Introduction

The early evolution of animals, or metazoans (Metazoa), can be traced back to the Ediacaran Period (635–541 Ma) based on the fossil record of macroscopic soft-bodied impressions, carbonaceous compressions, and subsequently mineralized and organically preserved organisms (Narbonne, 2005; Moczydłowska et al., 2014; Wan et al., 2016; Xiao et al., 2016; Schiffbauer et al., 2020). However, a more ancient metazoan origination event has been inferred using molecular clocks, which suggests emergence times ranging from between 833–650 Ma (dos Reis et al., 2015), to a more specific point estimate of up to 780 Ma (Erwin et al., 2011). The affinities of the various Ediacaran metazoan taxa are likewise unresolved and are often contested because of their unusual morphologies and lack of unambiguous diagnostic features in comparison to modern phyla; these have resulted in assorted identifications including sponges, placozoans, cnidarians, lophophorates, and putative bilaterian metazoans (Narbonne, 2005; Fedonkin et al., 2007a, b; Xiao and Laflamme, 2009; Sperling and Vinther, 2010; Erwin et al., 2011; Sperling et al., 2011; Zhuravlev et al., 2015; Hoyal Cuthill and Han, 2018; Wood et al., 2019; Schiffbauer et al., 2020; Shore et al., 2020).

The preserved remains of Ediacaran metazoans seem to represent body tissues that were largely composed of soft organic matter (Moczydłowska et al., 2014; Narbonne et al., 2014; Yin et al., 2015; Wan et al., 2016; Bobrovskiy et al., 2019; Schiffbauer et al., 2020), although calcified skeletons had appeared by ca. 550 Ma (Germs, 1972; Grant, 1990; Grotzinger et al., 2000; Wood et al., 2002, 2019; but see Schiffbauer et al., 2020, and Yang et al., 2020). The onset of metazoan mineralization is a precursor to the independent acquisition of biomineralized skeletons and of different mineralogies by multiple metazoan clades, an important feature of the Cambrian Explosion (541–485 Ma; Erwin et al., 2011; Erwin and Valentine, 2013; Murdock, 2020).

Here we report an assemblage of three-dimensionally (3-D) preserved macroscopic tubular and conical fossils from marine siliciclastic strata of Ediacaran age on the Digermulen Peninsula of Arctic Norway. These specimens represent three distinct morphotypes, which we recognize as distinct taxa, and are thus interpreted as components of a coeval assemblage that is stratigraphically restricted to the basal succession of the Stáhpogieddi Formation. This interval lies 40–42 m above the Mortensneset Diamictite—an uppermost marker of the Varangerian glaciations in Scandinavia (Føyn, 1985; Nystuen, 1985) and correlated with the regional Ediacaran Gaskiers glaciation at ca. 580 Ma (Halverson et al., 2005, 2010; Rice et al., 2011; Pu et al., 2016). The affinities of the new taxa remain uncertain, but collectively all features suggest eumetazoans. We apply a range of microscopic, computed micro-tomographic, petrographic, and geochemical techniques to determine their morphology and mineralogy. This reveals evidence of primary siliceous biomineralization, which to our knowledge, constitutes the earliest examples of silica-based animal skeletons yet documented from the fossil record. The global stratigraphic ranges of the Ediacaran impression fossils, including Dickinsonia Sprigg, 1947, Swartpuntia Narbonne et al., 1997, Hiemalora Fedonkin, 1982, and Aspidella Billings, 1872, and 3-D modular Palaeopascichnus Palij, 1976 occurring in the Stáhpogieddi Formation (Farmer et al., 1992; Högström et al., 2013, 2017; Høyberget et al., 2017; McIlroy and Brasier, 2017; Jensen et al., 2018; additional new data on Dickinsonia and Swartpuntia per A. Högström, personal communication, 2018; Fig. 2) are compiled and the sedimentary continuity of this formation is discussed with the purpose of estimating the age of the new taxa.

Geological setting, stratigraphy, and age

Tectonic setting.—The Stáhpogieddi Formation is well exposed across the Digermulen Peninsula in far northeastern Finnmark, Arctic Norway. Structurally, the region forms part of the Gaissa Thrust Belt (or Gaissa Nappe Complex) within the Lower Allochthon of the eastern Scandinavian Caledonides. The Digermulen Peninsula is in a continuous transition with the parautochtnonous and autochthonous areas around Tanafjorden and on the Varanger Peninsula (Gee and Sturt, 1985; Gayer et al., 1987; Siedlecka et al., 2006; Rice, 2014; Fig. 1).

Geologically, it has been subjected to limited tectonic, diagenetic, and thermal modifications in comparison to the highly metamorphosed nappes in the central Caledonides, which have a complex history of folding and décollement (Roberts and Gee, 1985; Gayer et al., 1987; Dallmeyer et al., 1989; Roberts, 2003; Rice, 2014; Meinhold et al., 2019a). The Gaissa Thrust Belt has experienced low angle thrusting over a relatively short distance (Rice, 2014) and is sequentially sub-divided into the Vadsø, Ekkerøya (elsewhere considered a formation-level unit of the Vadsø Group; Føyn, 1985), Tanafjorden, Vestertana, and Digermulen groups, with a total succession thickness of ∼5 km.

Figure 1.

Sketch-map of Scandinavia with studied area marked in northern Norway by arrow in (1), pointing to geological map of the Ediacaran-Cambrian strata on the Digermulen Peninsula coastal outcrop in (2), showing the new fossils locality. The outline of the Varanger and Digermulen peninsulas and the Tanafjorden area is in (3) with Gaissa Thrust Belt shaded gray.

The Vestertana Group, including the studied Stáhpogieddi Formation (Fig. 2), was deposited in a foreland basin under shallow marine to deep outer shelf sedimentary conditions on a passive continental margin (Roberts and Siedlecka, 2002; Rice et al., 2012; Rice, 2014; W. Zhang et al., 2015). This basin developed along the western Baltica paleocontinent at ∼50°S (orientation at the time; Torsvik and Cocks, 2017).

Stratigraphic information.—The Vestertana Group extends stratigraphically from the upper Cryogenian Smalfjord Formation, which incorporates mostly terrestrial fluvioglacial diamictite and glaciomarine sediments correlated with the Marinoan Glaciation at 635.5 ± 0.6 Ma (Edwards, 1984; Gayer and Rice, 1989; Condon et al., 2005; Rice et al., 2011, 2012; see Bechstädt et al., 2018 for a recent review; further referred to 635 Ma) to the Breidvika Formation of the lower Cambrian Terreneuvian Series (Reading and Walker, 1966; Farmer et al., 1992; Högström et al., 2013; Jensen et al., 2018; Fig. 2). The age range of the Vestertana Group is chemo- and biochronologically constrained to 635–521 Ma (Rice et al., 2011; Jensen et al., 2018). The Vestertana Group sequentially comprises the Smalfjord, Nyborg, Mortensneset, Stáhpogieddi, and Breidvika formations and is 1120–2025 m thick, depending on the lateral facies distribution across the depositional basin (Reading, 1965; Reading and Walker, 1966; Banks et al., 1971; Banks, 1973; Edwards, 1984; Føyn, 1985; Rice et al., 2012). The Mortensneset Formation contains an upper glaciomarine diamictite correlated with the Gaskiers Glaciation at 580 Ma (Halverson et al., 2005; Halverson and Shields-Zhou, 2011; Rice et al., 2011; Pu et al., 2016; see also discussion by Jensen et al., 2018). The two glacigenic formations were previously attributed to the Lower and Upper Tillites of the late Proterozoic Varanger Ice Age, respectively (Føyn, 1985; Nystuen, 1985).

Figure 2.

Ediacaran sedimentary succession on the Digermulen Peninsula showing the stratigraphic distribution of fossils and new mineralized body fossils. In the left column, the isotopic ages of the Marinoan and Gaskiers glaciations at 635 Ma and 580 Ma, respectively, and their time-equivalent diamictites are marked; the base of the Cambrian at 541 Ma, and the age of the topmost Manndrapselva Member at 545 Ma are all published data (Peng et al., 2012; Zhang et al., 2015; Pu et al., 2016; Xiao et al., 2016). Note the Cryogenian age of the Smalfjord Formation at the base of the succession. In the right column, the age of the Ediacara-type impression fossils is chronostratigraphically recognized from their global ranges at 558–555 Ma and at the lower occurrence of only Aspidella (referenced in the text). The age of new mineralized body fossils is interpolated herein at ca. 575 Ma.

The post-glacial Stáhpogieddi and Breidvika formations are marine siliciclastic successions that preserve age-diagnostic fossils (Farmer et al., 1992; Högström et al., 2013; McIlroy and Brasier, 2017; Jensen et al., 2018) and constitute a conformable succession with the underlying Mortensneset Formation (Reading, 1965; Reading and Walker, 1966; Banks et al., 1971; Banks, 1973; Edwards, 1984; Føyn, 1985; Gayer and Rice, 1989; Farmer et al., 1992; Rice et al., 2012; Högström et al., 2013; Rice, 2014; Jensen et al., 2018). Regional unconformities have been recognized at the bases of the Smalfjord and Mortensneset diamictites (Edwards, 1984; Rice et al., 2012; see compilation by Rice, 2014; Fig. 2), and thus have implications for determining the relative age of the Stáhpogieddi Formation.

