Co-evolving N-Fe redox processes controlled iron minerals in banded iron formation

doi:10.21203/rs.3.rs-3724120/v1

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Co-evolving N-Fe redox processes controlled iron minerals in banded iron formation

https://doi.org/10.21203/rs.3.rs-3724120/v1

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Biogeochemical cycles in the Precambrian ocean responded to the co-evolution of biosphere (microorganisms) and the physicochemical structure (e.g., redox, temperature) of the ocean, which closely link to the enigma of banded iron formations (BIFs) that primarily triggered by massive Fe(II) oxidation under anoxic-hypoxic condition for two-billon years (~3.8-1.8 Ga). The current Fe(II) oxidation models, however, rarely consider the effects of the evolution of coupled biogeochemical cycles on secular succession (shifting from magnetite to hematite) of dominant iron minerals in BIFs. Here, we investigated the evolution of coupled Fe-N redox processes under the simulated Precambrian ocean conditions, and propose a dynamic model for the origin of iron mineral succession in BIFs: During the early-mid Archean, NO₂^- was mainly produced by nitrification in the oceans of warm-hot temperatures (>50-60 ^oC), which favored the primary precipitation of Fe(II)-Fe(III) oxides (magnetite) and silicates (cronstedtite) in the early BIFs. Subsequently, the cooling and oxygenation of paleo-ocean near the GOE promoted the input of both NO₂^-and NO₃^-, resulting in co-precipitation of an increasing amount of Fe(III) minerals (goethite and lepidocrocite as precursors of hematite). This dynamic N-Fe coupling model explains the observed secular transition of iron mineral phases in BIF deposition.

Earth and environmental sciences/Biogeochemistry/Element cycles

Earth and environmental sciences/Planetary science/Mineralogy

The early Precambrian ocean, especially during Archean to Paleoproterozoic, was characterized by low oxygen, rich reductants, and probably warm-hot temperature (>50 ℃, probably higher than 70 ℃) conditions¹, but irreversibly evolved into oxic and temperate conditions in the late Neoproterozoic². Although many details remain controversial about the evolutionary history of the ocean³, there is no doubt that large environmental changes, especially the gradual cooling and stepwise oxygenation of the ocean, fundamentally shaped the physiochemical structure (e.g., ferruginous to stratified euxinic to oxic) of the oceans and the microbial world (e.g., thermophilic anaerobes evolving into mesophilic aerobes). The evolving marine microbiota thereby would have been responsible for the changing biogeochemical reactions (especially for redox sensitive elements such as N, Fe, C) and their products over geological time^{4, 5}. A major event closely intertwined with paleo-ocean biogeochemical evolution is the massive deposition of BIFs from the seawater column through extensive Fe(II)_aq oxidation during the early Precambrian (dominantly during 3.8-1.8 Ga)⁶. The geochemical, isotopic and mineralogical records of BIFs exhibit temporal and spatial variations. For example, the average valence state of Fe in BIFs gradually increased over time, because of a shift in dominant Fe minerals from magnetite (especially >2.9 Ga) to hematite (particularly after the Neoarchean)⁷, suggesting that Fe(II) oxidation process evolved. Currently photo- and/or O₂-dependent Fe(II) oxidation in the shallow euphotic zone of paleo-oceans is widely considered to account for BIF precipitation, but this model is difficult to explain the intriguing mineralogical or geochemical changes of BIFs over time⁸. How the evolving biogeochemical processes affect Fe(II) oxidation pathways and the resulting mineral products across anoxic-oxic transition was rarely considered in the previous models.

