7) with the

<figure><img src="image_18.png" /><figcaption>Fig. 15. Relationship between the δD and the δ18O values of kaolinite in the Presidente Figueiredo profile. Kaolinite isotherms (dotted line) were drawn using the equation of Sheppard and Gilg (1996) and considering the present-day global meteoric water line equation δD = 8 δ¹⁸O + 10. The Kaolinite Weathering Line (KWL) represents the isotopic value expected for kaolinite of supergene weathering origin (δD = 7.6 δ¹⁸O – 220; Sheppard et al., 1969).</figcaption></figure>

composition of modern rainwater collected in the Manaus area during the wet season, which represents on average 90 % of the annual rainfall (δ¹⁸O~w.Manaus~ = -5.2 ± 1.4 ‰ and δD~w.Manaus~ = -29.3 ± 11 ‰; Table S4). However, it is difficult to compare these two sets of data considering the significant inter-annual variability of present-day δ¹⁸O~w.Manaus~ – δD~w.Manaus~ (e.g., δ¹⁸O~w.Manaus~ = -2.70 ‰ and δD~w.Manaus~ = -9.50 ‰ in 1983 while δ¹⁸O~w.Manaus~ = -6.81 ‰ and δD~w.Manaus~ = -40.3 ‰ in 1976).

The calculated composition of the paleorainwater for the second weathering episode (Late Miocene) is statistically lighter (δ¹⁸O~Kw.~ = -5.7 ± 0.7 ‰ and δD~Kw.~ = -35.6 ± 5.4 ‰) than the rainwater composition inferred to the first episode (Late Eocene - Oligocene) with δ¹⁸O~Kw.~ = -4.3 ± 0.5 ‰ and δD~Kw.~ = -24.9 ± 4.3 ‰. Hematite from the duricrust PF2_870 have a δ¹⁸O~FeW~ similar to that of the kaolinite from Ferralsol (ca. -6.70 ‰, Fig. 7), which is consistent with their similar (U—Th)/He and EPR formation age (Fig. 13). Iron nodules from samples PF2_750, which is younger than kaolinite from the Ferralsol and the Fe-oxides from the duricrust, show even lighter parent water composition and indicate a change in the water source composition, similarly to kaolinite atop of the profile (Fig. 7). The observed δ¹⁸O and δD differences between weathering episodes are thus likely to reflect long-term changes in the water flow through the profile (i.e., rainfall quantity and intensity, infiltration rate), which can in turn provide constraints on climatic conditions (e.g., temperature, air moisture) under which these kaolinites were formed (Girard et al., 2000; Feng and Yapp, 2009; Clauer et al., 2015; Fang et al., 2019).

While it is often assumed that temperature is the main driver of the isotopic composition of pedogenic secondary minerals (Yapp, 1987; Savin and Hsieh, 1998), additional factors, like atmospheric circulation or vegetation cover, are known to affect the isotopic composition of
precipitation and are important to discuss prior to interpreting isotope data in terms of climate evolution.