The Stáhpogieddi Formation is a 530-m-thick succession of alternating mudstones, siltstones, and sandstones, with a few thin calcified siliciclastic beds in the uppermost part, and is undeformed, other than being tilted and dipped at 19–28°W, cleaved, and affected by diagenesis and low-grade thermal overprinting (Dallmeyer et al., 1989; Farmer et al., 1992; Meinhold et al., 2019a, b; Fig. 2). The formation has been subdivided into the Lillevannet, Indreelva, and Manndrapselva members, which reflect transgressive-regressive parasequences within a continuous succession; shallowing events mark the sedimentary parasequences, including two distinct upward-coarsening turbidite sandstone intervals occurring within the Indreelva Member and three within the Manndrapselva Member (Reading, 1965; Reading and Walker, 1966; Banks et al., 1971; Banks, 1973; Siedlecka, 1985; Farmer et al., 1992; Rice, et al., 2012; Högström et al., 2013; Jensen et al., 2018). The boundaries between the members of the Stáhpogieddi Formation are gradual or at sharp lithologic contacts but conformable.

The basalmost Lillevannet Member is a 40-m-thick quartzitic sandstone with cross-bedding and ripple cross-lamination. It is also interbedded with mudstone and occasional conglomerate layers that reflect a transition from fluvial to shallow marine, sub-tidal environments (Reading and Walker, 1966; Banks et al., 1971; Edwards, 1984; Rice et al., 2012).

The overlying Indreelva Member, which has yielded the new fossils, is up to 300-m-thick succession and consists of red and green, thin-bedded mudstone with subsequent alternating siltstone, mudstone, and sandstone beds (Fig. 2). A rapid transition from sub-tidal mudstone into a deeper, sub-wave-base facies occurs in the basal portion of the member. The succession was deposited in quiet water environments on an offshore deeper shelf (characterized by mudstone, siltstone) with two regressive intervals (sandstone) tracking a gradual change into shallow marine shelf conditions (Reading, 1965; Banks et al., 1971; Banks 1973; Edwards, 1984; Rice et al., 2012).

The uppermost Manndrapselva Member is a 190-m-thick succession of alternating coarse-grained quartzitic sandstone interbedded with greywacke and thin-bedded siltstone. It incorporates three shallowing-upward parasequences consisting of fine-grained distal to proximal turbidite sediments deposited on an outer marine shelf; the quartzitic sandstone intervals reflect shallower marine conditions (Reading, 1965; Reading and Walker, 1966; Banks et al., 1971; Högström et al., 2013; Jensen et al., 2018).

The only occurrence of carbonate sediments within the Stáhpogieddi Formation includes a series of discontinuous calcified siliciclastic beds up to 15 cm thick. These are randomly distributed as calcareous lenses and concretions over a 23-m-thick interval of the upper Manndrapselva Member, and are observable over a limited outcrop on the Digermulen Peninsula (Meinhold et al., 2019b). These upper Manndrapselva Member siliciclastic sediments formed during the late Ediacaran and were locally early diagenetically calcified and altered into concretions through the late Cambrian–Ordovician time (Meinhold et al., 2019b). The stratigraphic interval bearing these carbonates lies 40 m below the base of the Cambrian recognized in the succession, is of latest Ediacaran age and a rather insignificant sedimentary component within the 530-m-thick siliciclastic Stáhpogieddi Formation The upper boundary of the Stáhpogieddi Formation with the succeeding Breidvika Formation is lithologically abrupt but without any depositional break (Jensen et al., 2018).

The Stáhpogieddi Formation contains the diagnostic Cambrian-type trace fossils and organic tubular metazoan Sabellidites Yanishevsky, 1926 in its uppermost part, the Ediacara-type impression fossils in the middle, the new body fossils in the lowermost part, and casts and molds of modular Palaeopascichnus macrofossils extending throughout the formation (Farmer et al., 1992; Högström et al., 2013; McIlroy and Brasier, 2017; Jensen et al., 2018). The Ediacaran-Cambrian boundary is recognized in the uppermost Manndrapselva Member (Högström et al., 2013; McIlroy and Brasier, 2017; Jensen et al., 2018; Fig. 2).

Sedimentary continuity within the Stáhpogieddi Formation.—Regionally, in a scale of the marine depositional basin extending across the Digermulen and Varanger peninsulas, the entire succession, including the Mortnesneset Diamictite Formation, the Lillevannet Member, and the succeeding members of Stáhpogieddi Formation, is sedimentologically continuous. This has been examined by numerous studies and synthesized by Rice et al. (2012) and Rice (2014), see also discussion by Jensen et al. (2018). On the Digermulen Peninsula, the contact between the Mortensneset Formation and the Lillevannet Member is lithologically abrupt between diamictite and mudstone or sandstone, but without erosion and is conformable (Reading and Walker, 1966; Edwards, 1984; Gayer and Rice, 1989; Rice et al., 2011, 2012; Rice, 2014; Jensen et al., 2018). This contact was suggested to be a sharp erosional unconformity (McIlroy and Brasier, 2017, p. 355), but without indicating any sedimentary structures for erosion. The conglomerate layers in the middle of the Lillevannet Member, which occur only on the Varanger Peninsula, show a shallowing, regressive interval with channels cutting sub-surface and fluvial influx from the nearby continental margin of a high relief into marine basin, but these layers do not extend farther into shelf on the Digermulen Peninsula (Banks et al., 1971; Edwards, 1984). These conglomerates are not considered to represent an erosional unconformity involving time hiatus, but rather show lateral facies distribution within the Lillevannet Member.

McIlroy and Brasier (2017, p. 353) suggested that a significant depositional hiatus occurred within either the Lillevannet or Indereelva members based on their comparison of the 530-m-thick upper Ediacaran (ca. 580–541 Ma) succession of the Stáhpogieddi Formation with the ∼2500 m thick succession between the Gaskiers Formation and the Cambrian GSSP horizon in southeastern Newfoundland (McIlroy and Brasier, 2017 cited as Dalrymple et al., 2006, but Ichaso et al., 2007 is the correct citation for this data) corresponding to the same time interval. However, this is misleading because the Newfoundland strata (including the Conception and St. John's groups in west Conception Bay, the Avalon Zone) comprise deep-water, siliciclastic-volcanoclastic turbidites, contourites, and volcanic ash beds deposited in marine slope-to-basin settings. Such fore-arc, pull-apart depositional basins within the tectonic subduction zone of a spreading center (Ichaso et al., 2007) typically experience very fast sedimentation rates (Stow, 1986; Leeder, 1994). This clearly conflicts with the Stáhpogieddi Formation, which was otherwise deposited on a passive margin, and therefore is substantially thinner, but continuous with no demonstrable sedimentological unconformities (Farmer et al., 1992; Rice et al., 2012; Rice, 2014; Jensen et al., 2018; Fig. 2). We therefore favor a more feasible depositional comparison with the 500-m-thick upper Ediacaran succession in the Adelaide Rift Complex in South Australia; this fine-grained siliciclastic succession comprises minor dolomites and accumulated on carbonate shelf to delta front with some basinal sedimentation (Grey, 2005, p. 38, fig. 13) and has been bracketed at 580–541 Ma based on isotopic dating of the Acraman ejecta layer versus the Ediacaran-Cambrian boundary (Grey, 2005, p. 54).

The Ediacaran fossil ranges and chronostratigraphy.—The Ediacaran soft-bodied macroscopic impression taxa (some also known from 3-D casts and molds in other occurrences), including Dickinsonia, Swartpuntia, Hiemalora, and Aspidella, and 3-D preserved Palaeopascichnus are recorded in the Stáhpogieddi Formation (Fig. 2). Their first appearance datum (FAD), last appearance datum (LAD), and stratigraphic ranges are established from global records and used for our estimation of the age of the assemblage containing them on the Digermulen Peninsula. Furthermore, this age provides a biochronologic time frame for the new fossils, which underlie ∼130 m below in the rock succession (Fig. 2). The FADs and LADs of individual taxa may be approximately recognized from isotopic dates of the rock levels that are close to their records, and we follow those well-dated ranges.

Biostratigraphic range of Dickinsonia.––The impression fossil total grouping of Dickinsoniomorpha has been constrained to either 559–550 Ma (Fedonkin and Vickers-Rich, 2007; Grazhdankin, 2014) or 556–541 Ma time intervals (Narbonne et al., 2012; Xiao et al., 2016). The nominal genus Dickinsonia ranges between 558–555 Ma (Wood et al., 2019), with the FAD and LAD following isotopic calibrations from the White Sea assemblage in Russia (Martin et al., 2000; Fedonkin and Vickers-Rich, 2007).

Biostratigraphic range of Swartpuntia.––The monotypic genus Swartpuntia is isotopically dated to between 545.1 ± 1 Ma and 543.3 ± 1 Ma and up to 539.4 ± 1 Ma or 538 ± 1 Ma in the Nama Group of Namibia (Narbonne et al., 1997; Vickers-Rich, 2007), and recently updated to 538.99 ± 0.21 Ma (Linnemann et al., 2019). In the White Sea assemblage, Swartpuntia germsi Narbonne, Saylor, and Grotzinger, 1997 is recorded in the middle of the Verkhov Formation, which is isotopically dated to between 558.3 ± 1 Ma and 555.3 ± 0.3 Ma (Martin et al., 2000; Fedonkin and Vickers-Rich, 2007). Swartpuntia cf. S. germsi occurs together with the cloudiniids, which range at 550–541 Ma, in the Wood Canyon Formation of Nevada, USA, and shortly below the appearance of the Cambrian Treptichnus Miller, 1889 trace fossil, thus confirming its uppermost Ediacaran extension (Hagadorn and Waggoner, 2000) as known in Namibia. The Swartpuntia-like molds similarly occur alongside Treptichnus in the lowermost Cambrian Uratanna Formation in the Flinders Ranges of South Australia (Wood et al., 2019) and support a LAD of the genus that accords with the more reliably dated Nama assemblage (Linnemann et al., 2019). The range of Swartpuntia is 558–539 Ma and its FAD coincides with that of Dickinsonia (558–555 Ma).