The early anoxic ocean (> 3.0 Ga) was filled with ammonium-replete ferruginous waters, where the main anaerobic biological processes included N₂ reduction to NH₄⁺ and Fe(II)-based anoxic photosynthesis (photoferrotrophy) ^{9, 10}. At this time, vertical profiles of NH₄⁺ and Fe(II) in paleo-oceans were uniform (Fig. 1a). As oxygenic photosynthesis emerged to gradually consume reductants¹¹, the uniformly reducing water column evolved into a heterogeneous one due to the developments of redoxcline and chemocline (Fig. 1b)^{12, 13}. These developments would have greatly changed the pathways of Fe(II) oxidation¹⁴. Under the rising redox state of the paleo-oceans, NH₄⁺ would be continuously oxidized to nitrogen oxyanions (NO_x^-, mainly NO₂^- and NO₃^-) even in anoxic waters^{15, 16} that can be dated back at least to the Mesoarchean^{12, 17}. As a result, an intermediate nitrogenous zone gradually developed between the upper oxidizing surface water and the lower ferruginous water (Fig. 1b)^{18, 19}. This nitrogenous zone is commonly found in oxygen minimum zone (OMZ) of modern oceans and ferruginous lakes, both regarded as the analogues of the hypoxic Precambrian ocean^{20, 21}. This nitrogenous layer is crucial to Fe(II) oxidation because it would act as a redoxcline/chemocline barrier by consuming O₂ and CO₂ from above via nitrification and by oxidizing the upwelling Fe(II)_aq from below via N-dependent Fe(II) oxidation (general reaction: Fe(II) + NO_x^- → Fe(III) + N₂)^{22, 23}. The N-dependent Fe(II) oxidation mechanism would greatly decrease the dominance of the traditional photo- and/or O₂-dependent Fe(II) oxidation by spatially separating the reactants required in previously proposed Fe(II) oxidation mechanisms (Fig. 1b)⁸. Interestingly, the peak deposition of BIFs at 2.7-2.4 Ga, after the onset of aerobic nitrification (and the presence of nitrogenous layer)²⁴ but before the remarkable rise of atmospheric O₂ marked as the Great Oxygenation Event (GOE, 2.4-2.2 Ga)^{25, 26}, argues for the potential importance of coupled N-Fe redox cycling to BIF deposition. However, direct evidence is lacking. Here we hypothesize that an active, biologically-driven N-Fe redox process develops around the nitrogenous layer in the shallow water (Fig. 1) that is responsible for Fe(II) oxidation and substantially contributes to BIF formation. To test this hypothesis, we investigated how the evolution of coupled processes of Fe(II) oxidation and N redox cycling affect the formation of BIF-related mineral products under the simulated Precambrian ocean conditions.

Ammonia, formed abiotically on early Earth (e.g., volcanic, lightning, bolide^{30, 31}, serpentinization of ultrabasic rocks³²) and biotically at later times¹⁰, was the predominant form of fixed N pool in the early reducing ocean (Fig. 1a)^{9, 33}. Aerobic ammonia oxidation, which can occur under extremely low oxygen level (at nM level)³⁴ or even in anoxic waters¹⁶, was dated back to at least ~ 3.0 Ga¹³ and would have triggered the N redox processes by supplying NO_x^- species to the ancient oceans (Fig. 1b), but it was scarcely considered by current biogeochemical models. For example, at a low O₂ level, the end products of the N redox processes were highly sensitive to temperature, probably via the effect of temperature on microbial functions and incomplete nitrification and denitrification processes in response to temperature changes.

To simulate the N redox processes in a gradually cooling hypoxic Precambrian ocean, nitrifying and denitrifying microbial consortia were enriched from hot springs in Xizang and Yunnan Provinces, China, which span broad geochemical and temperature gradients (Fig. S1, Table S1 and S2). The reason to use hot spring as an analog of Precambrian ocean was because key physicochemical properties in hot springs, such as low contents of oxygen and organic carbon, and high contents of Fe, N and Si, are similar to those of the paleo-oceans³⁵. Furthermore, the microorganism-dominated community in hot springs, including many thermophilic ancestors, is a good analog of the simple biosphere on the early earth³⁶.

