We used petrographic microscopy, powder X-ray diffraction, and electron beam microanalysis to identify and characterize the structure and mineralogy of the singular marker horizon M in the Lower to Middle Eocene bituminous black pelite of the maar lake Messel. The marker horizon represents a short exceptional sedimentation episode with a supraregional origin, which cannot be explained by run-off debris from the crater wall or from an external inflow by chemical and textural interpretation of the sediment. The phosphate mineral assemblage of messelite, montgomeryite, and mantienneite indicates a significant change of the sediment composition characterized by an enrichment of, e.g., P and Ti, that is exceptional during a Lake Messel depositional history of about 640 kyr (Lenz et al. 2011, 2015; Lenz and Wilde 2018). The vital elements to reconstruct the marker horizon formation are the precursor composition, phosphorus sources and transport, and fluid–mineral interaction processes that control the mineralogy.
Marker horizon precursor sediment
The formation of abundant montgomeryite and mantienneite within the marker horizon M implies a bifurcation of the depositional process and sediment precursor composition. The rather uniform thickness of the marker horizon across the entire Lake Messel and compositional singularity argues against sedimentary input from the crater wall. Trace minerals such as alkali feldspar, apatite, and mixed Fe–Ti oxides occur throughout the black pelite of the Middle Messel Formation (Weber and Zimmerle 1985; Kubanek et al. 1988) and the ~ 9 mm thick central band of the marker horizon. However, only the marker horizon contains montgomeryite and mantienneite. This indicates a significant change of the sediment composition not found elsewhere within 640 kyr of Lake Messel sedimentation (Lenz et al. 2011, 2015; Lenz and Wilde 2018). Apparently, the sediment that preceded the marker horizon did not derive from the weathered crater rim, but entered the lake via an alternative pathway, for example, a tephra cloud, as suspected previously for the older marker horizons α and β (Weber 1988). Mantienneite enrichment in the upper half of the central band could reflect a change in tephra composition, with an upward increasing Ti concentration. A potential cause for increasing grain sizes and changing compositions toward the top of the central band could be a two-phase eruption with a stronger second eruption that more violently vented tephra in a larger eruption cloud. Another possibility is crystal fractionation in a magma chamber. In this scenario, an early tephra layer formed the lower part of the central band, while Ti-rich minerals with high density (e.g., ilmenite, Ti-rich pyroxenes, etc.) were assimilated as xenocrysts during replenishment of a magma chamber and transported during protracted explosive volcanism. Thus, a second tephra cloud with a higher proportion of these minerals may have formed during an immediately following eruption. Together with glass shards, Lake Messel may then have accumulated Ti-rich minerals in a distinct horizon within the tephra band as a source for later mantienneite formation (Fig. 3d).
The original petrographic composition of MH M can only be incompletely reconstructed due to the strong alteration and phosphatic replacement of the primary minerals. Only the stubby shape of the pore spaces in mantienneite-rich domains, resembling euhedral crystals of pyroxene and feldspar, and the ions provided for the formation of the various phosphates, such as Al (relicts of feldspar and/or clay-sized Al-rich components), Mg, K (from alkali feldspar, biotite), and Ti (e.g., from pyroxene, amphibole, biotite, ilmenite), indicate an originally volcanogenic composition of the source sediment.
The micrometer-sized cavities left by dissolved ash particles and minerals in MH M suggest that crystals and shards originate from the distal parts of a faraway wind-driven ash cloud that partly settled over Lake Messel. Therefore, the nearby Eocene eruption centers on the Sprendlinger Horst, in the Taunus Mountains, or within the Upper Rhine Graben Rift Zone, the latter today buried under kilometers of younger Tertiary and Pleistocene sediments (Lutz et al. 2013), are less probable as a cause. In terms of grain size, distant ashes from the post-breakup North Atlantic Igneous Province (NAIP) or the pre-rift Massif Central, which were active in the time of Messel (Michon and Merle 2001; Meyer et al. 2007; Wilkinson et al. 2017) are potentially to be considered (e.g., NAIP generated bentonites in Austria; Huber et al. 2003), but this is speculative, due to the intense diagenetic overprint of the precursor material.
