Samples were collected from a northwest-facing hillslope transect extending from the interfluve to the channel of the Hoëgne (LS1–LS6; Fig. 1; Table 1). Current peat depths along this transect (0.2–2.1 m) were previously characterized using ground penetrating radar imaging (Henrion et al., 2024) and are greatest at the footslope (LS5) and least at the topslope (LS2) (Table 1). Samples for determination of saprolite and bedrock geochemistry were collected in one or two soil pits excavated through the peat layer using a shovel at locations LS1–LS4 (Fig. 1B). Larger, quartz-bearing samples for cosmogenic nuclide analyses were later taken at 6 positions (LS1–LS6; Fig. 1B) along the transect by coring through the peat layer using a soil auger. At LS4, the saprolite yielded insufficient quartz, so a bedrock sample taken at the bottom of the soil pit was used in lieu of saprolite. The sample collected adjacent to the Hoëgne (LS6) we interpret as consisting of buried valley-bottom sediment. The thickness of the intervals sampled for quartz ranged between 5 cm (LS4) and 28 cm (LS6). Bulk stream sediment was collected from several locations spaced a few meters apart along the stream bed at the base of the study hillslope and amalgamated to obtain a well-mixed sediment sample (LSR-1). The stream's discharge was monitored between August 2023 and July 2024 using a water level gauge and the local stage-discharge relationship. Stream water samples were collected into acid-cleaned, low-density polyethylene bottles at the gauge at monthly intervals.

Stream water chemistry and saprolite, rock, and stream and valley-bottom sediment geochemistry were measured by inductively coupled plasma-optical emission spectrometry (ICP-OES). Stream solutes are derived partly from atmospheric deposition and partly from bedrock weathering. To isolate the bedrock signal, we focus on Si, which is the major dissolved constituent that is least impacted by atmospheric inputs. Dissolved Si concentrations were measured in the stream water samples using an Agilent 5800 spectrometer. The bulk saprolite samples were wet sieved to isolate the fine earth fraction (< 2 mm). This was heated to 550 °C for 24 h in a glass crucible to remove organic matter and loss on ignition was determined. An aliquot of the ash was weighed and fused in a 2:1 mixture of lithium metaborate and lithium tetraborate at 1000 °C in a graphite crucible and the fusion cake was dissolved in 1.6 M HNO₃. To determine bedrock composition, rocks embedded within the saprolite, the intact bedrock sampled at LS4, and the coarse fraction of the stream sediment (> 2 mm) were crushed in a ball mill, fused, and prepared for measurement in the same way as the saprolite samples. The concentrations of Si and the refractory elements Ti and Zr were determined using a ThermoFisher iCAP 6000 spectrometer.

A watershed-scale denudation rate (D) can be estimated from stream solute fluxes and weathering intensity. To be comparable with denudation rates determined from cosmogenic nuclide concentrations in saprolite samples, we calculate D for the saprolite. To gauge the weathering intensity of Si in the saprolite, a Si-specific chemical depletion factor (CDF_Si) was calculated following Riebe et al. (2004): 1CDFSi=1-Sisap[I]rockSirockIsap where [Si]sap is the concentration of Si in the saprolite, [Si]rock is the concentration in the bedrock from which the saprolite is derived, and [I]sap and [I]rock are the concentrations of a refractory element that is assumed to be immobile during weathering, such as Zr or Ti, in the saprolite and bedrock, respectively. In equilibrium, the denudation rate is proportional to the weathering flux of Si, which is equivalent to the product of specific discharge (Q, mL cm⁻²) and the Si concentration in stream water ([Si]w, µgmL-1), [Si]rock, and CDF_Si (Riebe et al., 2004): 2D=QSiwCDFSi[Si]rock Significantly, Eq. (2) assumes that the dissolved flux of Si is constant over regolith residence timescales. Specific discharge was estimated by taking the mean of the discharge measurements on the 12 sampling days (1 per month), multiplying by the length of the year, and dividing by catchment area.