Nitrifying enrichments were first recovered from diverse hot springs ranging from 75 ^oC to 30 ^oC (Fig. S2 and S3). Interestingly, the dominant product of NH₄⁺ oxidation gradually shifted from NO₂^- to NO₃^- in response to a temperature decrease, with a temperature cross-over at 60 − 50 ^oC (Fig. 2a) regardless of the spring water geochemistry. Consistently, ammonia oxidizing archaea (AOA, performing NH₄⁺ oxidation to NO₂^-) were ubiquitous in the enrichments, with thermophilic Nitrosocaldales being dominant in higher temperature (> 40 ^oC) and Nitrososphaerales at lower temperature. Nitrite-oxidizing bacteria (NOB), mainly Nitrospirales, that can oxidize NH₄⁺/NO₂^- to NO₃^- were only revealed at lower temperature (< 40 ^oC, Fig. S3). This observation is consistent with previous studies in that AOA were detected at temperature up to 90 ^oC, but most NOB isolates, if not all, were active under 60 ^oC^{37, 38, 39, 40}. Therefore, a transition of fixed N from NH₄⁺ to NO₂^- and then to NO₃^- may have occurred as a result of cooling of the Precambrian ocean.

NO_x^- generated by nitrification may be reduced by canonical denitrification under reducing condition. Therefore, canonical denitrifying consortia under the same temperature gradient were enriched to explore the stability of produced NO_x^- under different temperature conditions (Fig. S4 and S5). Interestingly, accumulation of NO₂^- resulting from NO₃^- reduction significantly decreased with temperature cooling (Fig. 2b), i.e., higher temperature favored production of NO₂^-. This was well illustrated in one representative hot spring flow channel, in which denitrifying consortia stoichiometrically accumulated NO₂^- from NO₃^- reduction at 65 ^oC, less at 55 ^oC, and undetectable at 50 ^oC (Fig. S5). The denitrifying consortia were dominated by Hydrogenobacter (more abundant at high temperature) and Thermus (Fig. S4), two types of common denitrifying bacteria in hydrothermal habitats^{41, 42}. In fact, incomplete denitrification of NO₃^- and accumulation of NO₂^- were commonly observed in thermophilic cultivations^{43, 44, 45} and oxygen-depleted waters⁴⁶. Therefore, NO₂^- accumulation from ammonia nitrification or nitrate reduction would be a substantial source of N species especially under elevated temperatures.

In summary, an evolving N inventory would have developed in the hypoxic cooling Precambrian ocean. A NO₂^--dominant nitrogenous layer may gradually form in the shallow water by incomplete nitrification (NH₄⁺→NO₂^-) and denitrification (weak NO₂^- reduction), especially in the early high temperature (> 50–60 ^oC) ocean. With the cooling of the ancient ocean, NO₃^- was gradually produced in larger quantities through complete nitrification at lower temperature (especially < 50 ^oC), leading to the pervasive mixture of NO₂^- and NO₃^- in the ocean, although an increasing proportion of NO_x^- was also reduced into gaseous N species (e.g., N₂O, N₂) through denitrification.

In addition to production of NO_x^- via nitrification and partial removal via denitrification, they can be reduced by an alternative and even more abundant electron donor in the ferruginous Precambrian ocean–the rich Fe(II)^{20, 27}. Fe(II) can be chemically oxidized by NO₂^- reduction (known as chemodenitrification) and/or microbially induced by NO₃^- reduction (known as nitrate-dependent Fe(II) oxidation, NDFO). Vertically, the NO_x^--rich layer just overlayed the lower Fe(II)-rich ferruginous water, providing an opportunity for chemo- and microbial denitrification. However, how NO_x^--dependent Fe(II) oxidation occurred under the scenario of the evolving NO_x^- pool and cooling temperature are unclear. Therefore, chemical and biological induced NO_x^- - dependent Fe(II) oxidation was simulated under a broad range of temperature schemes (25-75^oC), with consideration of co-effects on mineral formation of HCO₃^- and dissolved Si that were rich in the Precambrian ocean^{47,48, 49}.