Biostratigraphic range of Hiemalora.––The range of Hiemalora is within the time interval of 565–550 Ma (Hofmann et al., 2008; Chen et al., 2014; Grazhdankin, 2014). The LAD is within the Shibantan Member of the Dengying Formation in China, whose deposition is bracketed to between 551–541 Ma, with a precisely dated horizon of the formation basal age at 551.09 ± 1.02 Ma (Schmitz, 2012). Chen et al. (2014) alternatively listed a LAD of 550 Ma. The range of Hiemalora could be substantially wider, if one would accept this taxon to be synonymous with Aspidella and both to represent the preservation states of holdfasts of frondose taxa like Charniodiscus Ford, 1958 rather than discrete species (Burzynski and Narbonne, 2015; Burzynski et al., 2017a, b). This synonymy has not yet been supported by the record of Aspidella and Hiemalora being attached to specimens of the same frondose taxon. On the other hand, the connection between Hiemalora as being the holdfast of Primocandelabrum hiemaloranum Hofmann, O'Brien, and King, 2008 is documented in a fossil specimen preserving those two taxa together (Hofmann et al., 2008). Additionally, Primocandelabrum Hofmann et al., 2008 is not attributed to any recognized Ediacaran clade, whereas Charniodiscus is associated with Arboreomorphs (Burzynski and Narbonne, 2015), suggesting that Hiemalora may not be the preservation state of Aspidella.

Biostratigraphic range of Aspidella.––The type species Aspidella terranovica Billings, 1872 has a global stratigraphic range of 579–550 Ma (Grazhdankin, 2014), with slightly narrower of 571–560 Ma range in Avalon Zone type area of SE Newfoundland (Liu et al., 2015; Pu et al., 2016), or 580–560 Ma in the Mackenzie Mountains of NW Canada (Burzynski et al., 2017a). Wood et al. (2019) listed Aspidella within a simplified category of “discs” at 571–540 Ma, but without including Hiemalora and seemingly without its older part in NW Canada and globally. Two specimens of Aspidella from the Mackenzie Mountains among numerous “discs” are reported from Cryogenian strata and suggest extension this taxon's stratigraphic range by ca. 80 Ma (Hofmann et al., 1990; Burzynski et al., 2017a). Regardless of the Cryogenian record and the late Ediacaran acme range of Aspidella at ca. 571–540 Ma, the taxon is less stratigraphically significant. We treat the two taxa, Aspidella and Hiemalora, as individual genera and their ranges are only partially overlapping at 565–550 Ma (i.e., coinciding with the range of Hiemalora).

Biostratigraphic range of Palaeopascichnus.––This modular fossil, also broadly treated within the group of palaeopascichnids, has been taxonomically synonymized with several taxa and its range suggested to span the entire Ediacaran Period (Kolesnikov et al., 2018). The junior synonym Orbisiana Sokolov, 1976, emend. Wan et al., 2014 occurs in the lowermost Ediacaran Lantian Formation of South China, which is estimated at 635–590 Ma (Yuan et al., 2011; Wan et al., 2014, 2016; Wang et al., 2017; Kolesnikov et al., 2018), but with re-evaluation to ca. 620–580 Ma based on correlation with the member II of the Doushantuo Formation (Yuan et al., 2011; Wang et al., 2017), for which the cited age is estimated (Liu and Moczydłowska, 2019). Palaeopascichnus has been identified in Baltica and Avalonia at 565–541 Ma (Grazhdankin and Maslov, 2009, 2015; Jensen et al., 2018) and in Australia at 560 Ma (Haines, 2000; Bowring et al., 2007). The overall global range of paleopascichnids is estimated at 580–540 Ma (Grazhdankin, 2014). We follow the range of Palaeopascichnus to span most of the Ediacaran Period, ca. 620–540 Ma (not all), and the taxon has limited stratigraphic significance.

Based on the current recognition of stratigraphic ranges in isotopically dated sedimentary intervals containing the fossils, the overlapping ranges of Aspidella, Hiemalora, Dickinsonia, and Swartpuntia at 558–555 Ma provide the closest available age that is inferred for the strata recording these taxa in the Stáhpogieddi Formation (Fig. 2).

Age of the Stáhpogieddi Formation and recorded fossils.—The base of the Stáhpogieddi Formation correlates with the top of the Mortensneset Formation diamictite at 580 Ma (Halverson et al., 2005; Halverson and Shields-Zhou, 2011; Rice et al., 2011, 2012; Pu et al., 2016). Detrital zircon grains of Baltican provenance recovered from the medium-grained turbiditic sandstone in the upper Manndrapselva Member of the Tanafjorden area, which extends laterally to the Digermulen Peninsula, have provided ages of 542.1 ± 2.8, 542.4 ± 4.8 and 544.9 ± 5.1 Ma (W. Zhang et al., 2015). The maximum age of this sediment interval is thus 545 Ma (W. Zhang et al., 2015; Fig. 2), which concurs with the 541 Ma age estimate based on trace fossils from the uppermost horizons of the Manndrapselva Member (Högström et al., 2013; Jensen et al., 2018). These ages bracket the Stáhpogieddi Formation to between 580–541 Ma and slightly younger where it extends into the Cambrian (Jensen et al., 2018; Fig. 2). The uppermost Manndrapselva Member records the first appearance of Treptichnus pedum (Seilacher, 1955) and other Cambrian-type trace fossils (Jensen, 1997; Högström et al., 2013; McIlroy and Brasier, 2017; Jensen et al., 2018). Treptichnus pedum is also the index taxon for the base of the Cambrian System at the GSSP in Newfoundland (Narbonne et al., 1987; Brasier et al., 1994; Landing, 1994). U-Pb dates from the Ediacaran-Cambrian transitional strata (Grotzinger et al., 1995; Bowring and Schmitz, 2003; Bowring et al., 2007) have refined the lower Cambrian boundary to 541 ± 0.63 Ma (Peng et al., 2012; further referred to 541 Ma; Fig. 2) or possibly to ca. 539 Ma (Linnemann et al., 2019). The latter age of the stratigraphic level suggested for the Ediacaran-Cambrian boundary in the Nama Group of Namibia (in a composite section) is within 538.99 ± 0.21 Ma to 538.58 ± 0.19 Ma (Linnemann et al., 2019; further referred to 539 Ma). Before the new age may be accepted by the International Commission of Stratigraphy (ICS), we will use the age at 541 Ma for the lower Cambrian boundary and the age at 539 Ma for the ranges of certain fossil taxa recognized in Namibia (Linnemann et al., 2019) and extending from the Ediacaran.

Classic Ediacaran impression fossils, including Aspidella, Hiemalora, Dickinsonia, and Swartpuntia, all occur in the middle Indreelva Member of the Stáhpogieddi Formation (Farmer et al., 1992; Högström et al., 2013, 2017; Høyberget et al., 2017; new identification of Dickinsonia and Swartpuntia; Fig. 2) and provide a demonstrable age range of 558–555 Ma (Fig. 2), and certainly not younger relying on the LAD of Dickinsonia. The FAD of Dickinsonia and Swartpuntia, which is isotopically dated (Martin et al., 2000; Fedonkin and Vickers-Rich, 2007), constrains the maximum age. The time interval 558–555 Ma is the biochronologically estimated age of this stratigraphic horizon in the Stáhpogieddi Formation The lowermost occurrence of Aspidella in the succession (Högström et al., 2017; Høyberget et al., 2017) may be close to this taxon acme occurrence known from 571 Ma (Wood et al., 2019), and is shortly above the new fossils (Fig. 2).

Our new Indreelva Member assemblage co-occurs with Palaeopascichnus, which prompted a specific age correlation for this interval at ca. 565 Ma (Jensen et al., 2018). As mentioned, this taxon has wide stratigraphic range spanning 620–540 Ma and we decline from using this taxon for age determination in the Stáhpogieddi Formation.

The age of the new fossil assemblage may also be estimated from its relative stratigraphic position within the Stáhpogieddi Formation and calculated rate of deposition within the succession. Examples and arguments exist that the rates of deposition may be very variable and change in marine environments. Variables that are considered in estimating the rate of deposition include record fidelity, complex and incomplete nature of stratigraphic record owing to changing dynamics of the depositional system, variation in rate of deposition for elements in the basin, time and length scales, and stochastic processes (“noise”) (Tipper, 2016; Davies et al., 2019; Holbrook and Miall, 2020; Straub et al., 2020). However, in long timescales, the rates of deposition are statistically buffered over periodicity (Straub et al., 2020) and cyclic variability sums produce the average rate of deposition, which can be plausibly estimated from measured parameters of sediment thickness over the time span of the section (Holbrook and Miall, 2020; Straub et al., 2020). The accuracy of average rates of deposition depends on the vertical thickness and longer durations of time (Holbrook and Miall, 2020). The measured rates of deposition are balanced among states of deposition, erosion, and stasis (Davies et al., 2019) and are normalized for longer measurement duration (Straub et al., 2020). Despite the limitations of the incomplete records and variation of rates of deposition, the estimation of rate of deposition is worthwhile and possible (Tipper, 2016).