To isolate quartz for cosmogenic nuclide measurements, the saprolite (LS1–LS3, LS5), bedrock (LS4), valley-bottom sediment (LS6), and stream sediment (LSR-1) samples were first washed and (except LS4) sieved. Vein quartz and quartzite pieces were hand-picked from the > 2 mm fraction of the saprolite and valley bottom sediment because of the small available sample sizes and, along with the rock sample from LS4, crushed and cleaned of meteoric ¹⁰Be by repeated leaching in 2 % HF/1 % HNO₃. Quartz was separated from the 0.25–2 mm fraction of the stream sediment through treatment with H₂O₂ to remove organics, leaching in 2 % HF/1 % HNO₃, and density separation in lithium heteropolytungstate.

The quartz separates were spiked with low-level ⁹Be carrier (1052 µgBeg-1) and dissolved in concentrated HF / HNO₃. After digestion was complete, volatile fluorides were removed by evaporation and the residues dissolved in HNO₃. A liquid aliquot was taken for determination of total sample Al content by ICP-OES using an Agilent 5800 spectrometer. The Be and Al were then isolated and purified using a simplified separation scheme. In brief, non-amphoteric species were first precipitated at pH 14 and removed by centrifugation. The Be and Al were then precipitated at pH 8 and the precipitates dissolved in 0.4 M oxalic acid. The oxalic-acid solution was passed through a 2 mL anion exchange column (Dowex 1x8, 100–200 mesh). In 0.4 M oxalic acid, Be²⁺ forms a neutral oxalate complex and passes through the column, whereas negative oxalate complexes of Al³⁺, Fe³⁺, and Ti⁴⁺ are retained on the resin. The Al was then selectively eluted from the column in 0.4 M HCl. The Be and Al were precipitated as Be(OH)₂ and Al(OH)₃ with NH₄(OH). The hydroxides were washed with 18.2 MΩ-cm water, transferred to quartz glass crucibles, dried, and oxidized at 900 and 1000 °C, respectively, in a muffle furnace. The BeO was mixed with Nb powder and the Al₂O₃ with Cu powder. The mixtures were loaded into Cu cathodes for measurement of ¹⁰Be / ⁹Be and ²⁶Al / ²⁷Al by accelerator mass spectrometry (AMS) at the Laboratory of Ion Beam Physics, ETH Zurich using the compact MILEA system (Gautschi, 2024). The AMS measurements were normalized to the in-house standards S2007N and S2010N (¹⁰Be) and ZAL94N (²⁶Al) (Christl et al., 2013), which have been recalibrated against the KN standard series (Nishiizumi, 2022).

Despite a large water flux from the study catchment, the solute flux carried by the stream is low. The annual specific discharge was 1214 mm yr⁻¹, which is equivalent to 70 % of precipitation (P) during the August 2023–July 2024 gauging period (1736 mm) measured at the Mont Rigi weather station. Given the water balance relationship for a watershed with no change in groundwater storage with time (Q = P - ET), this suggests that ET is 522 mm yr⁻¹. To obtain a longer-term estimate of Q, we subtract this ET value from the 30-year normal MAP of 1440 mm yr⁻¹ (Mormal and Tricot, 2004) giving a long-term specific discharge of 918 mm yr⁻¹. Concentrations of Si measured in the stream water range from 0.84 to 1.59 µgmL-1 and do not vary with the stream discharge on the sampling days: the unweighted mean concentration is 1.20 µgmL-1, which is nearly identical to the flow-weighted mean of 1.22 µgmL-1 (Data Tables in the Supplement; Moore, 2026). The annual Si weathering flux from the study catchment calculated using the long-term specific discharge is 1.12 tons Si km⁻² yr⁻¹, equivalent to 2.39 tons SiO₂ km⁻² yr⁻¹. That the gauging record misses high discharge events is unlikely to appreciably bias this estimate if overland flow becomes significant during these events and the excess discharge does not interact with the soil.