Chemodenitrification was tested at temperatures ranging from 25 ^oC to 75 ^oC, which showed higher Fe(II) oxidation and NO₂^- reduction (both the rate and extent) by higher temperature (Fig. S6). Mineral formation also changed systematically with temperature cooling (summarized in Table 1). In the absence of dissolved Si, magnetite that can reach micrometer-scale size was the main or even only mineral product at temperature higher than 55 ^oC (Fig. S7, S9-a3, a4, b3, b4; Fig. S8-a, b). When HCO₃^- was present, magnetite started to form at 40 ^oC (Fig. S7, S9-b2; Fig. S8-a). In contrast, lepidocrocite and goethite predominantly precipitated at temperatures of 25 and 40 ^oC when supplemented with Cl^- and HCO₃^-, respectively (Fig S7, S9-a1, a2, b1, b2). Addition of dissolved Si did not prevent Fe(II) oxidation by NO₂^- but significantly altered the iron mineral phases and crystallization, as the main minerals turned to be weakly crystallized cronstedtite (empirical formula: Fe^2 + ₂Fe^3 + ₂SiO₅(OH)₄, an Fe(III)-containing greenalite-like serpentine) and greenalite while a small part of magnetite only existed at higher temperature (Fig. S7, S8, S9-c).

Biologically induced NO_x^- - dependent Fe(II) oxidation was first tested with the use of a typical thermophilic NO₃^--reducing isolate from a Xizang hot spring, Thermus sp., at 55 ^oC and 75 ^oC. Fe(II) was effectively oxidized by NO₃^- reduction at both temperatures, regardless of the presence of Si (Fig. S10). NO₂^- was accumulated from NO₃^- reduction without Fe(II) (Fig. S10a), but it was not observed in the presence of Fe(II) (Fig. S10-c, d), suggesting that Fe(II) oxidation either consumed the produced NO₂^- or bypassed it during the NO₃^- reduction pathway. Magnetite was the only mineral product at both temperatures without Si (Fig. S11-a1, a2). When Si was present, cronstedtite was the dominant well-crystalline mineral at 55 ^oC (Fig. S11-b1), accompanied with some greenalite (Fig. S8-d). Magnetite was co-precipitated with cronstedtite and greenalite at 75 ^oC (Fig. S11-b2, Fig. S8-e). Notably, microbially formed silicate minerals in Si-amended treatment were more crystalline compared to those formed by chemodenitrification (compared Figs. S11-b and S7-c). Many of these minerals tightly encrusted the cell surface (Fig. S12), suggesting that microbial activity greatly stimulated mineral crystallization.

Three representative denitrification consortia, enriched from a spring channel (Fig. S1-b1) of different temperatures but similar water chemistry, were tested for their performance in Fe(II) oxidation. The three consortia were dominated by Hydrogenobacter sp. at high temperature (65 ^oC) and by Thermus sp. at lower temperature (55 and 50 ^oC, Fig, S4). Fe(II) oxidation occurred in consortia at higher temperature (> 50 ^oC) when NO₂^- was produced as a byproduct of NO₃^- reduction (Fig. S13-b, c), but did not occur without consortia or in consortia of lower-temperature (50 ^oC) (Fig. S13-a), suggesting that Fe(II) was oxidized metabolically and/or chemically by NO₂^-. The mixture of cronstedtite and magnetite was the main mineral products (Fig. S11-c), similar to that formed by pure Thermus sp. strain in the presence of Si, probably due to the fact that the hot spring water contained a considerable amount of Si (Table S1). Mineral particles, especially for the cronstedtite formed by the denitrifying consortia (Fig. S14), however, were larger in size than those formed by pure Thermus strain, perhaps due to slower Fe(II) oxidation rate or interspecies interaction in the microbiota.