A direct paleo-geographical effect is unlikely because the studied area has stayed in the same equatorial position for almost 100 Ma (Sheppard et al., 1969; Siegenthaler, 1979; Tardy et al., 1991), discarding any potential latitudinal effects on paleo-rainwater δD and δ¹⁸O values. An altitudinal effect on δD and δ¹⁸O (−1.5 to −2.5 ‰/km for δ¹⁸O, Gonfiantini et al. 2001) could be linked with the uplift history of the Central Amazon basin. Major uplift and physical erosion of the Guiana Shield has been recorded during the opening of the equatorial Atlantic Ocean at the end of Cretaceous (60–80 Ma), based on sediment accumulation in the Guiana-Suriname basin (Rouby et al., 2023) apatite fission tracks (Harman et al., 1998), Sr–Nd isotope signatures (Roddaz et al., 2021), or on low denudation rate reconstruction of the Guiana Shield (Derycke et al., 2021). This event can explain the deposition of the Alter do Chão sediment but precedes the main weathering episodes recorded in the PF2 profile. From 60 Ma to Quaternary, both denudation rates and siliciclastic accumulation in the Guiana-Suriname basin have been considerably reduced, arguing in favour of limited uplift and erosion in the Central Amazon region (Rouby et al., 2023). The tectonic pulses during the Andean orogeny has certainly caused minimal uplift in the Central Amazon (Hoorn et al., 1995; Garzione et al., 2008; Sundell et al., 2019), especially after the onset of the transcontinental Amazon River (8–11 Ma). However, we contend that these altitudinal modifications are not strong enough to be a significant driver of the shift observed in the δD – δ¹⁸O signatures of water flowing 8 m depth. The consequences of the Andean orogeny are most likely rather indirect, affecting the atmospheric circulation and precipitation distribution throughout South America as proposed by sedimentary investigations and atmospheric modelling (Lamb and Davis, 2003; Insel et al., 2010; Sepulchre et al., 2011). An effect of a change in the origin of atmospheric moisture can also be discarded as atmospheric circulation models reveal that water moisture above the Amazon basin should have remained of Atlantic origin with or without the Andean orographic barrier (Insel et al., 2012; Jeffery et al., 2012; Pöhlker et al., 2019). Finally, the extent of the Amazonian rainforest cover might have impacted the regional hydroclimate via enhanced or reduced evapotranspiration and recycled rainfall, and therefore should have affected rainwater isotopic composition (Matsui et al., 1983; Ampuero et al., 2020; Costa et al., 2021). But palynological studies showed that the Amazon rainforest have persisted in this area since at least the beginning of the Cenozoic (Hoorn, 1994; Colinvaux and De Oliveira, 2001; Jaramillo et al., 2011), even though changes in vegetation diversity could have occurred during regional cooling periods (Hoorn et al., 2010, 2023). Consequently, the δD and δ¹⁸O compositions of investigated secondary minerals is most likely the results at the regional-scale of climatic perturbations in temperature, or water source and recycling.

The effect of temperature on secondary mineral formation can be discussed by reconstructing the paleo-temperatures of kaolinite crystallization using the δD – δ¹⁸O signatures, considering the kaolinite-water fractionation factor from Sheppard and Gilg (1996) and assuming that the water composition at the time of kaolinite formation followed the present-day Global Meteoric Water Line (GMWL; δD = 8 x δ¹⁸O + 10) (Clauer et al., 2015):

$$3.04 \times 10^6 \times T^{-2} = \delta^{18}O_k - 0.125 \times \delta D_k + 7.04 \quad (6)$$

Calculated temperatures range from 19 °C to 27 °C (Table 1), as expected for the tropical conditions (20 °C – 30 °C) that have prevailed since ca. 100 Ma at the sampling site (Tardy et al., 1991; Zhang et al., 2019). A homogenous crystallization temperature (T<sub>average</sub> = 23.4 °C; 1σ = 2.3 °C) has been obtained across the whole Presidente Figueiredo profile despite significant differences in kaolinite crystalline order and δ³⁰Si, δ¹⁸O and δD compositions. This observation implies that the fluid composition and crystallization kinetics, rather than temperature are the main driver of the observed changes in kaolinite crystalline order and

isotopic composition during the formation of the lateritic cover. Such past variations in δ¹⁸O and δD composition of the weathering fluid are more likely due to a change in local or regional hydrological regime (water saturation in the profile, temporal rainfall distribution) through the Cenozoic (Zhang et al., 2019; Scotese et al., 2021).

### 5.3.2. Paleoclimatic and paleoenvironmental implications for the Cenozoic Central Amazon basin

The minimal influence of global climatic shifts on low-latitude regions, combined with the scarcity of archival records, has led to poor documentation of paleoenvironmental conditions in the Amazon basin during the Paleogene. Recent attempts have aimed to provide a better understanding of the Cenozoic history of the Amazon basin using palaeobotanical, paleofauna, sedimentary or pedological records (Hoorn, 1994; Colinvaux and De Oliveira, 2001; Jaramillo et al., 2006; Klaus et al., 2017; Boivin et al., 2019; Gautheron et al., 2022; Salazar-Jaramillo et al., 2022; Hoorn et al., 2023). Our new mineralogical and isotope data on the Presidente Figueiredo lateritic profile contribute to a refinement of these paleo-environmental scenarios for the Central Amazon basin.