Sources of phosphate
Various autochthonous and allochthonous sources contributed to the organic matter and associated phosphorus content of the Messel black pelite over time (Bauersachs et al. 2014). A considerable amount of phosphate input into Lake Messel results from internal organic matter, solutions from the paratropical soils by run off and groundwater, subaquatic alteration of pyroclastic particles, and weathered basanite-nephelinite crater rim debris (Kubanek et al. 1988). According to Porder and Ramachandran (2013) basanitic–nephelinitic basalts have median P concentrations from 3000 to 4000 µg/g. To a much lesser extent, crater wall material such as granodiorite (median 698 ppm), diorite (median 1004 ppm) and amphibolite (median 655 ppm) may have contributed to the phosphate content of the Messel black pelite (Mezger et al. 2013). The crater wall lithologies could partly provide P for phosphate precipitation and other elements of the marker horizon. However, the absence of a horizon similar to MH M suggests that the P content of the basanitic–nephelinitic precursor of the Messel oil shale was not enough to achieve the complete transformation of the exceptional marker horizon sediment into phosphate minerals. For this, an additional P supply was necessary.
This supply is likely related to the abundant sedimentary organic matter in the Middle Messel Formation. The high molar Corg/Ntot ratios of the Middle Messel Formation black pelite reflects a predominantly terrestrial origin of the preserved sedimentary organic matter. This includes plant litter from a dense paratropical rainforest near the maar lake. Comparatively heavy δ15Ntot as well as high Corg/Ntot ratios and HI values of the of the sedimentary organic matter suggests a high number of lipid-rich constituents of vascular plants of the total organic matter (Bauersachs et al. 2014).
Autochthonous sources, for example, bacteria (stored as polyphosphates in granules, e.g., Cosmidis et al. 2014) and algae from the epilimnion of Lake Messel contributed in variable proportions to the organic matter content over time. Stable carbon isotope excursions to heavier δ13Corg values indicate periods of increased algal productivity by different (pico-)phytoplankton species (e.g., cyanobacteria, Chrysophyta indet., the dinoflagellate Messelodinium thielepfeifferae, the coccal green algae Tetraedron minimum, Botryococcus sp. as well as Coelastrum sp.) and increased input of autochthonous organic matter to the lake sediments (Goth 1990; Lenz et al. 2011; Richter et al. 2013; Bauersachs et al. 2014). The overall depletion of δ13Corg values together with an increase in δ15Ntot values throughout the black pelite evidence microbial reworking of the sedimentary organic matter by methanogens, methanotrophs and denitrifiers (Bauersachs et al. 2014). Later, anaerobic digestion liberated organically bound phosphorus as orthophosphate into solution that may bind with cations including Mg, Ca, Al, Fe, and Ti to form the phosphate minerals that dominate the marker horizon M.
The exceptional phosphate accumulation associated with MH M is potentially linked to the input of the volcanic ash precursor into the water column. It has been shown that volcanic ash dissolving in lake- or seawater modifies the nutrient budget of the surface lake and ocean and stimulates the growth of phytoplankton, for example, in iron-limited lake/oceanic areas (e.g., Mattews-Bird et al. 2017; Duggen et al. 2007; Browning et al. 2015). The amount of bioavailable elements is directly proportional to the volcanic ash thickness (Duggen et al. 2007). Since ocean production and export of organic carbon transfers CO2 from the atmosphere to the ocean interior, volcanic ash fertilization may play a vital role for the ocean–atmosphere gas interchange and ultimately the development of the global climate (Hamme et al. 2010; Hamilton et al. 2022). In oligotrophic lakes, volcanic ash input can result in 1.5- to eightfold increases in total suspended solids, light extinction, phosphorus concentrations, and phytoplankton biomass relative to pre-eruption conditions (Modenutti et al. 2013). Potentially, a transient micronutrient enrichment supplied by a volcanic ash resulted in intense bacteria and algae blooms in Lake Messel as well. Such blooms significantly enhance the diffusive fluxes of soluble reactive phosphate and iron from sediment pore water to the overlying water (Chen et al. 2018; Wang et al. 2022). While an increased pH value of the water column and upper sediment layer promotes the desorption of phosphate from mineral surfaces (Gao et al. 2014), anoxic conditions at the sediment–water interface cause the reductive dissolution of iron oxide minerals and the coupled release of iron and surface-bound P into pore water (Smith et al. 2011; Cosmidis et al. 2014).