In summary, higher temperature (over ~ 50 ^oC) favored the precipitation of Fe(II)-Fe(III)-mixed minerals, especially magnetite, either via chemical or biological Fe(II) oxidation when coupled with NO_x^- reduction; in comparison, lower temperature was more beneficial to the formation of Fe(III) minerals (Table 1). This variation in mineralogy in response to a temperature gradient is consistent with the trend observed in natural BIFs, that is, early BIFs (at higher temperature) were dominated by magnetite, and later (at lower temperature) gradually shifted to hematite (as partially summarized in Table S3). In addition, the presence of denitrifying microbial consortia tends to promote the formation of magnetite at lower temperature, with a higher proportion and better crystallinity relative to their absence, suggesting the importance of direct (mediate N-Fe reactions) and indirect (such as providing mineral adsorption sites for crystallization) biological effects on iron mineral precipitation in BIFs.

Table 1

Iron minerals precipitated by chemical and biological N-Fe redox coupling under temperature gradient ^a
Temperature (^oC)	Reactions	Iron minerals
Temperature (^oC)	Reactions	Lepidocrocite	Goethite	Magnetite	Silicates (Cronstedtite, Greenalite)
< 40	Chemodenitrification	+ +	+ +	-	+ + +?
< 40	Microbial NDFO	/	/	/	/
40–55	Chemodenitrification	+ +	+ +	+ +	+ + +
40–55	Microbial NDFO	-	-	+ + +	+ + +
> 55	Chemodenitrification	-	+	+ + +	+ + +
> 55	Microbial NDFO	-	-	+ + + + +	+ +

^a Summarized based on the mineral composition in Figs. S7-S9, S11, S12, S14. More signs of “+” denote higher abundance. Black “+” symbols show the mineral composition in Si-absent treatments, while red “+” symbols show that in Si-rich systems; “-” denotes not detectable, and “/” denotes not tested in this study.

Reconstruction of the effect of biogeochemical evolution on BIF across geological timescale is extremely difficult due to the uncertainty of changes in key environmental factors. For example, oxygen level is a crucial factor determining whether Fe(II) oxidation is influenced by oxic or anoxic processes. However, it is difficult to determine the evolution of oxygen level in paleo-oceans with heterogeneous chemical structures, despite the widespread consensus of an overall rise of oxygen in the atmosphere of the early Earth. In comparison, temperature cooling is more predictable. From a long-term geological perspective, temperature cooling of the paleo-oceans is an indisputable fact, although the absolute magnitude of cooling is still controversial. Nevertheless, the cooling trend of temperature of the Archean-Proterozoic oceans would greatly reshape the structure and functional composition of microbial communities, thereby altering biogeochemical processes in the ocean, including the N and Fe cycling and BIF deposition. In the present work, by integrating all the results, we developed a four-stage evolution scenario for the coupled N-Fe redox processes, providing a possible explanation for the secular changes of iron mineral precipitation in BIFs in the Archean-Paleoproterozoic ocean (Fig. 4):

Small scale BIFs (i.e., with typical laminated Si-rich and Fe-rich layers) began to deposit as early as 3.8 Ga. ago.⁵⁰ The ocean was hot (most probably higher than 55 ^oC, even over 75 ^oC)^{1, 47, 51}, highly reducing, and almost devoid of any microbial activity (if any). Iron minerals in these oldest BIFs are overwhelmingly dominated by magnetite, accompanied by some Fe-silicates, such as grunerite, but few contain hematite⁵². During this prebiotic era, abiotic N₂ fixation was the main process of introducing a low flux of fixed N, (initially NO₂^- ) (~ 10⁶-10¹⁰ mol N∙year^-1 or even higher), into the ferruginous ocean^{30, 53}. The NO₂^- would be rapidly reduced through chemical oxidation of Fe(II) especially at high temperature, followed by precipitation of iron in the form of magnetite (2NO₂^- + 9Fe²⁺ + 16HCO₃^-→N₂ + 3Fe₃O₄ + 8H₂O + 16CO₂, as shown in Fig. S6-c), and/or cronstedtite in the presence of Si. Therefore, although NO₂^- would not be accumulated to a considerable concentration, its abiotic supply was sufficient to precipitate a large proportion of the oldest BIFs (totally ~ 10¹¹-10¹² mol Fe in the Paleoarchean BIFs estimated by the BIF deposit size and average 20 wt% Fe content, Fig. 3a). Moreover, the primary production of Fe(II)-Fe(III)-mixed minerals (magnetite, silicates) eases the vexatious requirement of a reduction step of ferric iron precursors (e.g., ferrihydrites), as proposed previously⁵⁴. This scenario offers a reasonable explanation for the higher abundance of magnetite in the early BIFs⁷.