Depositional environments of the Stáhpogieddi Formation in the lowland basin were relatively stable on the passive continental margin, and cyclic shallowing and deepening events due to sea-level changes were of low amplitudes within shelf settings without tectonic disturbance or erosional unconformities (Gayer and Rice, 1989; Rice et al., 2012; Rice, 2014). This formation consists of transgressive-regressive parasequences, in which the rate of deposition between alternating units of mudstone, siltstone, and sandstone certainly varied in small-scale intervals. However, in a substantial thickness (530 m) of sedimentologically continuous succession it is rather uniform. In a cyclic repetition of fine- to coarser-grained siliciclastic rocks, the average rate of deposition is normalized and similar throughout the succession. Plausibly therefore, the long deposition time within 39 Ma (580–541 Ma and slightly younger; Fig. 2), together with its continuous sedimentation at similar average rate of 13.59 m/Ma, allows the interpolation of the ages within various portions of the formation. Most likely, the basal Indreelva Member, containing the new body fossils and overlying 40–42 m above the Mortensneset Formation, can be estimated at as little as 3 Ma, and rounded to as much as 5 Ma younger than the base of the formation assuming a degree of error (Fig. 2). Notably, this estimated ca. 575 Ma age limit is consistent with the FAD for typical Ediacara-type impression fossils in the middle Indreelva Member, which occur ∼130 m above, and their overlapping ranges at 558–555 Ma. Calculating independently from the rate of deposition in the narrower sediment interval limited by the age-controlled horizons between 558–555 Ma and 580 Ma (Fig. 2), this 172 m of continuously deposited sediments in 22–25 Ma allow estimation of the time horizon of new fossils to be 5.46–6.10 Ma younger than 580 Ma. This conforms to the age estimated from the entire succession at ca. 575 Ma. Therefore, the average rate of deposition may be considered here as reliable.

Materials and methods

Twenty-two individual fossil specimens were collected from natural exposures along bedding plains in mudstone of the lowermost Indreelva Member in the Stáhpogieddi Formation (Fig. 2). The host sediment laminae bent plastically around the skeletons, which were slightly distorted by tectonic compression but retained hollow lumen or were mostly post-mortem infilled with the host sediment. Three different morphotypes are recognized at the rank of monotypic genera (of single species) and taxonomically described.

The fossils were studied under a reflected light microscope (RLM), a transmitted and polarized light microscope (PLM), an environmental scanning-transmission electron microscope (STEM) equipped with energy-dispersive X-ray detector for energy dispersive spectroscopy (STEM EDS), and a field emission electron probe microanalyzer (EMP) also equipped with EDS detector. Three-dimensional rendering and visualization of internal structures was also undertaken using a Nikon H225 ST Micro-Computed Tomography (CT) at the University of Oslo (Supplementary Data). Petrographic thin-sectioning was also carried out with geochemical and acid tests, as well as STEM EDS, EMP EDS, and Laser-Raman spectroscopy to determine mineralization (Supplementary Data). The mineral composition of the surrounding mudstone matrix was examined using Laser-Raman spectroscopy and X-ray diffraction (XRD; Supplementary Data).

Figure 3.

Sketch-drawings and morphologic terminology of new fossil taxa drawn proportionally to their dimensions; interrupted lines define the lumen within walls. (1, 2) Holotype of Anulitubus n. gen. Moczydłowska in Moczydłowska et al.; (3) holotype of Coniculus n. gen. Moczydłowska in Moczydłowska et al.; (4) representative specimen of Fistula n. gen. Moczydłowska in Moczydłowska et al. Thick black outline of aperture and body surface in (3, 4) marks the wall outer veneer layer.

Repository and institutional abbreviation.—The material is reposited in the paleontological collections of the Museum of Evolution at Uppsala University (PMU). The collection is assigned as the Digermulen Peninsula fossils, Finnmark, Norway, and carry the catalogue numbers PMU 34736–34757, in which the last two digits refer to the individual specimen number.

Systematic paleontology

Terminology.—Although the higher taxonomic affinities of the fossils described herein are unresolved, they are hierarchically assigned to Metazoa and Eumetazoa based on the eukaryote classification of Adl et al. (2012, 2019) and dos Reis et al. (2015) (see also Cunningham et al., 2017). The term “animal” is used as an informal equivalent of metazoans in accordance with Blair (2009). Morphologic terminology (Fig. 3) follows the recommendations of the International Zoological Code of Nomenclature ( www.iczn.org). For example, we use “tube,” “tubular wall,” or “tubular skeleton” to describe cylindrical structures and “conus,” “conical wall,” or “conical skeleton” for conical structures. We have therefore avoided the otherwise ambiguous terms “conotubular” or “conical tube.” Similarly, we prefer “funnel-shaped segments or elements” (sensu Zhuravlev et al., 2012) instead of “funnel-shaped cylinders,” and have refrained from using “apex,” “apical end,” or “basal apex” (sensu Cai et al., 2010, 2011, 2017; Cortijo et al., 2015) because of possible conflicting interpretation as embryonic segments that may have been embedded in the substrate. Finally, the term “apical flaring” was consistently used for defining portions of segments in Cloudina (Zhuravlev et al., 2012), which are oriented towards the distal end of the body.

Type locality.—All the taxa described in this study are from the coastal rock exposure at Árasulluokta Cove between the Árasulluokta and Manndrapselva rivulets on the southeastern coast of Tanafjorden, Digermulen Penninsula, Finnmark County, northeastern Norway. Mapsheet Langfjorden 223611, coordinates 420 303.

Type stratum.—All taxa described in this study are from the red-brown, thin-bedded slaty mudstone in the lowermost part of the Indreelva Member, Stáhpogieddi Formation. The type horizon is ∼40–42 m above the base of the Stáhpogieddi Formation coinciding with the top of the underlying Mortensneset Diamictite Formation; Ediacaran ca. 575 Ma.

Figure 4.

Macroscopic, mineralized tubular annulate Anulitubus formosus n. gen. n. sp. Moczydłowska in Moczydłowska et al., holotype specimen PMU 34736. Images from RLM (1–4) showing (1) front side view, (2) opposite side view, (3, 4) apertural views with free lumen and visible incomplete septum. Images from STEM (5–7) showing (5) front side view, (6) opposite side, and (7) aperture. Scale bars = 1 mm.

Figure 5.

Macroscopic, mineralized tubular annulate Anulitubus formosus n. gen. n. sp. Moczydłowska in Moczydłowska et al. holotype specimen PMU 34736. Images from RLM in (1), and CT in (2–7); (1, 2) front view with three annuli, two grooves visible on the outer surface, and one groove on the inner surface within the lumen; (3) apertural view showing incomplete septum with central opening; (4) front view with lower portion that is embedded in the sediment (marked in gray color), showing the first annular segment together with the basal cup-shaped element; (5) mirror orientation of image in (4) with sediment entombing the specimen, but clearly showing the central knob at the tube base; (6, 7) mirror images of the cup-shaped element displaying the concave surface facing the tube lumen and radial symmetry. Scale bars in (1–7) are 1 mm.

Figure 6.

Wall composition of Anulitubus formosus n. gen. n. sp. Moczydłowska in Moczydłowska et al., holotype specimen PMU 34736, shown by elemental content analyzed by STEM EDS on the wall surface in (1) and the incomplete septum in (2). Measured points are marked by red asterisks. The point spectra are characteristic of silica forming the wall with high content of element carbon indigenous to the wall and remnant of organic matter.

Figure 7.

The wall composition of Anulitubus formosus n. gen. n. sp. Moczydłowska in Moczydłowska et al., holotype specimen PMU 34736, shown by elemental mapping analyzed by STEM EDS on the wall aperture (1–5) and the incomplete septum surface (6). Asterisk in (5) indicates the measurement site of incomplete septum in (6). The intensity and homogenous distribution of Si and O are consistent with siliceous wall and incomplete septum, with some accessory elements derived from the host sediment. Scale bar in 1 for all images = 1 mm.

Figure 8.

The wall composition of Anulitubus formosus n. gen. n. sp. Moczydłowska in Moczydłowska et al. revealed by the Laser-Raman spectra acquired on the wall surface in several points and showing the signatures of quartz and anatase minerals, and spectral disordered D-band and graphitic G-band of carbonaceous material that is admixed in the mineralized wall. Holotype specimen PMU 34736 illustrated in Figure 4. The specimen wall is composed of quartz that is recrystallized from the original amorphous opaline silica into microcrystalline phase during diagenesis. The anatase derived from the burial environment and is a contaminant from microcrystals attached on the specimen wall surface. The green and orange color tick marks are the positions of strong Raman peaks of anatase (TiO₂, RRUFF ID: R060277) and quartz (SiO₂, RRUFF ID: R040031), respectively.

Figure 9.

Macroscopic, mineralized conical Coniculus elegantis n. gen. n. sp. Moczydłowska in Moczydłowska et al., holotype specimen PMU 34737. (1, 2) Images from RLM in a front side view; (3) opposite side view; and (4, 5) apertural view. Distinct outer veneer layer of the wall is seen as a sharp outline in the apertural, circular cross sections (4, 5) and can be chipped off at the edge (1, 5). The thicker inner wall layer is seen around the aperture (1, 4, 5) and the lumen is infilled with sediment in the lower portion (5). The end of cone is entombed in the sediment at approximate level marked by white arrow in (2) and recognized by the relief and by comparison with CT images. Scale bars for (1–3) = 6 mm, and for (4, 5) = 1 mm.

Figure 10.

Coniculus elegantis n. gen. n. sp. Moczydłowska in Moczydłowska et al. CT images in a front side view of the specimen (1) and a mirror and turned side view (2) showing constant wall thickness and sharp edges, and the lumen lower part infilled with the sediment (yellow color in images). Scale bar = 5 mm.

Figure 11.

Coniculus elegantis n. gen. n. sp. Moczydłowska in Moczydłowska et al. Images from RLM (1–5) and STEM (6). Conical wall with tapered closed proximal ends that are entombed in sediment (1, 2) or broken apart (3), and apertural ends showing sharp-edged outer veneer layer (4, 5). Lumen is infilled with host sediment (4, 5). The surface of specimens may be partly coated by the sediment aluminosilicates, including chlorite (6, enlarged from 1, right side). White arrow in (2) indicates position of proximal end. Specimens PMU 34738 (1), PMU 34740 (2), PMU 34739 (3–5). Scale bars in (1–5) = 2 mm, in (6) = 3 µm.

Figure 12.