The decomposition of algae during or after an algae bloom results in the release of P from degraded algal cells and re-release to the sediment and overlying water (Wang et al. 2022). Settling tephra particles in Lake Messel could also have adsorbed P from the water column and contribute significantly to the synsedimentary, exceptional P accumulation in the sediment within and adjacent to the later marker horizon. Early diagenesis mobilizes these accumulations and phosphorus diffuses and reacts with the volcanic ash. In this context, high-resolution palynological study (e.g., Moshayedi et al. 2020) of the sediment near the marker horizon M may contribute to our understanding of the response of ecosystems in a volcanically disturbed habitat, e.g., harmful bacteria or algae blooms in a lacustrine environment or the abrupt scarceness of Messel pit fishes, plants, and arthropod remains after a volcanic eruption (e.g., Micklich, 2012; Lu et al. 2021).
Conclusively, the formation of the marker horizon requires a relative phosphorus enrichment to the typical black pelite sediment. Indicators for an exceptionally thick organic layer, such as a microbial mat grown on the sediment surface (Felder 2007), and bacteriomorphic phosphatic compartments (Schmitz 1991; Liebig et al. 1996) were not observed. The occurrence of black pelite in the messelite bands shows that it is unlikely that event sedimentation of weathered crater rim material directly provided the amount of phosphorus required for the formation of the marker horizon. Our results indicate that, contrary to earlier studies that favored a direct precipitation of the marker horizon at the sediment–lake interface (Schaal et al. 1987; Goth 1990; Felder 2007), it is more likely that phosphorus enrichment resulted from alteration of P-rich precursor sediment, e.g., phosphatized Coelastrum-algae layers (Richter et al. 2013) and P adsorbed to mineral surfaces, and subsequent pore water diffusion toward and within the marker horizon during diagenesis. Later phosphate mobilization from sediment below the marker horizon resulted in the formation of vertical messelite veins that crosscut the entire marker horizon.
Element transport and phosphate precipitation in the marker horizon
The diverse mineralogy of MH M indicates that the pore fluid composition changed continuously during diagenesis. The phosphate precipitation sequence records these changes that are likely related to, e.g., the specific surface area of primary minerals, primary mineral dissolution rates and sequence, and kerogen maturation (Curtis 1983; Weibel and Friis 2004; Wilson 2004; Sindern et al. 2019). Among the detrital minerals, those with a large specific surface area are most likely to be preserved as negative crystal shape, e.g., associated with mantienneite (Fig. 4d). Since the marker horizon mineralogy lacks indicators for vertical diffusion gradients, most mineral reactions can be considered to result from element re-distribution within the marker horizon. The low mobility of Al and Ti indicates that the marker horizon itself is the source and sink of these elements.
During diagenesis of the marker horizon, the pore fluid will repeatedly be out of equilibrium with primary and authigenic minerals. Thus, it is unavoidable that fluid-mediated, coupled mineral dissolution and precipitation processes will take place (Ruiz-Agudo et al. 2014). It should be noted that only trace amounts of minute non-phosphate minerals are preserved in the marker horizon. Thus, their scarcity lowers their value as geochemical indicators. Among the trace minerals, silica spheres in messelite are by far most abundant. They document the important role of silicate dissolution during phosphatization of the marker horizon.