After the deposition of the Alter do Chão Formation, the first weathering episode recorded in the Presidente Figueiredo profile is attributed to the formation of well-ordered kaolinite in the Oligocene. In the Guiana Shield, the Paleogene has been documented as a period of bauxite formation (Bardossy and Aleva, 1990; Wong, 1994; Ansart et al., 2022), associated with intense weathering and very humid climatic conditions in well-drained environments. However, the formation of well-ordered kaolinite in the Presidente Figueiredo profile is at odds with such a scenario (Tardy and Nahon, 1985). The formation of large and well-ordered kaolinite booklets rather indicates slow crystallization under stable chemical conditions. Kaolinite crystalline order and oxygen and hydrogen isotopic signatures are consistent with a poorly contrasted climate indicating a weak seasonal variability of rainfall. Even if most of the kaolinite ages recorded in the PF2 profile point to 20–30 Ma, the onset of this first weathering episode cannot be dated precisely. Indeed, older kaolinites formed by weathering of the Alter do Chão sediments have been found by Balan et al. (2005) in other lateritic profiles of the Manaus region. This means that some records of ancient weathering remain, albeit scarce. This is also the case from a few iron duricrusts around 40 Ma developed on the Alter do Chão formation and dated by (U–Th)/He (Gautheron et al., 2022). Such relatively old Paleogene records are so far specifically associated to the weathering of the Alter do Chão formation. Indeed, this weathering episode is not encountered in the profiles developed on granite from the Western Central Amazon basin, where the Neogene rejuvenation of kaolinites was more complete (Mathian et al., 2020). Generally, this weathering episode aligns well with a shift from an erosion-dominated regime in the Late Cretaceous to a weathering-based regime during the Paleogene (Rouby et al., 2023). At a global scale the Oligocene is marked by the opening of the Drake passage, ca. 34 Ma, leading to the formation of the circumpolar Antarctic current, and allowing for the build-up of a permanent Antarctic ice sheet (Sijp and England, 2004; Livermore et al., 2005). Aside from inducing a global decrease in sea level, this event is suspected to have caused an episode of a less humid environment in western Amazonia, which initiated taxonomic fauna diversification (e.g., Boivin et al., 2019), a drop in flora diversity (Jaramillo et al., 2006; Hoorn et al., 2010), and a decrease in continental drainage (Tardy and Roquin, 1998).

The relative drying of western Amazonia during this period was followed by an episode of sea level rise in the mid-Miocene (de Rossetti, 2001; Miller et al., 2020) initiating the formation of the Pebas system, an large region of wetlands extending from the present Orinoco basin to the western Amazonia (Antoine et al., 2006; Wesselingh, 2006; Hoorn et al., 2010; Leite et al., 2017). The formation of the duricrust of the PF2 profile during the mid-Miocene (~15 ± 2 Ma), is contemporaneous with the Pebas formation, and thus may potentially be linked with the same climatic event. These events coincide in time with the Middle Miocene
Climatic Optimum (MMCO), characterized by an increase in global temperature (Miller et al., 2020), and are synchronous with an enhancement in continental drainage according to Tardy and Roquin (1998). According to Scholz et al. (2020), in Columbia, this episode is synchronous to an increase of precipitation associated to a higher seasonality and was proposed to arise from a northward shift of the Intertropical Convergence Zone. This timing also corresponds to the deposition of the Miocene Novo Romanso Formation in the central part of the sedimentary basin (Dino et al., 2012), linked with a higher water base level in the Central Amazon basin. As discussed in section 5.2, we have seen that the duricrust formation likely results from lateral Fe²⁺ fluxes during wet seasons and subsequent precipitation in the form of Fe-(oxyhydr)oxides, when the water table fluctuates during dry season. The Alter do Chão Formation is known as one of the most important aquifers of Brazil (Souza et al., 2013; Henrichs De Azevedo and Guimaraes Campos, 2021) and thus could represent appropriate environment to allow for water circulation at high piezometric elevation during the mid-Miocene and for Fe-immobilization as a duricrust (Tardy and Nahon, 1985). The homogenous ages (~15 ± 2 Ma Ma) determined in the Presidente Figueiredo duricrust constitute evidences for the