The wall composition of the Coniculus elegantis n. gen. n. sp. Moczydłowska in Moczydłowska et al., holotype specimen PMU 3434737 analyzed by STEM EDS on the cross section of the wall inner layer (1) and wall surface veneer layer (2). The highest elemental content of Si and O is indicative of siliceous composition with admixture of C and N from organic matrix. Red asterisks mark the measurement points on the insets of STEM images. Scale bars in STEM images = 1 mm.

Figure 13.

The wall composition of the Coniculus elegantis n. gen. n. sp. Moczydłowska in Moczydłowska et al., holotype specimen PMU 34737, illustrated in Figure 9, shown by the Laser-Raman point spectra from the wall surface in the distal and proximal portions and characteristic signatures of quartz and anatase minerals. The wall was siliceous and anatase derived from the burial environment. The blue and orange tick marks are the positions of strong Raman peaks of anatase (TiO₂, RRUFF ID: R060277) and quartz (SiO₂, RRUFF ID: R040031), respectively.

Figure 14.

Petrographic thin section images of Coniculus n. gen. Moczydłowska in Moczydłowska et al. (PMU 34739), (1) parallel light, (2) crossed polarizers, and Fistula n. gen. Moczydłowska in Moczydłowska et al. (PMU 34744), (3, 5) parallel light and (4, 6) crossed polarizers, in longitudinal sections showing two layers of the wall and internal lumen infilled by host sediment. The wall layers are composed of microcrystalline phases dominated by silica (quartz, opal) and minor sheet silicates. The infilling of lumen is penetrated by coarse chalcedony veins. The outer veneer layer (dark line) is thin and compact and the inner layer (light gray band) is wide and microcrystalline. The aperture of Coniculus n. gen. Moczydłowska in Moczydłowska et al. (1, 2) shown in the upper edge of the specimen is infilled by elongated crystals of chalcedony. In Fistula n. gen. Moczydłowska in Moczydłowska et al. (3, 4) lumen is shown at the right side, to the left is wall with internal and veneer layers, and a chip of host sediment attached on the surface, respectively. The opposite side of specimen in (5, 6), the lumen with chalcedony veins is at left side, then inner layer and veneer layers to the right. Numbers with arrows in yellow color in (6) refer to EMP EDS analyses shown in Figure 15.

Figure 15.

Analyses of elemental composition and mineralogy of Fistula n. gen. Moczydłowska in Moczydłowska et al. wall and lumen infilling minerals by a field emission electron probe microanalyzer equipped with energy-dispersive spectroscopy detector (EMP EDS) (1–6). (1–4) Minerals infilling the fossil lumen are diagenetic and comprise the first generation of Ca-carbonate (calcite) associated with Mn-oxide (1, 2) and silica that are, in turn, partially replaced by the second generation of silica (chalcedony) (4), Ca-carbonate (calcite) and barite (3). (5, 6) The fossil wall is composed of pure silica (5) with an admixture of a Mg-Fe aluminosilicate, probably mixed-layer clay mineral (5, 6).

Figure 16.

Macroscopic, mineralized cylindrical and segmented Fistula crenulata n. gen. n. sp. Moczydłowska in Moczydłowska et al. All are RLM images. (1, 2) Holotype specimen PMU 34743, side view in (1) with a sediment chip at left lower portion, and apertural view in (2) showing the sharp edge of the wall outer veneer layer and the thick inner layer around free lumen. (3, 5) Specimen PMU 34744, showing the faint crenulation on wall surface and two wall layers; the sediment chip is attached to the wall surface on the right side and is clearly distinguishable from the veneer layer; white arrows point to the surface wrinkles. (4, 6) Specimen PMU 34741 showing the groove between the segments on the wall surface (6) and the two wall layers in cross section (4); the sediment is attached to the left side of specimen (black triangles) and easily recognizable from the specimen surface by peeling off (4, 6); black asterisks indicate outer veneer layer, and black dots the wall inner layer. Specimen cross sections demonstrate the radial symmetry of cylindrical exoskeletons. Scale bars = 5 mm.

Figure 17.

Macroscopic, mineralized, cylindrical, and segmented Fistula crenulata n. gen. n. sp. Moczydłowska in Moczydłowska et al. All are RLM images with the exception of CT image in (6). (1) Specimen PMU 34749, in major part exposed from the reddish sediment; the largest specimen with surface crenulation (faint wrinkles indicated by white arrows); circular cross section on the left end; covered by diagenetically crystallized quartz (white color). (2, 4) Paratype specimen PMU 34748 lying on the bedding plane and demonstrating longitudinal polarity; wall segments delimited by incised sutures (4); proximal end is rounded (2, bottom), and aperture is open (2, upper part). (3, 5, 6) Specimen PMU 34742, showing cylindrical exoskeleton with segments (5, 6), circular cross section (3), and constant wall thickness at fractures between the segments (6); free lumen is infilled with sediment and visible under the fractured wall (6) and at the apertural end covered by diagenetic quartz (5, left end, white in color). (7, 8) Specimen PMU 34745, undeformed and circular in cross section demonstrating its radial symmetry (8), lying on the bedding plane and partly exposed from the host sediment, with sediment layers bent around specimen (8); both wall layers are visible and lumen is infilled with sediment. Scale bars = 5 mm for all images.

Figure 18.

CT images of the Fistula crenulata n. gen. n. sp. Moczydłowska in Moczydłowska et al., holotype specimen PMU 34743, revealing the internal features of the tube wall and its two layers. Sediment partly covering the specimen in front view is gray in color in (1, 2), as well as infilling the breakage of the wall, which is blue in color (4). The wall is marked by artificial colors: outer veneer layer is gray in (3, 4); wall inner layer is green in (2–4). The knob-like structure is visible at the tube bottom (3, 4) and broken portion of the veneer layer on the left side in (3). The central lumen infilled with the sediment appears as vertically stacked yellow bands in (4). Scale bars equal 5 mm.

Figure 19.

Wall composition of Fistula crenulata n. gen. n. sp. Moczydłowska in Moczydłowska et al. revealed by the Laser-Raman spectra acquired on the wall surface in several points on different portions of specimen PMU 34752, and showing the signatures of quartz and anatase minerals. The wall is siliceous in the quartz microcrystalline phase, and anatase is a contaminant from the sediment. The blue and orange tick marks are the positions of strong Laser-Raman peaks of anatase (TiO₂, RRUFF ID: R060277) and quartz (SiO₂, RRUFF ID: R040031), respectively.

Figure 20.

The Laser-Raman spectra, for comparison of mineralogical composition and differentiation of primary mineralogy of the fossil, acquired (1) on the wall of Fistula crenulata n. gen. n. sp. Moczydłowska in Moczydłowska et al., specimen PMU 34750, (2) on the entombing sediment around the specimen, and (3) on another sample of host sediment a few meters above the specimen. The strong quartz spectra of the wall (N = 4) is different than the weak spectra in the entombing (N = 3) and overlying sediments (N = 5), showing the primary siliceous skeleton of the fossil in quartz microcrystalline phase. The signatures of anatase and hematite minerals are from the sediment and contaminant in the fossil. The carbonaceous material in disordered D-band indicates the organic matter admixture in the fossil wall and its presence as disseminated organic matter in the sediments. The green and orange tick marks are the positions of strong Laser-Raman peaks of anatase (TiO₂, RRUFF ID: R060277) and quartz (SiO₂, RRUFF ID: R040031), respectively; blue tick mark is for hematite.

Figure 21.

Mineralogic composition of the sediment entombing the fossils examined by X-ray diffraction (XRD). Red, thin-bedded mudstone is composed of detrital quartz, muscovite, and feldspars. Other identified minerals are diagenetic.

Metazoan affinity and taxonomic comparisons

Phenotypically, the new taxa described manifest a range of features consistent with metazoan affinity: macroscopic sizes (0.5–7.0 cm), cylindrical and conical shapes, annulated or segmented structure with smooth to crenulate surface ornamentation, single- or double-layered skeletal wall, radial symmetry, and longitudinal polarity, including closed proximal versus open apertural (oral) ends. There was also clearly staged skeletal accretion and increase in diameter from a cup or knob-like proximal initial skeleton in Anulitubus n. gen and Fistula n. gen., or a simple cone in Coniculus n. gen. The variation in skeletal morphology and complexity observed between these taxa is comparable with modern marine benthic organisms, such as cnidarians and lophophorates, as well as vermiform animals and tubular annelids (especially polychaetes), or even bryozoans. However, individual morphological elements observed in new taxa, their overall body habit, and some structural similarities exist with Ediacaran body fossils such as Cloudina, Sinotubulites, Conotubus, Multiconotubus, Gaojiashania, and Shaanxilithes. All the latter taxa possessed either organic integument or mineralized skeletons, and have been interpreted as possible representatives of currently extant phyla (Grant, 1990; Cai et al., 2010,, 2014, 2015; Cortijo et al., 2010, 2015; Zhuravlev et al., 2012, 2015; Wood et al., 2017, 2019; Schiffbauer et al., 2020; Shore et al., 2020; Yang et al., 2020).