Diagenetic processes alter the mineral components of the marker horizon as well as the organic material. Kerogen maturation and hydrocarbon mobilization is an omnipresent process in the Messel oil shale and can play an important role for pore fluid composition, mineral dissolution/precipitation, and element mobility (Curtis 1983; Waldmann and Gaupp 2016; Sindern et al. 2019). High molecular weight organic matter readily loses carbon dioxide, which is soluble and dissociates (Curtis 1983). This process not only produces acidified, reducing pore fluids but also releases large quantities of organically and mineral surface-bound phosphate. Further kerogen maturation is a constant source of these acidic, reducing fluids until the reducing agent (organic material) is exhausted (Weibel and Friis 2004).
Messelite has been known from its type locality, the Messel pit, since 1890 (Muthmann 1890). Other sedimentary occurrences are the equally old deposits of the Prinz von Hessen mine (Dietrich 1978) as well as lacustrine argillaceous sandstones intercalated with Jurassic brown coal seams in the Kostanay region of Kazakhstan (Vertushkov 1952). Previous work mainly dealt with the crystallographic classification (e.g., Taborszki 1977; Dietrich 1978; Goth 1990; Fleck and Kolitsch 2003).
In the marker horizon M, messelite in the upper and lower band encloses single black pelite layers and packages of up to five laminae (Fig. 2). All platelets show the same degree of compaction, regardless of their spatial position. The intercalated black pelite shows that portions of the marker horizon precursor material had a composition like the bulk black pelite of the Messel pit. Evidently, the sedimentation of black pelite, algae layers, and marker horizon precursor material lasted several years. The subsequent diagenetic messelite formation resulted from the reaction of migrating phosphorus from organic material layers and mineral surfaces, and elements crucial for messelite precipitation, e.g., Fe and Ca, from the tephra (Table 1). Small ash particles with a large specific surface area and high dissolution rate are especially susceptible to early diagenetic dissolution (Wilson 2004). These reactive minerals comprise, e.g., olivine, pyroxene, and amphibole. Dissolution of these minerals readily liberates Fe and Ca for messelite formation into solution. The low Al concentration of messelite suggests that plagioclase dissolution played only a minor role in Ca liberation, while the presence of ferrous Fe (Table 1) indicates reducing conditions likely associated with hydrocarbon mobilization. Continuous messelite crystal growth in the lower and upper band fractured and disturbed the black pelite layers within the marker horizon. This indicates that the formation of the black pelite was completed when messelite crystal growth began and that, contrary to previous interpretations (Schaal et al. 1987; Goth 1990; Felder 2007; Tütken 2014), the marker horizon M did not precipitate directly at the sediment–water interface but formed diagenetically.
Central montgomeryite globule band
Sharp mineralogical and textural boundaries between the messelite bands and the central messelite–montgomeryite–mantienneite band indicate an abrupt change of the marker horizon precursor composition. The central band is characterized by a specific order of phosphate precipitation, as well as trace mineral type and abundance, and the absence of black pelite relicts. The textural relationships suggest the following precipitation sequence (Fig. 5): (1) a messelite groundmass, (2) montgomeryite globules, (3) mantienneite, and (4) late vertical messelite veins with similar composition to the earlier bands.
The precursor to messelite in the central band is completely altered, with only a few silica spheres left. The dissolution of messelite and replacement by montgomeryite (Fig. 3b) releases Fe to the pore fluid that may migrate and form iron oxide/hydroxide elsewhere. Phosphorus from the alteration of messelite and organic matter is incorporated into montgomeryite. At low pH, the alteration of detrital apatite (Fig. 3c) can also provide P for montgomeryite formation (Nriagu 1976). Compared to messelite, montgomeryite contains large amounts of aluminum, ~ 0.1 vs ~ 16 wt% Al2O3, respectively (Table 1). Since messelite dissolution provides only negligible amounts of aluminum, dissolving feldspar and other Al-rich minerals are required as a local source for an Al-rich microenvironment. While aluminum is considered fluid immobile at circumneutral pH in low-temperature settings, its mobility increases significantly at low pH (10–3 mol/kg at pH 3; Curtis 1983). Migrating hydrocarbons from remnant dissolving organic compounds in- and outside the marker horizon are a likely source of this acidic (probably reducing) pore fluid. Subsequent montgomeryite precipitation may induce a concentration gradient that promotes Al diffusion toward mineral surfaces and facilitates montgomeryite globule formation (Fig. 3a).