After the Mesoarchean, deposition of BIFs increased slightly, until the Neoarchean. Magnetite is still the predominant iron mineral in most BIFs (particularly > 2.9 Ga)⁷. During the Mesoarchean, biological N₂ fixation bloomed and greatly enlarged the NH₄⁺ flux to the ocean^{9, 10}. Microbial oxidation of NH₄⁺ may also emerge as early as in the Mesoarchean⁵⁵, as AOA could have oxidized ammonia at extremely low levels of O₂ or even in the anoxic water ^{15, 16}. Moreover, local “oxygen oasis”, due to the emerging oxygenic photosynthesis⁵⁵, may have gradually stimulated NH₄⁺ oxidation¹². Therefore, microbial nitrification provided additional NO_x^- flow, mainly NO₂^-, as the temperature was still likely much higher than modern oceans (over 50 ^oC, as indicated in Figs. 2a and 4). With the increase of NO₂^- flux, chemical oxidation of Fe(II) by NO₂^-, which primarily forms magnetite and cronstedtite, was still the predominant process of coupled N-Fe processes. This scenario offers an explanation for the increase of magnetite-dominant BIFs volume (totally ~ 10¹⁴-10¹⁵ mol Fe in Mesoarchean BIFs as estimated by the BIFs deposit size and average 20 wt% Fe content, Fig. 3b)⁶.

BIF deposition increased significantly after the Neoarchean, reaching a peak from the mid-Neoarchean to the early Paleoproterozoic (2.7–2.4 Ga), right before GOE. Intriguingly, hematite content increased dramatically in these BIFs, especially after 2.6 Ga, suggesting that certain important transition of biogeochemical processes might have occurred and significantly affected the Fe redox cycle. Indeed, the oxygenation of surface water was more pervasive at this time, which greatly stimulated aerobic nitrification and further increased the NO_x^- flux. For example, the extremely positive excursions in δ¹⁵N was observed in the 2.7–2.5 Ga sediments, which was interpreted as a great rise in N loss via coupled nitrification-denitrification processes^{56, 57}. Therefore, the abundant NO_x^- might be extensively reduced by Fe(II) oxidation considering the super-large BIF deposition during this period. However, as temperature continued to decrease (lower than 50 ^oC), microbial nitrification of NH₄⁺ would produce a mixture of NO₂^- and NO₃^- with an increasing share of NO₃^-. With the increased NO_x^- pool/flux, Fe(II) oxidation coupled with microbial NO₃^- reduction and chemical NO₂^- reduction produced a mineral assemblage of magnetite, silicates, and Fe(III) hydroxides (i.e., goethite, lepidocrocite, Fig. 3c). Over time at certain temperature and pressure, these ferric hydroxides can transform to hematite⁵⁸. Furthermore, the relative dominance of Fe(III) hydroxides produced by N-Fe redox coupling would gradually increase at lower temperatures (Table 1), which offers a plausible explanation for the transition of magnetite-dominant to hematite-dominant BIFs during this period.

A dramatic decrease or even stop of BIF deposition was observed after 1.8 Ga when the temperature further cooled down, suggesting a quiescence of Fe(II) oxidation at this time (Fig. 3d). Two possible reasons may account for this transition. First, canonical nitrification-denitrification cycles might outcompete the coupled N-Fe cycles, due to rapid and complete oxidation of NH₄⁺ (into NO₃^-) and complete reduction of NO₃^- to N₂O and N₂ gases (Fig. S1d, S4c) at low temperatures. Second, the development of an euxinic zone (S^2- rich layer, caused by increased sulfate input from oxidative weathering of continental sulfides) between the nitrogenous and ferruginous layers provided an important reductant (i.e., S^2-) to consume NO_x^- and/or precipitate Fe(II)⁵⁹, both of which would weaken the coupling of the N-Fe cycles. This scenario offers a possible explanation for the extremely low BIFs during the Meso- and Neoproterozoic (without considering the Rapitan IFs deposited during the Neoproterozoic).