Cloudina has nested funnel-like elements with high Mg-calcite content (Grant, 1990; Zhuravlev et al., 2012, 2015; Cai et al., 2014; Cortijo et al., 2015; Wood et al., 2017), as preserved in carbonate rocks, and was interpreted as biomineralized exoskeleton from the total-group cnidarians, and thus crown-eumetazoans and non-bilaterian (Wood et al., 2019; Shore et al., 2020). This was based on a combination of morphological characteristics, growth and reproduction patterns, as well as microstructure of the tube, and inferred paleoecology. Yet, Cloudina has also been compared with serpulid worms that are polychaetes in the Phylum Annelida (Hua et al., 2005; Cai et al., 2014), but this affiliation has been questioned and Cloudina was considered to be a cnidarian (Zhuravlev et al., 2012; Warren et al., 2017). This was concluded because of the microstructure of Cloudina and the pattern of budding in asexual reproduction that are characteristic of cnidarians (Zhuravlev et al., 2012). The multiple (polytomous) branching of tube and attachment structures in Cloudina, and cloudinomorphs in general, were supporting the cnidarian and non-bilaterian affinity (Shore et al., 2020). Alternatively, Cloudina and new cloudinid Zuunia have been further interpreted to be originally organic and non-mineralizing stem-group annelids (Yang et al., 2020). Other cloudinomorphs, exemplified by Saarina Sokolov, 1965, emend. Selly et al., 2020 and Costatubus Selly et al., 2020, that were inferred to have been organic in original composition and preserving internal cylinders representing guts have been placed among the stem-annelids, and thus bilaterians (Schiffbauer et al., 2020; Selly et al., 2020). These opposing interpretations may be reconciled if cloudinomorphs show convergent morphology of funnel-stacked elements in tubes and represent polyphyletic origins (Shore et al., 2020). With regards to the new genus Anulitubus, the presence of morphological elements of segments, although ring-like and not funnel-like in the tubular body, is demonstrably a metazoan feature.

Sinotubulites is a multilayered cylindrical fossil interpreted to have been a biomineralizing organism with a probable animal affinity (M. Chen et al., 1981; Z. Chen et al., 2008; Cai et al., 2015) or of uncertain affinity (Wood et al., 2017). Its morphological similarity to organic-walled Corumbella has been recognized, and the latter taxon inferred to be a scyphozoan cnidarian (Van Iten et al., 2014), with Wood et al. (2019) positing that Sinotubulites is a mineralized counterpart of Corumbella Hahn et al., 1982, thus supposedly cnidarian. The double-layered wall in new conical Coniculus n. gen. and cylindrical Fistula n. gen. is present and is consistent with metazoan affinities.

Conotubus is tapering along its length and has nested funnel-shaped elements, and was grouped with Cloudina and Sinotubulites to constitute the family Cloudinidae (Hua et al., 2007; Cai et al., 2011), and thus related to cnidarians as Cloudina was recognized (Wood et al., 2017, 2019; Shore et al., 2020), or Cloudina as newly attributed to stem-annelids (Yang et al., 2020). Regardless of the closer phylogenetic relationships, Conotubus is accepted to be metazoan. Comparably, the overall conical body habit and wall layering structure of Coniculus n. gen., although without segmentation, is shared with metazoans.

Multiconotubus has a multilayered, cone-in-cone wall structure and is diagnosed as a conical biomineralizing organism with “apical” end closed and open aperture (Cai et al., 2017). The fossil is 3-D preserved and diagenetically phosphatized (Cai et al., 2017, 2019), but its primary mineralogy and possible affinity have not yet been determined. This taxon resembles Coniculus n. gen. morphologically by conical shape and layered wall, and both are considered metazoans.

Gaojiashania is preserved as carbonaceous compressions of tubes with characteristic circular rings (Cai et al., 2010)). This taxon has been variously interpreted as worm (Hua et al., 2004), or tubular metazoan in general (Weber et al., 2007). The segmented flexible tube structure has been compared specifically with modern siboglinid polychaetes (Cai et al., 2013). The fossil was originally organic or lightly mineralized (Hua et al., 2004; Cai et al., 2010,, 2012, 2013; Smith et al., 2016), although without indication of its primary mineralogy.

The segments of the tubular body in Anulitubus n. gen. resemble the morphology of Gaojiashania, which is however non-mineralized.

Shaanxilithes is a ribbon-shaped and annulated or corrugated compression fossil (Hua et al., 2004; Shen et al., 2007; Cai et al., 2010)) that may be metazoan in origin (Tarhan et al., 2014). Its organic composition has been established by environmental SEM EDS spectra and elemental mapping and Laser-Raman spectroscopy from specimens preserved in a variety of sediments, including dolostone, siltstone, and phosphatic shale (Meyer et al., 2012; Tarhan et al., 2014). This taxon body annulation may resemble the surface of Fistula n. gen.

In summary, the terminal Ediacaran body fossils are recognized as metazoans based on their morphology, symmetry, wall construction and structure, and inferred benthic mode of life and paleoecology (Cai et al., 2011, 2012; Wood, 2011; Zhuravlev et al., 2012, 2015; Wood and Curtis, 2015; Penny et al., 2017; Wood et al., 2017, 2019; Schiffbauer et al., 2020; Shore et al., 2020; Yang et al., 2020). By the same characteristics of having complex morphology, comparable individual features, radial symmetry and polarity, and structure of mineralized wall, the new taxa represent a metazoan grade of organisms, as was their mode of life within the benthos.

Skeletal biomineralization

We considered a potential non-mineralized organic construction of the new fossils' walls or their diagenetic mineralization versus primary biomineralization. As aforementioned under each taxon, the lack of even accessory elements Ca or P in the walls and the host sediment excluded potential diagenetic replacement by silica of calcareous or phosphatic skeletons. The only exception is the diagenetically induced, secondary calcareous mineral in the infilling of lumen of Fistula n. gen. (Figs. 14, 15). However, this is clearly distinguished from the wall mineralogy, which is siliceous with an admixture of aluminosilicate clay mineral. There is no record of any significant pyrite crystallization in the fossils, as it would be in taphonomic pyritization of organic walls by bacterial decay (seen in SEM images, but not illustrated). Most significant is the very different degree of host-rock early diagenetic silicification or admixture of silicate clay minerals, which is low and not exceeding the common content of SiO₂ in typical siliciclastic rocks. To the contrary, as shown by the series of Raman spectroscopic, petrographic, and SEM- EMP EDS analyses, the fossils' walls are siliceous and differentiated from the sediment (Figs. 14, 15, 21), thus supporting primary biomineralization. The diagenetic history of the succession that contains the fossils and in a regional scale provides no evidence for the diagenetic silicification of strata and taphonomic alteration of fossils. The lesson is learned from the interpretations of biomineralized versus primary organic construction of Cloudina and some other cloudinomorphs, which were thought to be the earliest metazoans with biomineralized calcareous skeletons (Grant, 1990; Zhuravlev et al., 2012; Cai et al., 2014; Cortijo et al., 2015; Wood et al., 2017) and turned out to be originally organic (Schiffbauer et al., 2020; Yang et al., 2020). Expectedly, the next finds of described fossil taxa will provide additional preservation modes to test the interpretation favored here.

Our microscopic, petrographic, spectroscopic, and geochemical analyses all infer the presence of primary biomineralized silica in the skeletal walls of the new Digermulen fossils. The additional supporting evidence is also the lack or only accessory admixture of digenetic silicate clay minerals in the host matrix, as well as the absence of any chert nodules, intrusive siliceous layers, or veins. There is likewise no indication of mineral dissolution (mobilization), solution migration, or recrystallization (Dallmeyer et al., 1989; Rice et al., 1989; Meinhold et al., 2019a), and the fossils themselves are not diagenetically replaced or encrusted.

Biomineralization is biologically induced or organic matrix-mediated mineralization (Lowenstam, 1981; Lowenstam and Weiner, 1989), and the first type is characteristic of single-celled organisms and the second of eukaryotes (Veis, 2003). In forming of the mineral phase by organisms, the mineral is sequestered from the aqueous environment depending on temperature, pH, water chemistry, and ion saturation (Veis, 2003). The process of silicification in organisms always proceeds in confined spaces in which organic membranes or extracellular matrices govern the precipitation of biosilicas (Perry, 2003). The precipitated biosilicas are amorphous and silicified organisms have complex structural features (Perry, 2003). The complexity of walls containing organic matter in studied taxa is in accord with the inferred biomineralization, which is suggested from the chemical elemental (STEM- EMP EDS) and primary mineral (petrography, Laser-Raman signatures) composition of the wall.

Laser-Raman spectra show markedly more intense quartz peaks in the fossils (a typical diagenetic transformation phase of originally precipitated opaline silica) versus host sediment samples. These spectra indicate that the fossils were silicified before the sediment was lithified and its detrital quartz grains and cement could be diagenetically altered and affect the buried fossils (Fig. 19; Supplementary Data). Differentiation between the fossil versus host sediment mineral signatures and intensity (mineral concentration) if containing the same mineral, allows recognition of the fossil primary composition (Cai et al., 2012;; Meyer et al., 2014; Schiffbauer et al., 2014; Muscente et al., 2015; Shang et al., 2018).

These geochemical-spectroscopic methods of determining the material composition have been applied in studies of both mineralized and organic fossils of Ediacaran and other ages and for distinguishing in-life from taphonomic mineralization (Schopf et al., 2005; Cai et al., 2012;; Meyer et al., 2012, 2014; Moczydłowska et al., 2014; Schiffbauer et al., 2014; Tarhan et al., 2014; Muscente et al., 2015; Pruss et al., 2015, 2018; Shang et al., 2018). Laser-Raman spectroscopy is a nondestructive optical method that provides the information to identify the chemical composition of fossils and allows their mineral identification (Kudryavtsev et al., 2001; Schopf et al., 2002, 2005; Wacey, 2009; Olcott Marshall and Marshall, 2015). In the present case study, this method supports biomineralization, as does the petrography.