The authigenesis of mantienneite documents an exceptional sedimentation episode during the formation of the Messel black pelite and the marker horizon M. To date, mantienneite has been analytically identified at three locations worldwide (including this study)—Anloua (Cameroon), Pinciná (Slovakia), and Messel (Germany). All three locations share a common association of lithologies dominated by clay minerals and organic matter. Mantienneite was first described by Fransolet et al. (1984) from the vivianite deposit of Anloua (Cameroon), a series of lacustrine deposits of Upper Tertiary to Quaternary age. In this series spherulites of mantienneite, a clay fraction, and siderite cement a black bituminous argillite with some thin sandy layers. A second occurrence is an alginate deposit near Pinciná, Slovakia (Vavrová et al. 2006). The alginate, a mixture of organic matter and clay minerals, was deposited in the Pinciná Maar lake that formed during Upper Miocene–Pliocene basalt volcanic activity (Hurai et al. 2021). Well-developed mantienneite spheres composed of radiating crystals occur in trace amounts in some horizons dominated by smectite and kaolinite. A hydrothermal overprint of microbially degraded algae and basaltic volcaniclastics at the maar surface has been suspected to induce mantienneite precipitation directly from solution (Vavrová et al. 2006). However, a hydrothermal overprint of the Middle Messel Formation black pelite has not been reported yet. We infer that the textural dissolution–precipitation relationships between the phosphate minerals in the central band show that mantienneite did not crystallize directly from a solution at the lake-sediment surface.
The dissolution of messelite and montgomeryite can provide Al and P required for mantienneite precipitation (Table 1). Mantienneite authigenesis also requires a local increase of the Mg, K, and Ti activity. The dissolution of minerals, such as pyroxene, alkali feldspar, biotite, and amphibole likely provided Mg and K. BSE images show that mantienneite replaces all other phosphates, which means that local dissolution of detrital Ti-rich minerals, e.g., Ti-augite, Ti-rich amphibole, biotite, and ilmenite must have liberated additional Ti. This agrees well with previous studies that interpreted mantienneite authigenesis in sediments rich in clay minerals and organic matter to result from the reaction of phosphorus-rich solutions with altering minerals, e.g., titanian augite and ilmenite (Fransolet et al. 1984). The silica sphere-filled pore space resembling stubby euhedral crystals in mantienneite (Fig. 4d) could be remnants after dissolution of such detrital minerals. Similarly to aluminum, Ti is usually fluid immobile in low-temperature geological systems, but its mobility increases considerably in low pH fluids rich in dissolved organic compounds (Hausrath et al. 2009; Fuchs et al. 2015; Schulz et al. 2016; Liu et al. 2019; Sindern et al. 2019). Kerogen maturation likely generates an acidic, reducing fluid, which also intensifies the leaching of detrital minerals, e.g., ilmenite (Weibel and Friis 2004). However, an increased titanium mobility at low pH, at least on the microscale, is commonly associated with oxidizing conditions (Weibel 1998; Sindern et al. 2019). Further, the presence of ferric Fe in mantienneite and precipitation of iron oxide/hydroxide indicates oxidizing conditions during mantienneite formation. The mantienneite-forming reaction, thus, marks the change from a regime of reducing to oxidizing conditions, possibly upon exhaustion of the dissolved organic compounds. Vertical messelite veins cross-cutting the entire marker horizon demonstrate that later kerogen maturation created another pulse of reducing pore fluid from below the marker horizon.