Calculating CDF_Si (Eq. 1) for a unit of saprolite requires knowing the elemental concentrations in the parent material. Although concentrations were determined in rock samples embedded within the saprolite (Data Tables in the Supplement; Moore, 2026), these might not accurately reflect the parent material because the bedrock at the study site consists of interbedded argillite and quartzite, whereas the sampled rocks were exclusively quartzite, which more effectively resists weathering. The ratio between two immobile elements should not change between the parent material and saprolite. However, TiO₂ / Zr ratios in the analyzed rocks are, on average, lower than those in the saprolite (mean and standard deviation of 9.8 ± 3.6 vs. 19.3 ± 4.7), demonstrating that the rocks do not fully capture the parent material composition, assuming that Ti and Zr are equally refractory. Conversely, the coarse fraction of the stream sediment (> 2 mm), which consists of a mixture of argillite and quartzite gravel, has a TiO₂ / Zr ratio of 17.8. This is comparable to the mean ratio in the saprolite. Thus, we use the stream sediment elemental concentrations to approximate the average bedrock composition along the transect when calculating CDF_Si. This approach introduces considerable uncertainty because of fluvial sorting and because the ratio of argillite to quartzite at each hillslope position may differ from the catchment average. This yields CDF_Si ranging from 0.25 to 0.57 between all soil pits when using Ti as the immobile element and 0.23 to 0.54 when using Zr. The mean CDF_Si along the hillslope is 0.43 (Moore, 2026). Using the mean CDF_Si, the measured riverine flux of Si, and the stream sediment Si concentration in Eq. (2) yields a watershed-average denudation rate of 7.2 tons km⁻² yr⁻¹ (2.8 m Myr⁻¹ at a rock density of 2.6 g cm⁻³). This estimate of the denudation rate has large, but poorly quantified, uncertainties arising from long-term temporal variability in runoff and the assumptions underpinning the calculation of CDF_Si.

Concentrations of ²⁶Al and ¹⁰Be vary with hillslope position (Table 1). Concentrations of both nuclides are highest at the summit (LS1), although this sample was collected from the greatest depth below the surface. Concentrations measured in the valley-bottom sediment sample (LS6) are intermediate between the summit (LS1) and the remaining hillslope positions (LS2–LS5). Neither ²⁶Al nor ¹⁰Be concentrations are significantly correlated with the samples' depths below the modern surface (r2 = 0.05 for ²⁶Al and r2 = 0.15 for ¹⁰Be). The concentrations of both nuclides in the stream sediment sample are broadly comparable with the mean concentrations in the hillslope samples (LS1–LS5) (252 × 10³ atoms g⁻¹ vs. 292 × 10³ atoms g⁻¹ for ¹⁰Be and 1.39 × 10⁶ atoms g⁻¹ vs. 1.30 × 10⁶ atoms g⁻¹ for ²⁶Al).

Aluminum-26 and ¹⁰Be are produced together in quartz at a characteristic production ratio (7.2 at the surface at the study site). The measured ²⁶Al / ¹⁰Be ratios are lower than the surface production ratio and range from 3.29 ± 0.12 at the summit (LS1) to 6.06 ± 0.36 at the toe-slope (LS5). The ²⁶Al / ¹⁰Be ratio in the stream sediment is 5.52 ± 0.19, approximately equivalent to the hillslope positions LS2–LS5, but significantly higher than the ratio measured at the summit (LS1) and in the valley-bottom (LS6). All samples plot further below the steady-state denudation line on an ²⁶Al / ¹⁰Be two-nuclide plot than can be accounted for by the shielding provided by the modern peat cover (Fig. 2). Ratios below the steady-state denudation line are conventionally interpreted to result from deep burial and descent from the line along a decay trajectory. However, the geomorphic position of the study site on a hillslope near a local topographic maximum in a non-glaciated landscape precludes deep burial by sediment or ice during the Pleistocene. Instead, we interpret the depressed ²⁶Al / ¹⁰Be ratios to result from the average peat thickness across the exposure duration of the samples exceeding the modern thickness and apply the models described in Sect. 3.5.