Independently of geochemical analyses, we thoroughly reviewed the diagenetic history of the host sediment in respect to fossil taphonomy and mode of preservation. The tectonic and diagenetic changes associated with the Caledonide orogeny had relatively limited influence on the studied sedimentary succession (Dallmeyer et al., 1989; Rice et al., 1989; Meinhold et al., 2019a) and the constituent fossils. Mineral composition of shales by XRD analysis determined largely quartz and muscovite, with subordinate chlorite, plagioclase, and hematite. The detrital quartz grains are modified to elongate shape produced by cleavage-forming processes associated with low-grade metamorphism and tectonothermal activity, age-constrained to 440.2 ± 9.4 Ma (Dallmeyer et al., 1989). This tectonothermal alteration at the stage of diagenesis transitional to anchizone (∼200° C), together with pressure-related dissolution of micas but without mobilizing silica, precluded silicification of the sediment. There is no evidence of recrystallization or penetrative migration of any mineral solutions. These conditions were key for preserving the primary mineralogy of the fossils and preventing its replication. The shale from the upper Indreelva Member analyzed by XRD consists of 71% phyllosilicates (mainly illite), 18% quartz, 10% feldspar, and 1% other minerals (Meinhold et al., 2019a). This analysis shows a relatively low content of quartz and does not support the possibility of syngenetic or diagenetic silicification. Our XRD study of laminated mudstone entombing the new fossils in the lowermost Indreelva Member also documents the presence of quartz, kaolinite, muscovite, feldspars (K-feldspar), hematite, leucite, and smectite, and possibly a zeolite group of minerals (Fig. 21; Supplementary Data). This composition is typical of fine-grained siliciclastic rock comprised of detrital, diagenetic, and weathered aluminosilicates (phyllosilicates), and detrital and authigenic quartz (Supplementary Data).

The mineralogical and diagenetic features of the host sediment are congruent with the interpretation that the new fossils preserved primary mineralogy. Comparably, the organic-walled metazoan Sabellidites, macroscopic filamentous vendotaenids, and microscopic unicellular acritarchs preserved in the formation (Farmer et al., 1992; Högström et al., 2013; Jensen et al., 2018; unpublished data on acritarchs) remained organically preserved, thus independently attesting against the sediment syngenetic or diagenetic silicification. There could not be selective permineralization of one type of fossil. The casts and molds of co-occurring Palaeopascichnus show no sign of diagenetic silicification or other re-mineralization (Jensen et al., 2018), and this state of preservation is in contrast to, and also supports, the primary mineralogy of the new fossils.

Diagenetic to the lowermost metamorphic grade of alteration and without remineralization or recrystallization affected the sediment long after (135 and 175 Ma) the studied organisms were buried ca. 575 Ma. They were originally resistant, three-dimensional, and mineralized by primary silica. These organisms were the earliest silica biomineralizing macroscopic organisms that formed skeletons.

Environments facilitating silica biomineralization

Accepting our age estimates, biogenic skeletal silicification would appear to have been the first step in animal biomineralization at ca. 575 Ma. Previously, aragonite-calcite biomineralization was inferred to appear at ca. 550 Ma (Wood et al., 2017), with phosphatization emerging shortly after 541 Ma (Erwin et al., 2011; Erwin and Valentine, 2013). However, the nominal taxon Cloudina and some other cloudinomorphs, among those animals interpreted to biomineralize ca. 550 Ma, have been revised as organic and non-biomineralized but taphonomically altered into various modes such as phosphatic, organic/pyritic, calcareous, limonitic/pyritic, and siliceous preservations (Schiffbauer et al., 2020; Yang et al., 2020). Namacalathus Grotzinger et al., 2000 and Namapoikia Wood et al., 2002 remain as possible calcareous skeletal organisms in the terminal Ediacaran (Grotzinger et al., 2000; Wood et al., 2019). The initial diversification of skeletonized animals then considered has been variously attributed to changing oceanic chemistry, redox conditions, and predation (Bengtson, 2004; Wood, 2011), as well as based on the record of calcareous fossils (Cui et al., 2016; Schiffbauer et al., 2016; Bowyer et al., 2017; Wood et al., 2017, 2019). No predation is evident at that time or in the Ediacaran Period (Narbonne, 2005; Gibson et al., 2019). Silica biomineralization has been posited in sponge spicules (Sperling et al., 2010), but not yet considered a factor in integument of metazoan body fossils in this developmental transition, but it might also had been environmentally triggered by the ready availability of silica. The concentration of silica in cold seawater is higher, facilitating sequestration into siliceous skeletons (Perry, 2003), and such conditions existed during the Gaskiers glaciation (580 Ma) and shortly after, where biomediated silica uptake may have been more efficient. Indeed, previous studies have determined that oceanic silica concentrations were unusually high in the Precambrian (Siever, 1992; Maliva et al., 2005; Konhauser et al., 2007; Tarhan et al., 2016) and promoted chert precipitation in the absence of hydrothermal input (Shen et al., 2011; Ramseyer et al., 2013; Dong et al., 2015; Shang et al., 2018; Penman and Rooney, 2019) or biogenic mediators (e.g., siliceous sponges, foraminiferans, radiolarians, diatoms, and chrysophytes, all of which appeared much later). Early Ediacaran cap chert formation in Africa and Mongolia has been linked to the global post-Marinoan deglaciation and demonstrates the oversaturation of silica in the ocean and perturbations in the silica cycle, as was in the carbon cycle (Penman and Rooney, 2019), at least regionally. Similar conditions might therefore have existed in the post-Gaskiers deglaciation interval, which likewise evinces carbon cycle disturbances (Halverson et al., 2010; Zhou et al., 2017; Cui et al., 2019).

High oceanic silica concentration in the Ediacaran has been further implicated as a cause of chert precipitation and enhancement of fossil preservation by syngenetic and diagenetic mineralization of organic remains particularly in reducing microenvironments, such as that reconstructed for the Doushantuo Formation of South China (Muscente et al., 2015; Shang et al., 2018) and Ediacara Member of South Australia, where pervasive silica precipitation cemented casts and molds of soft-bodied organisms (Tarhan et al., 2016). We suggest that not only the preservation mode has been prompted by high silica concentration, but that metazoans invented the biochemical sequestration of silica into their walls for constructing the skeleton due to this condition, and much earlier as recorded in the Digermulen Penisula.

The cold-water ocean with high silica concentration flooded the margin of Baltica during the post-Gaskiers deglaciation, and such marine conditions are in concert with the position of Baltica in the high paleolatitudes (30–60°S; Torsvik and Cocks, 2017). Globally, the cold-water masses that are prone to silica saturation extend in high-latitude oceans (deep and shallow layers) and in the deep bottom layers in equatorial to low latitudes. Oceans have been stratified with respect to the thermal gradient throughout most of geologic history, and the surface temperature either follows climatic zones (Leeder, 1994; Garrison, 1999) or is uniformly low during the glacial periods, as in the Cryogenian (Hoffman and Schrag, 2002; Hoffman et al., 2017). The middle Ediacaran environmental conditions and marine silica saturation in the basin of the Digermulen Peninsula could have imposed and been opportunistically utilized by macroscopic animals, which needed support for their soft bodies by constructing skeletons.

The primary siliceous skeletons of the new taxa appear to be in contrast to exclusively calcareous skeletons of the terminal Ediacaran animals (Namacalathus, Namapoikia), given that calcite/aragonite are currently thought to be the only minerals precipitated by early biomineralizing organisms (Wood et al., 2019; but see Yang et al., 2020). The paleogeographic position of successions bearing these calcareous skeletal animals is in, or restricted to, the intermediate to equatorial paleolatitudes (Zhuravlev et al., 2012; Warren et al., 2017), thus both climate and marine water were warm. Widespread marine cements facilitated the calcareous mineralogy of animal skeletons and were suitable for development of reefal and biohermal ecosystems at this time (Wood et al., 2019). The environmental conditions, ocean water chemistry, and Ca ion concentration have been inferred to trigger skeletonization in these animals (Grotzinger et al., 2000; Wood, 2011; Penny et al., 2017; Wood et al., 2017, 2019). However, diametrically different conditions existed when (25 Ma earlier) and where (high paleolatitudes, cold climate, and deglaciation zone) the new taxa appeared, and their ecology and metabolism have to be assessed accordingly.

Paleoecological implications

Deduced from the depositional setting of the Stáhpogieddi Formation (Reading and Walker, 1966; Banks, 1973; Farmer et al., 1992; Rice et al., 2011; Torsvik and Cocks, 2017) and preservation of fossils, the newly described organisms inhabited the marine offshore shelf of a marginal inland basin located in high southern paleolatitudes, in quiet and below wave-base conditions during the transgressive, cold-water flooding in the post-glacial time (Reading and Walker, 1966; Banks, 1973; Farmer et al., 1992; Rice et al., 2011; Torsvik and Cocks, 2017). These environments of non-agitated and light-penetrated clear, well-oxygenated water column above the siliciclastic, laminated muddy bottom hosted the sea floor-dwelling, sessile organisms, such as co-occurring Palaeopascichnus and slightly later appearing Aspidella are interpreted to have been (Gehling et al., 2000; Gehling and Droser, 2009; Wan et al., 2014; Burzynski and Narbonne, 2015; Tarhan et al., 2015; Jensen et al., 2018; Kolesnikov et al., 2018). The fossils of the new genera Anulitubus, Coniculus, and Fistula were preserved in situ, although not in a life position (post-mortem collapse) and without transport (autochthonous assemblage) because there is no record of sedimentary structures indicative of currents, wave activity, water flow, or dragging. The lifestyle of these organisms, based on functional morphology of their skeletons with closed proximal and opened apertural ends, distinct longitudinal polarity, but without ventral-dorsal orientation, was most likely that of a benthic, sessile organism.

The skeleton aperture functioned likely as oral opening, which was probably lined with tentacle-like organs or chaetae to actively collect or passively uptake nutrient particles during suspension or filter feeding. From the initial skeleton with probable attachment structures (cup- and knob-shaped elements) or simple proximal cone point that was formed in the nepionic larval stage and embedded in the sediment, the organism grew above the substrate and added to the skeleton upward around the apertural opening. Such reconstructions imply an upright epibenthic life position. Forming skeletons to support their sizable soft bodies in an upright position might be beneficial for the uptake of nutrients from the water column and still being firmly settled at the sea floor. These organisms were probably feeding on bacteria and unicellular algae that formed microplankton (“marine snow”) and not on the ground bacterial mats (neither recognized by the presence of organic films nor the sediment texture).