In summary, our proposed evolution model of the coupled N-Fe cycling processes, as a response to temperature decrease, provides a reasonable explanation for the successive change of abundance and mineralogy of BIFs. An important reason for the overlooked role of the coupled N-Fe cycles in the BIF formation in most previous studies is the assumption of the low accumulation of oxidative NO_x^- species in the reducing Precambrian ocean. However, we argue that the increasing flux and changing speciation of NO_x^- are extremely important for triggering extensive Fe(II) oxidation and mineral formation. A recent study found that the networked N-Fe reactions would provide substantial N₂O emission and N burial (by association with Fe minerals) in the Precambrian marine ecosystem⁶⁰, supporting the pervasive N-Fe reactions. Here we further demonstrated that N and Fe cycles could also be directly coupled to affect Fe mineralogy. In addition, our observed iron mineral assemblages formed by the dynamically evolving N-Fe coupling also have some other under-appreciated geological and environmental consequences, such as the bioavailability of N (via change of N speciation) and P (via adsorption to iron minerals) overtime, thereby controlling primary productivity and emission of greenhouse gasses N₂O and CO₂⁶⁰. Finally, further work is required for additional assessment of the contribution of N-Fe coupling to BIF deposition, which include, but not limited to, the competition by different Fe(II) oxidation pathways, quantitative impacts of other evolutionary factors (e.g., pH, Ca/Mg), and the N and Fe isotopic footprints recorded in geological archives in response to temperature gradients.

Thermophilic nitrification and denitrification consortia enrichment and analyses

Hot springs in the geothermal fields in Tengchong, Yunnan Province (Rehai, Ruidian and Houqiao), China, with diverse temperature and geochemical gradients, were sampled in September 2019, August 2020 and November 2021. Xizang (Dagejia, Quzhuomu and Qucai) hot springs were sampled in August 2020 (detailed in the Supplemental text, Fig. S1, Tables S1 and S2). Thermophilic nitrifying and denitrifying consortia were preliminarily enriched in the field. At each selected hot spring, spring water mixed with sediment suspension was aseptically injected into pre-autoclaved Balch glass bottles with needles, which served as medium and microbial inoculant. The bottles were filled with air for nitrification enrichment but degassed by N₂ gas for denitrification enrichment. The main reason for using hot spring water as a medium was to minimize any bias caused by defined media. However, enrichment was also performed with defined media⁶¹ in representative hot springs to test the stability of nitrification and denitrification. NH₄⁺ and NO₃^- (final concentrations 0.5-1 mM) were amended as substrates for nitrifying and denitrifying cultures, respectively, and the bottles were incubated in situ for 3–5 days in hot springs to stimulate microbial growth. The in-situ enrichments were subsequently transported to laboratory at ambient temperature in less than a week. In the lab the enrichments were incubated in incubators with temperatures similar to the corresponding springs. The concentrations of NH₄⁺, NO₂^- and NO₃^- were periodically monitored by spectrophotometric methods. The positive cultures that showed NH₄⁺ oxidation or NO₃^- reduction were then transferred to new media. Transfer continued for 2–5 times to obtain stable nitrification and denitrification consortia. Microbial compositions in successful enrichments were analyzed for representative temperatures (75 ^oC, 65 ^oC, 55 ^oC and 40 ^oC) by using the 16S rRNA gene pyrosequencing approach. Microbial and chemical analyses are provided in the SI.