The results of the inverse-model for the hillslope transect are given in Table 3 and Figs. 3 and 4. The median modelled overburden thicknesses (z2) exceed modern overburden thicknesses by 190–350 g cm⁻² along the hillslope transect. This is equivalent to ca. 1.8 to 3.4 m of peat at a saturated peat density of 1.04 g cm⁻³ (assuming 90 % porosity and a solids density of 1.4 g cm⁻³; Lal and Shukla, 2004). Inferred peat thickness is greatest at LS3, on the hill's shoulder, and least at LS4, where the hillslope is steepest. Under these thicknesses, spallation production is severely attenuated and muon reactions are proportionately more important to total production. Thus, the results are only weakly sensitive to the choice of spallation scaling model or reference production rate.

Modelled denudation rates of the saprolite are low: a median rate of 0.32 tons km⁻² yr⁻¹ (equivalent to 0.12 m Myr⁻¹ at a rock density of 2.6 g cm⁻³) was determined for the summit position (LS1) and median rates between 1.8 and 4.9 tons km⁻² yr⁻¹ (0.69 and 1.9 m Myr⁻¹) for the four downslope positions. Integration timescales of the signal range between 0.72 and 1.8 Myr. The long integration timescales result from the slow denudation rates and the greater relative importance of muon-induced production under the peat layer (Table 3; Text S2). This implies that the denudation rates integrate over all glacial-interglacial cycles since the mid-Pleistocene and that any signal from climate-driven oscillations in denudation rate is highly damped. Denudation rate and overburden thickness covary negatively at LS2–LS5 (Fig. S2). Only the summit sample (LS1) clearly resolves a change in overburden thickness. The model for this sample indicates that overburden thickness increased at a median probability age of 1.0 Myr and a modal age of 0.93 Myr (Fig. 5). There is a trade-off between the time of the change and the overburden thickness after the change such that ages greater than ca. 1.4 Myr are allowable only at the lowest possible overburden thicknesses (Fig. S3). The signal integration timescales of LS2–LS5 are too short to fully capture this change.

Modeling of the stream sediment sample (LSR-1) is broadly consistent with a thick long-term peat layer and slow erosion rate at the watershed scale (Fig. 6). Inversion of the ²⁶Al and ¹⁰Be concentrations using Eq. (6) gives a median catchment-averaged erosion rate of the saprolite of 0.64 tons km⁻² yr⁻¹ (equivalent to 0.49 m Myr⁻¹ for a dry saprolite density of 1.3 g cm⁻³). The modelled median peat layer/saprolite quartz ratio is 4.1 × 10⁻³ and mixed layer depth 490 g cm⁻² (equivalent to 4.7 m of saturated peat). The peat thickness and erosion rate covary negatively. Additionally, the posterior distribution is multi-modal and should be treated with caution (Sect. 5.4). For the valley-bottom sediment sample (LS6) modelled using Eq. (7), the erosion rate, quartz ratio, and mixed-layer thickness are essentially unconstrained, as is the burial age above 1 Ma (Table 3; Fig. S4). This indicates that the problem, which requires inverting for 5 parameters from two nuclide measurements, is non-identifiable. Measurement of additional nuclides would be required to constrain the model.

The modelled hillslope denudation rates are more than one order of magnitude below the global mean (Wittmann et al., 2020) and are similar to rates in extremely slowly denuding landscapes in the interior of Brazil (Shuster et al., 2012; Vasconcelos et al., 2019), Australia (Flatley et al., 2025; Struck et al., 2018), and South Africa (Richardson et al., 2025). Are such low rates consistent with independent evidence? The estimate of the watershed-averaged denudation rate from stream solute fluxes and regolith geochemistry is 7.2 tons km⁻² yr⁻¹ (2.8 m Myr⁻¹) (Sect. 4.1), which agrees in order of magnitude with the mean of the median probability rates for the five hillslope positions (2.7 tons km⁻² yr⁻¹ or 1.0 m Myr⁻¹). However, the climate of northwestern Europe during glacial periods was drier than modern climate (Kjellström et al., 2010; Strandberg et al., 2011). When considering that discharge over the timescales of regolith production and erosion, including glacial periods, was lower than modern discharge and that the solute-flux based denudation rate is therefore an upper estimate, the accordance between the two metrics may be even closer. Similarly, the cosmogenic nuclide denudation rate is likely a minimum bound, intermediate between the true denudation rate and the erosion rate. This is because some chemical weathering occurs below the depth of the sample (Riebe and Granger, 2013). If the solutes generated by this weathering are eventually discharged into the stream through groundwater flow, the weathering missed by the cosmogenic nuclide signal will be recorded in the solute flux. Thus, the denudation rates inferred from the hillslope cosmogenic nuclide model are broadly supported by the solute flux.