preservation of these environmental conditions over Myr times scales (Fig. 16). Similar ages for the formation of the lateritic duricrust (Late Oligocene – Early Miocene) in Central Amazonia (Gautheron et al., 2022) could indicate a regional extent of this episode. In addition, the preservation of the ferruginous lateritic duricrust of Presidente Figueiredo weathering profile until today shows that conditions enabling its total dissolution have not been encountered for long enough since its formation.

Kaolinite found atop the profile, associated with a late Miocene weathering episode, have δD – δ¹⁸O values consistent with previous studies performed across the Amazon basin and the Guiana Shield regarding the same period of time (Bird et al., 1993; Giral-Kacmarcík et al., 1998; Girard et al., 2000; Horbe, 2011). This observation implies regional homogeneity in hydrologic conditions associated with this weathering episode. The timing of kaolinites formation follows a period of significant geodynamic and climatic perturbations with a major Andean uplift starting ca. 15 Ma in the central part of the Andes and progressing toward Northern Cordillera at 5 Ma (Eude et al., 2015; Perez et al., 2016; Sundell et al., 2019). This continental-scale event is known to have changed the atmospheric circulation in South America and to

<figure><img src="image_19.png" /><figcaption>**Fig. 16.** Compilation of climate data (Westerhold et al., 2020), hydrologic and geologic events in South America (Hoorn et al., 2010; Boschman, 2021) and available geochronological data on secondary minerals compared to the new data obtained in this study. Note that mixing EPR ages are interpreted as resulting from the mixture of the two identified kaolinite generations. With EECO: Early Eocene Climate Optimum; MECO: Middle Eocene Climate Optimum; TEE: Terminal Eocene Event; MMCO: Middle Miocene Climate Optimum. Geodynamic events refer to periods of intensified uplift in Central Andes (C.A.) or in Northern Andes (N.A.).</figcaption></figure>
have initiated the onset of the transcontinental Amazon drainage system ca. 8–11 Ma (Hoorn et al., 2017). At the same time, ca. 13 Ma, the formation of a perennial Antarctic ice-sheet in the southern hemisphere drove the shift of the Intertropical Convergence Zone (ITCZ) northward (Holbourn et al., 2010). The light δ¹⁸O and δD composition of kaolinite in the Ferralsol could be representative of this climatic change, under which intense and fast drainage conditions have led to the formation of poorly crystallized kaolinite, for example during a climate with a short but marked wet seasons. Indeed, secondary minerals with light δD – δ¹⁸O signatures in Amazonia have been interpreted as the result of crystallization under highly contrasted climate, e.g., monsoonal climate (Bird et al., 1993; Girard et al., 2000; Kaandorp et al., 2005), which could be responsible for the depleted Late Miocene kaolinite oxygen and hydrogen isotopic composition in this study.

The modern O and H isotopic composition of rainwater obtained in the region of Manaus (Table S4) cannot yield reliable information on long-term climatic trends in the Central Amazon basin, considering the high temporal variability of the available data over the recording period. However, modern drainage conditions are suspected to be different from those prevailing in the Late Miocene weathering episode. As an example of these distinct environmental conditions, more frequent occurrence of young Fe-(oxyhydr)oxides ages found in the Central Amazon basin (Fig. 16) compared to old generations could be the result of weathering of an old population of secondary minerals and recrystallization of a younger ones which could explain why old records are scarce in tropical areas (Ansart et al., 2022; Heller et al., 2022). This implies that conditions of old secondary mineral preservation were not encountered and multiple changes in paleoclimatic and/or paleoenvironmental conditions occurred since the beginning of lateritization in this area.