As for Anulitubus n. gen., the life history may be further suggested. This organism larva formed the cup element and the tube was secreted into adult stage and achieved the full diameter after the first annular segment either was ready or increased minimally. The tube compartments, equivalent to the annular segment spaces, were separated by incomplete septa and the organism occupied the last, distal compartment at the time while constructing the wall of the next segment, and having a widely open aperture to the environment outside. The central hole in incomplete septum provided the free connection between compartments and possibly the way for attachment to the tube bottom by a sort of organic thread or peduncle-like extension to keep the organism permanently within the tube and not risking it being expelled through the aperture and lost.

Specimens of Anulitubus n. gen., Coniculus n. gen., and Fistula n. gen. were all found in close proximity as solitary individuals, but in contrast to other Ediacaran organisms that were living in more abundant and inter-related communities (Cai et al., 2014; Wood et al., 2019), they do not appear to have been communal. For example, Cloudina, Conotubus, and Sinotubulites have been recognized as sessile, gregarious organisms living on the bottom floor covered by microbial mats and in association with thrombolitic and stromatolitic reefs or bioherms on carbonate platform ramp (Germs, 1972; Grant, 1990; Cai et al., 2014, 2015). Shaanxilithes and Gaojiashania are also thought to be epibenthic organisms (Cai et al., 2010,, 2013; Tarhan et al., 2014).

The late Ediacaran ecosystems (ca. 570–540 Ma) of soft-bodied and biomineralizing animals alongside the Ediacara-type organisms have been reconstructed as diverse and highly integrated communities in the shoreface sandy shelf, restricted lagoonal to inner and outer ramp of carbonate platform settings, with a sea floor rich in microbial mats (Wood et al., 2019). These various settings were paleogeographically located in the intermediate to low latitudes (Zhuravlev et al., 2012; Torsvik and Cocks, 2017; Warren et al., 2017) and in warm climate zones with intense carbonate precipitation in warm-water environments.

The early, ca. 571 Ma ecosystem reconstructed from Avalonia, Newfoundland, is dominated by rangeomorphs and consists of taxonomically rich, leaf-like, frondose morphotypes with holdfasts, that were benthic and oriented upright in the water column (Liu et al., 2015; see Pu et al., 2016 for age; Butterfield, 2020). This biota lived in deep-water marine, aphotic environments at the time and later on (ca. 560–555 Ma) in other shallower-water environments (Narbonne et al., 2014; Liu et al., 2015; Wood et al., 2019). The Avalonian deep basinal setting recorded by siliciclastic hemipelagites (mud, silt) that developed in the turbiditic facies was located in high paleolatitudes (40–60°S; Liu et al., 2015). Similarly, the diverse frondose and modular rangeomorphs, arboreomorphs, and erniettomorphs thrived on the sea floor, in aphotic zone of the deep-water setting in the Mackenzie Mountains of Laurentia, at ca. 580–560 Ma interval (Narbonne et al., 2014). This basin was equatorially positioned at the time (Torsvik and Cocks, 2017). The deep marine environment in the aphotic zone has been inferred to be the site of the origin of the rangeomorph biota and, accordingly, their osmotic mode of feeding (Narbonne et al., 2014; Liu et al., 2015). Recently, rangeomorphs have been revised as being not osmotrophic, but rather heterotrophic and digesting substrates, likely living in symbiosis with a resident microbiome, and within the total-group Eumetazoa (Butterfield, 2020). Although the Laurentian and Avalonian basins were widely paleogeographically spread across equatorial to high paleolatitudes, they both represent deep-basinal and thus cold-water environments. Subsequently, the same taxa existed in shallow marine environments ca. 560 Ma (Wood et al., 2019), which were warm water and in the photic zone. The puzzle remains how the rangeomorphs and similar biota adapted without change in morphology and physiology from their initial deep, cold-water, aphotic environments to shallow, warm-water, and photic environments?

To these reconstructions, the ecosystem estimated to ca. 575 Ma from the shallower marine, offshore shelf but cold-water setting on the Digermulen Peninsula has to be added. This ecosystem comprised epibenthic macroscopic organisms and neritic microplankton. Epibenthic were the new genera Anulitubus, Coniculus, and Fistula animals, Palaeopascichnus (Gehling and Droser, 2009; Wan et al., 2014), which is a protistan foraminiferan (Antcliffe et al., 2011; Kolesnikov et al., 2018), and slightly above appearing Aspidella, which is a member of the Ediacara biota of unknown affinity (Gehling et al., 2000; Burzynski and Narbonne, 2015). The planktonic photosynthesizing microorganisms were organic-walled acritarchs that are predominantly of algal affinities (Moczydłowska, 2016; Liu and Moczydłowska, 2019) and filamentous cyanobacteria (unpublished data). Benthic bacterial mats are not yet recorded, but this may reflect a taphonomic bias because of the sediment oxidation, and they are more susceptible to decay in comparison with resistant cysts of acritarchs.

The Digermulen ecosystem extended in the high paleolatitude (30–60°S), cold-water marine shelf with siliciclastic sedimentation, in the photic zone, and above the oxic chemocline. During the post-Gaskiers deglaciation, the revived hydrological cycle that caused mixing of the well-oxygenated water and increased input of mineral nutrients into the marine basin from weathering of the uncovered Baltica paleocontinent created habitats for oxygen-demanding animals and favorable living conditions. The autotrophic photosynthesizers forming the marine snow and living on the sea floor generated the organic matter for the consumers. This ecosystem hosted the earliest silica biomineralizing animals described here.

We suggest that the cold-water environments, not necessarily restricted by the depth but by the temperature, were the sites of the earliest frondose-like biota expansion. If the rangeomorph biota is epifaunal (Narbonne et al., 2014) and among total-group eumetazoan (Butterfield, 2020), the Avalon assemblage included candidate cnidarians (Liu et al., 2014; Wood et al., 2019) and the animal Dickinsonia (Hoyal Cuthill and Han, 2018; Bobrovskiy et al., 2018), and the new biomineralizing taxa were animals, therefore the records at ca. 575–571 Ma show the radiation of animals. There were both small but macroscopic skeletal animals and large soft-bodied biota occurring in shallow- and deep-marine environments, but all cold-water. The evolution of biomineralization by silica, as recorded in the Digermulen Basin, seems to be confined to cold-water and oxic environments.

The Cold Cradle hypothesis was proposed for the evolution of the Ediacara biota and facilitated by the increased O₂ availability in cold water environments (Fedonkin, 2003). More recently, it has been suggested that estimates of O₂ level in the ocean would be not sufficient to develop diverse ecosystems and, instead, thermal stability, not necessarily cold conditions, has been hypothesized for the deep-water origin of Ediacara biota (Boag et al., 2018). We suggest the stenothermal (Boag et al., 2018) but cold-water (Fedonkin, 2003; the Digermulen Basin herein) environments for the origin of the frondose Ediacara biota (deep water but expansion into shallow water) and for the earliest biomineralization (shallow, cold-water environments).

Conclusions

The three newly described monotypic genera Anulitubus, Coniculus, and Fistula of macroscopic body fossils with mineralized walls display morphologic features, radial symmetry, and longitudinal polarity that are consistent with being skeletons of animals. Their paleoecology and the paleoenvironment of a marine shelf with well-oxygenated, cold quiet water in the photic zone above the muddy sea-floor support this interpretation and, together with characteristic functional morphology of the exoskeletons, suggest epibenthic mode of life and suspension/filter feeding by these animals. The new species differ from any previously described Ediacaran body fossils by spacial arrangement of annuli and segments, and by being uniquely primarily silicified.

The single- and double-layered skeletons are siliceous and their composition and primary mineralogy are revealed by petrography and various geochemical analyses, and supported by diagenetic processes of the host sediment. Biomineralization by silica was likely facilitated by the environmental conditions of cold, marine waters with high silica concentration, which are typical of silica sequestration into organismal skeletons. Such environments of a cold-water marine shelf existed in the Digermulen Basin during the post-Gaskiers (>580 Ma) deglaciation time and contemporaneously with the appearance of the new taxa. These environments could have forced and been opportunistically utilized by the increase in size of the animals, which needed support for their soft bodies by constructing skeletons. The new species evolved biochemical sequestration of silica into their walls and preceded significantly calcareous biomineralizing sponge (Namapoikia) and putative lophophorate animals (Namacalathus) at the 550 Ma terminal Ediacaran.

The age of the new fossil taxa is broadly bracketed within 580–541 Ma, but shortly post-dating the 580 Ma underlying Mortensneset Diamictite, which is coeval with the Gaskiers Glaciation. This age is estimated at ca. 575 Ma by the age interpolation within the sedimentologically continuous succession bracketed within 580–541 Ma and further constrained by the 558–555 Ma chronostratigraphic ages of the overlying Ediacaran biota.

The reported fossils are the earliest known in the evolutionary history biomineralizing animals that secreted siliceous skeletons and the only ones known from the Ediacaran Period. The phylogenetic affinity of the new genera remains un-resolved, but the body plan, complexity of skeleton with annulate elements, and initial larval structures indicate eumetazoan grade of organismal organization. The new biomineralizing animals evince the radiation among other Ediacaran organisms that are recorded by soft-bodied rangeomorphs of unknown affinities and appearing at 571 Ma. The oldest putative animals, including cnidarians and bilaterians preserved as carbonaceous compressions, the Lantian biota, are estimated here to ca. 620–580 Ma in age. The divergence of diverse Ediacaran animals that were soft-bodied and subsequently biomineralizing by silica and calcareous minerals is recorded by fossils in the evolutionary stepwise pattern since ca. 620 Ma until the beginning of the Cambrian.