The magnitude and spatial patterns in denudation that emerge from the cosmogenic nuclide model also align with the geomorphic interpretation of the site as spanning a transition from a relict surface into a slowly incising valley. The lowest denudation rate (0.32 tons km⁻² yr⁻¹, 0.12 m Myr⁻¹) was obtained at the summit position (LS1), which is located within the mapped extent of the pre-Senonian surface (Demoulin et al., 2018b). Preservation of supergene Mn-oxide minerals that yield Cretaceous K-Ar ages indicate that the down-wasting of this surface since the Cretaceous has been insufficient to completely strip the Mesozoic weathering profile (Demoulin et al., 2010). Although the surface's stability may be partly explained through burial by marine sedimentary cover for much of the Cenozoic, the denudation rate determined at the summit is consistent with minimal down-wasting during the Pleistocene. The downslope positions (LS2–LS5) have higher modelled denudation rates (3.2 tons km⁻² yr⁻¹, 1.2 m Myr⁻¹), which supports the incision into the paleo-surface mapped in this area. This incision must have begun before Pleistocene uplift of the Ardennes because the Pleistocene knickzone on the Hoëgne is located downstream of the study area. This is consistent with the establishment of the Ardennian drainage network in the Late Miocene to Pliocene proposed by previous workers (Rixhon et al., 2020; Rixhon and Demoulin, 2018).

Given the low denudation rates, explaining the measured cosmogenic nuclide concentrations requires that the samples experienced greater shielding by overburden, averaged over their exposure histories, than is provided by the modern peat layer (Figs. 3 and 4). Across the transect, the difference between modeled and measured overburden thickness ranges from 190–350 g cm⁻². Might snow or loess cover explain this signal, rather than a thicker peat blanket? On the Hautes Fagnes plateau, current snow thickness rarely exceeds 1 m (ca. 10 g cm⁻² snow water equivalent). Even if maximum snow cover persisted for 3 months per year, mean annual snow shielding would amount to only 2.5 g cm⁻², 2 orders of magnitude lower than the signal. The density of loess greatly exceeds that of snow, creating the potential for more substantial shielding. Accounting for the signal would require continuous cover of 1.1–2.0 m of loess over the integration timescale of the signal (assuming a saturated loess density of 1.8 g cm⁻³). In the central Belgian loess plain deposits exceed this depth (Haase et al., 2007), yet loess thicknesses diminish southwards towards the Ardennes (Rixhon et al., 2011) and, at present, no loess cover exists at the site (Lehmkuhl et al., 2021). Moreover, the peat layer contains a low ash (mineral) fraction, and the coarse bedload fraction (> 2 mm) of the stream sediment, which likely approximates watershed-averaged bedrock (Sect. 4.1), has TiO₂ / Zr ratios that are within the range of the saprolite's ratios (Data Tables in the Supplement; Moore, 2026), supporting that the saprolite originates from in situ weathering of bedrock. This implies that any loess cover deposited after the LGM on the study hillslope must have been sufficiently thin to be completely removed by Holocene erosion. Therefore, it is unlikely that snow or loess cover accounts for the bulk of the missing time-averaged overburden.

The Hautes Fagnes were under permafrost conditions at the LGM (Demoulin et al., 2018a; Lindgren et al., 2016; Pissart et al., 2011), and presumably during earlier glacial maxima. During these periods, segregation ice was likely present at the site and may have contributed to the overburden thickness. For example, the lithalsa ramparts preserved in the Hautes Fagnes are evidence for ice lenses several meters thick (Pissart et al., 2011). Still, the local nature of these lenses, which were ca. 10s of meters in diameter in the Hautes Fagnes area and widely spaced, makes them an unlikely source of the modeled signal, which is similar in magnitude across the hillslope. Furthermore, permafrost would have existed at the study site only near glacial maxima, limiting the time-averaged shielding effect. Thus, ground ice is also unlikely to fully account for the modeled overburden deficit.