## 6. Conclusion

Here, a unique combination of mineralogical, geochemical and geochronological investigations has been deployed to reconstruct the weathering history and the associated environmental and climatic conditions prevailing during the formation of a lateritic profile across the Cenozoic in the Central Amazon basin. The extensive constraints on kaolinites and hematite (EPR and U-Th/He ages, respectively) along with the reliability of the methodology yield a robust chronological framework for the entire lateritic profile. The comparison of these ages and of the weathering proxies with data obtained in other profiles of the region validates the regional extent of the weathering processes. Although the Amazon basin has endured tropical climate conditions with relatively constant temperatures for more than 100 Myrs, the three main weathering episodes highlighted in this study, over 40 Myrs in the Central Amazon basin, present strong variabilities in their weathering intensity, duration, water sources and thus in the associated climatic and environmental conditions. Interestingly, the homogenous O and H isotope signatures together with the contrasting ages of each population of kaolinites highlight the relevance of these archives to record long-term climatic trends, providing a perfect complement to biostratigraphic or palynological tracers used to reconstruct the Amazonian evolution through time.

### CRediT authorship contribution statement

**C. Ansart:** Writing – review & editing, Writing – original draft, Visualization, Validation, Methodology, Investigation, Formal analysis, Data curation. **D. Guinoiseau:** Writing – review & editing, Writing – original draft, Visualization, Validation, Methodology, Investigation, Formal analysis, Data curation. **C. Quantin:** Writing – review & editing, Visualization, Validation, Supervision, Methodology, Funding acquisition, Conceptualization. **D. Calmels:** Writing – review & editing, Validation, Supervision, Investigation, Funding acquisition, Conceptualization. **C. Gautheron:** Writing – review & editing, Validation, Supervision, Resources, Project administration, Funding acquisition, Conceptualization.

**P. Agrinier:** Writing – review & editing, Validation, Methodology, Data curation. **J. Bouchez:** Writing – review & editing, Validation, Supervision, Methodology, Investigation, Funding acquisition, Conceptualization. **Z. Fekiacova:** Writing – review & editing, Validation, Supervision, Funding acquisition, Conceptualization. **R. Pinna-Jamme:** Validation, Supervision, Data curation. **D. Demri:** Data curation. **E. Balan:** Validation, Data curation. **B. Heller:** Data curation. **Q. Bollaert:** Data curation. **Ö. Ataytür:** Data curation. **T. Allard:** Writing – review & editing, Writing – original draft, Visualization, Validation, Supervision, Methodology, Investigation, Funding acquisition, Formal analysis, Data curation, Conceptualization.

### Declaration of competing interest

The authors declare the following financial interests/personal relationships which may be considered as potential competing interests.

C. Gautheron reports financial support was provided by French National Research Agency. If there are other authors, they declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

### Acknowledgements

The samples used in this study have been collected in the framework of the RECA ANR-17-CE0101CE-0012-01 project. The authors wish to thank A.M.C. Horbe for field trip organization and her precious comments that improved the original manuscript. The authors would like to thank J. Nouet (GEOPS) for the development of the JaHDe software and his help and advice on the mineralogical quantification. The XRD engineers from GEOPS (S. Miska) and from IMPMC (L. Delbes) laboratories are also warmly thanked for their contribution on this work. We acknowledge A. Somogyi for the access to the NANOSCOPIUM beamline at SOLEIL synchrotron facility (Saint Aubin, France) and the beamline staff for the assistance during the experiments. We also appreciate the support provided by P. Le Pape and M. Chassé from IMPMC for the SOLEIL synchrotron experiment on Nanoscopium beamline. The authors are grateful to F. Genezini for assistance on the neutron reactor at IPEN (Sao Paulo, Brazil). The authors want to thank P. Louvat and P. Burckel for technical support with MC-ICP-MS and ICP-QMS, respectively. This work was supported by IPGP multidisciplinary program PARI, and by Paris region SESAME Grant no. 12015908. This work was financially supported by the French Ministry of Higher Education and Research (C. Ansart PhD grant). The authors thank C. Yapp and two anonymous reviewers for their constructive comments.

## Appendix A. Supplementary data

Supplementary data to this article can be found online at https://doi.org/10.1016/j.palaeo.2024.112674.

### Data availability

The data will be shared on an open-access platform (Pangaea repository) for free access

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