Fe(II) oxidation coupled with NO reduction (chemodenitrification)

Fe(II) oxidation by NO₂^- reduction (NO₂^- + Fe²⁺ → N₂ + Fe³⁺) is also known as chemodenitrification. To mimic iron mineral formation by chemodenitrification under conditions of cooling and water chemistry rich in HCO₃^- and dissolved Si of the Precambrian ocean, four experimental sets were established at the temperature range of 25–75 ^oC (as listed in Table S4) with and without Si (2 mM⁴⁸). The reaction was buffered in PIPEs buffer to maintain the pH at 7, and the levels of Fe(II) (1 mM) and NO₂^- (0.2 mM) were selected to simulate their low concentrations in the Precambrian ocean and ensure enough mineral formation for analysis. The medium was degassed with pure N₂ gas on a gassing station, autoclaved, and sampled in an anaerobic glove box. Fe(II) and NO₂^- levels were regularly monitored and the final mineral products were collected for X-ray diffraction, scanning electronic microscopy, and Mössbauer analyses as detailed in the SI.

Fe(II) oxidation coupled with NO ₃ ^- reduction by Thermus sp.

Microbial Fe(II) oxidation coupled with NO₃^- reduction (NO₃^- + Fe²⁺ \(\underrightarrow{\varvec{c}\varvec{e}\varvec{l}\varvec{l}}\) N₂/N₂O/NO₂^- + Fe³⁺) is common among denitrifying bacteria⁶². A typical thermophilic denitrifier, Thermus sp. isolated from Tibet hot spring (DGJ, 70 ^oC) and having a broad range of growth temperature (50–80 ^oC, optimal 60 ^oC), was used to test its capability of Fe(II) oxidation. Thermus sp. was grown in T5 medium (Table S5), which was then washed three times by 10 mM PIPEs buffer to remove any organic substrates. The collected Thermus sp. cells were used for NDFO test in a final concentration of ~ 1×10⁷ cells per mL. The PIPES-buffered non-growth medium, the same as that for chemodenitrification, was used, with NO₂^- replaced by NO₃^-. Four experimental sets were established at two temperatures (55 ^oC and 75 ^oC) to compare the effects of cells and silica on mineral formation (Table S5).

Fe(II) oxidation coupled with NO reduction by denitrifying consortia

Three canonical denitrifying enrichment consortia from the same hot spring stream with different temperatures (DG1-1, DG1-2 and DG1-3 at 65, 60 and 50 ^oC, respectively, Fig. S1) were selected to test their potential capability on Fe(II) oxidation and mineral formation. The water collected from the source spring was used as the medium for all three consortia. Three experimental sets were established for each consortium to compare their effects on Fe(II) oxidation (Table S6): a) abiotic control without inoculation, b) Fe(II) + inoculated consortia to test any electron acceptor possibly present in the hot spring water itself, and c) NO₃^- + Fe(II) + inoculated consortia.

Competing interests

The authors declare no competing interests.

Author contributions

LH and YD contributed equally to conceptualization, data analysis and draft writing. LH, YD, LL, JY and LM performed data collection. LH, HJ and HD acquired fundings. LH, YD, HD and HJ wrote the original draft, and all authors reviewed and edited the manuscript.

Acknowledgments

This work was supported by the National Natural Science Foundation of China (NSFC- 42192503, 42172340, 42192500, 91951205), and the “Deep-time Digital Earth” Science and Technology Leading Talents Team Funds for the Central Universities for the Frontiers Science Center for Deep-time Digital Earth, China University of Geosciences (Beijing) (Fundamental Research Funds for the Central Universities; grant number: 2652023001)”.

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Co-evolving N-Fe redox processes controlled iron minerals in banded iron formation

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Introduction

Response of nitrification and denitrification to temperature cooling

Temperature-driven Fe(II) oxidation coupled with NO reduction

Coupled N-Fe redox cycles for the mineralogical succession of BIFs

Methods

Thermophilic nitrification and denitrification consortia enrichment and analyses

Fe(II) oxidation coupled with NO reduction (chemodenitrification)

Fe(II) oxidation coupled with NO reduction by denitrifying consortia

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