The peat depths required to explain the signal are consistent with the limits to peat thickness predicted by theory. Climate in Belgium is modeled to have been broadly conducive to peat formation over most of the last glacial cycle (Marine Isotope Stages 5-2), and thus likely for most of the period since the mid-Pleistocene transition (Fig. S.5 in Treat et al., 2019). Over sufficiently long timescales peat thickness tends towards steady state, where inputs of organic matter are balanced by decomposition within the peat deposit (Clymo, 1984). The stability of peat is strongly controlled by the height of the water table, which is a function of net precipitation (P-ET) and the groundwater flow rate within the peat body. The latter depends on the hydraulic conductivity and head gradient. If peat is preserved in the long-term only below the water table, then a bog's surface topography will mimic the topography of the water table. This principle underlies the “groundwater-mound” model of bog morphology. In this model, the thickness and surface taper of a peat bog over a relatively impermeable substrate are explained by the need to accommodate a lateral water flux that increases with distance from the drainage divide, which requires a corresponding increase in the bog-surface (hydraulic head) gradient (Cobb et al., 2024; Ingram, 1982). The steady-state peat thickness at the center of a bog predicted by this model depends on the water balance, but is typically ca. 10 m.

The modeled long-term overburden thicknesses correspond to reasonable steady-state peat depths at the study site. Peat thicknesses greater than 8.5 m have been reported for raised bogs on plateau surfaces in the Hautes Fagnes (e.g., Misten Bog; Allan et al., 2013). These bogs must be at or below the modern steady-state peat thickness on a flat surface for the local climate. The median long-term peat thicknesses inferred from the cosmogenic nuclide model are below this threshold, reaching a maximum of 4.2 m at LS3 and a minimum of 2.5 m at LS4, consistent with the expectation of thinner peat on sloping landscapes. These thicknesses are broadly comparable to values reported for undisturbed blanket peat on similarly low-gradient slopes (< 3°) under comparable climatic conditions in, for example, the Wicklow Mountains of Ireland (Holden and Connolly, 2011). Additionally, the “groundwater mound” model requires that horizontal discharge increases linearly downslope from the drainage divide to accommodate water inputs across the bog's surface. The horizontal flux of water (Q) through a peat layer can be described using the Dupuit-Forchheimer approximation to Darcy's law (Belyea and Baird, 2006). In its simplest form, this gives Q = kDΔh, where k is hydraulic conductivity, D is peat thickness, and h is the hydraulic head (or surface) gradient. For spatially homogeneous hydraulic conductivity and net precipitation, the product of peat depth and the head/surface gradient must increase linearly downslope to accommodate the increasing discharge (Ingram, 1982). The product of the overburden thickness and surface gradient inferred from the cosmogenic nuclide model follows this trend (Fig. S5), indicating that the modeled long-term overburden thicknesses meet the physical constraints of a steady-state peatland.

The typical climate in northwestern Europe over the averaging timescale of the cosmogenic nuclide signal was colder and drier than the modern climate. How would cooler and drier conditions impact steady-state peat thickness? For example, modeling of European climate during a representative Marine Isotope Stage 3 stadial at 44 ka indicates that MAP was ca. 25 %–30 % lower and MAT ca. 5 °C lower than at present (Kjellström et al., 2010). If these relative differences are applicable at the scale of the Hautes Fagnes, then MAT was ca. 1.7 °C and MAP ca. 1050 mm yr⁻¹. This places the site near the high-precipitation end of the distribution of peatlands in modern precipitation-temperature space (Cobb et al., 2024), and slightly above the temperature threshold for permafrost formation (MAT < 0 °C). Moreover, because ET at 44 ka would have been lower than modern ET (ca. 500 mm yr⁻¹, Sect. 4.1) due to lower MAT, the precipitation to ET ratio would have exceeded 2.1, which is believed to be the threshold for blanket peat formation (Gallego-Sala and Prentice, 2013). In the “groundwater-mound” model, the steady-state peat thickness at the center of a bog is proportional to the square root of net precipitation (Ingram, 1982). Thus, a < 40 % reduction in net precipitation (> 550 mm yr⁻¹ vs. 918 mm yr⁻¹) during permafrost-free stadials would reduce the peat thickness by only ca. 20 % at the drainage divide. During even colder phases (e.g., the LGM) the site would have been under permafrost conditions. Under permafrost, peat stability ceases to be controlled by saturation and is instead maintained by low temperatures. This enables preservation of thick organic layers even under low precipitation rates. For example, in modern high-latitude permafrost, peat may extend to depths of greater than 3 m (Hugelius et al., 2013).

We suggest that the difference between the overburden thickness inferred from cosmogenic nuclides and the modern peat thicknesses measured by ground penetrating radar mostly reflects anthropogenic disturbance. The land use history of the Hautes Fagnes is well-documented in the historical and palynological records. Although cereal pollen is first detected in the Neolithic, evidence for human activity rises markedly in the High Middle Ages (the 13th century) and significant exploitation of the Hautes Fagnes peatlands began at this time (De Vleeschouwer et al., 2012). For example, extensive mowing and raking of litter and sphagnum for animal fodder and bedding, livestock grazing, slash-and-burn cultivation of rye on peat soils, and peat wildfires are documented in historical records and peat cores (Damblon, 1996; Froment, 1968). The net effect of these practices would have been to reduce the supply of fresh organic matter and shift the system to a negative mass balance, wherein peat decomposition exceeds litter inputs. The rate of anaerobic decay occurring below the water table is proportional to peat mass and a decay constant on the order of magnitude of 10⁻⁴ yr⁻¹ (Clymo, 1984). For a peat layer with a thickness of 400 cm, this suggests loss of ca. 30 cm in 800 years, if the supply of fresh litter is shutoff. Peat extraction for fuel is also recorded in historical documents since the 16th century (Paulissen et al., 2021). Annual peat extraction from the Hautes Fagnes at the beginning of the 19th century is recorded as nearly 200 000 quintals (or ca. 200 000 m³ of peat at a density of 100 kg m⁻³) (Froment, 1968). As an illustrative calculation, if this rate of extraction was maintained from the beginning of the 16th to the end of the 19th century and distributed evenly across the 15 000 ha of peatland documented on the Hautes Fagnes plateau in the late 18th century (Froment, 1968), this would amount to 50 cm of peat loss. Because extraction likely targeted thick and accessible deposits, local losses could be significantly larger. Although no peat-cutting scars are evident in the study transect, it is possible that these were destroyed during drainage and afforestation in the early 20th century, which indisputably accelerated peat degradation. Peat subsidence rates in blanket and raised bogs after drainage and afforestation reported in the literature range from ca. 1 to 2 cm yr⁻¹ (Adetsu et al., 2024; Oleszczuk et al., 2021; Shotbolt et al., 1998). If a similar rate were sustained since the early 20th century across the study transect, then this implies 1–2 m of peat subsidence. Thus, the combined effects of drainage, peat cutting, and other land-use practices may have driven > 2 m of peat loss. This likely explains most of the discrepancy between the modern measured and modeled long-term overburden thicknesses.

The inverse-model for the summit position (LS1) indicates that this sample captures a secular thickening in the overburden. The greatest probability density for the timing of this change is between 1.4 and 0.8 Myr (inter-quartile range), with a median of 1.0, and a mode of 0.93 Myr (Table 3; Fig. 5). This may represent the initial development of a thick peat blanket at the study site. Interestingly, the timing of this event corresponds closely to the uplift of the Ardennes recorded by the fluvial terraces of the Meuse and Rhine rivers, which cross the Ardennes-Rhenish massif. Abandonment of the East Meuse valley and an increase in Meuse terrace gravel content indicate that gradual uplift began in the early Pleistocene (da Silva Guimaraes et al., 2024). An acceleration in uplift is recorded by the abandonment of the Younger main terrace, a prominent terrace at mid-elevation in the Meuse and Rhine valleys. This has been dated to 0.73 ± 0.12 Myr in the Meuse valley using cosmogenic nuclide depth-profile dating (Rixhon et al., 2011) and 0.52 ± 0.11 Myr in the Meuse valley and 0.76+0.19/-0.17 Myr in the Rhine valley using ²⁶Al / ¹⁰Be burial dating (da Silva Guimarães et al., 2026; Terraza et al., 2026). This uplift pulse, combined with the subsequent isostatic response to erosional unloading, amounted to ca. 100 m of total uplift in the vicinity of the Hautes Fagnes plateau (Demoulin and Hallot, 2009). The summit position (LS1) is presently at an elevation of 630 m above sea level and thus was at ca. 530 m immediately before the uplift pulse. Today, the cool, moist climate of the study area, which is conducive to peat accumulation, is a result of moist air from the Atlantic rising over the Hautes Fagnes plateau and the resulting orographic precipitation, persistent cloudiness, and low ET. These orographic effects become pronounced above ca. 500 m. This indicates that the site may have reached sufficient elevation for peat accumulation shortly preceding or during the uplift pulse recorded by the Younger main terrace. The modeled timing of the overburden thickening event is in good agreement with this interpretation. Although for simplicity the ca. 10 % increase in spallation production rates associated with 100 m of uplift has not been incorporated into the model, considering this change would modestly lower the initial overburden thickness estimate and not significantly impact the remaining parameters.

The inversion of the stream sediment nuclide concentrations is roughly consistent with the inferences drawn from the hillslope samples. Converting the median probability catchment-averaged erosion rate of the saprolite (0.64 tons km⁻² yr⁻¹) to a denudation rate using the mean CDF_Si on the study hillslope as an approximation of catchment-average bulk weathering intensity gives a denudation rate of 1.5 tons km⁻² yr⁻¹ (0.57 m Myr⁻¹). This is intermediate between the denudation rate obtained at the summit position (0.32 tons km⁻² yr⁻¹/0.12 m Myr⁻¹) and the mean of the downslope positions (3.2 tons km⁻² yr⁻¹/1.2 m Myr⁻¹) and is of the same order of magnitude as the solute-flux based denudation rate (7.2 tons km⁻² yr⁻¹/2.8 m Myr⁻¹). The median probability catchment-averaged peat thickness is 490 g cm⁻² (4.7 m), greater than the mean of the modeled thicknesses along the hillslope of 370 g cm⁻² (3.6 m) and consistent with thicker peat on the low-relief surface in the headwaters of the catchment (Fig. 1). The median quartz concentration ratio between the saturated peat layer and saprolite is low (4.1 × 10⁻³), which is expected for ombrotrophic peats with minimal input of mineral material from below.

Nevertheless, there are several problems with inferring catchment-scale process rates from detrital sediment at the study site that complicate the sediment model. Notably, the site violates the assumption that radioactive decay is negligible and that the nuclide concentration is a simple, inverse function of the denudation rate, which is required for basin averaging (Granger et al., 1996). Furthermore, the depressed ²⁶Al / ¹⁰Be ratio in the valley-bottom sediments indicates that storage times, at least locally, may be > 1 Myr (Fig. S4). Remobilization of fluvial sediments that experienced significant radioactive decay during storage has the potential to bias the results. Therefore, the results from the stream sediment should not be over-interpreted. Yet, they broadly caution against inferring bedrock erosion rates in peatland environments from the concentration of a single cosmogenic nuclide in fluvial sediment. At the study site, simple interpretation of the ¹⁰Be concentration in the stream sediment as an erosion rate gives 43.5 tons km⁻² yr⁻¹ (ca. 16.7 m Myr⁻¹), which would overestimate the landscape lowering rate by an order of magnitude. Similarly depressed ²⁶Al / ¹⁰Be ratios in stream sediment and spuriously high single nuclide denudation rates have been reported in other European massifs, supporting this cautionary note (Jautzy et al., 2024).