Does Phytoplankton Bloom Management Provide an Opportunity for Greenhouse Gas Mitigation?

Abstract

A growing body of evidence indicates that freshwater bodies, particularly eutrophic systems, are significant sources of the greenhouse gases (GHGs) carbon dioxide (CO₂), methane (CH₄), and nitrous oxide (N₂O). Unlike marine environments, freshwater systems are generally shallower and more directly influenced by terrestrial inputs, including nutrient enrichment, organic matter deposition, and steep redox gradients in both the water column and sediments. These conditions promote intense phytoplankton growth, including massive harmful cyanobacterial blooms (HCBs), and stimulate microbial processes that drive GHG production and release. This opinion article examines the biogeochemical mechanisms underlying these emissions and evaluates the potential of mitigation treatments to both enhance carbon sequestration and reduce CH₄ and N₂O emissions. We argue that effective control of HCBs, whether through nutrient load reduction or direct mitigation protocols, would not only provide communities with toxin-free water but also significantly lower GHG emissions from eutrophic waterbodies. As this is an opinion paper rather than a comprehensive review, we intentionally avoided citing widely accepted concepts, since doing full justice to the many excellent contributions across all relevant subfields would not be possible within the scope of this work.

1. Brief Introduction

The rising atmospheric CO₂ concentrations from pre-industrial levels of about 280 ppm to over 420 ppm today, primarily due to fossil fuel combustion and deforestation, and that of other (methane and nitrous oxide) greenhouse gases (GHG) is widely recognized as a major driver of global warming. Along with other GHGs, elevated CO₂ levels contribute to a range of adverse effects, including ocean acidification, extreme weather events, and loss of biodiversity, all of which have serious implications for both human and ecological systems. International agreements, such as the Paris Agreement (https://unfccc.int/files/meetings/paris_nov_2015/application/pdf/paris_agreement_english_.pdf?utm_source=chatgpt.com, accessed on 20 July 2025) aim to limit the global temperature increase to well below 2 °C above pre-industrial levels. Achieving this target requires not only substantial reductions in ongoing emissions but also active removal of existing GHG from the atmosphere. However, despite global efforts and commitments, the concentration of GHG continues to rise and the trend is accelerating highlighting the urgent need for more effective and immediate action.

Various frameworks have been developed to sequester atmospheric CO₂, broadly classified into two main categories: engineered technologies and nature-based solutions. Engineered approaches such as direct air capture, carbon mineralization, and geological storage typically involve substantial energy requirements, complex infrastructure, and extended development timelines. In contrast, nature-based solutions leverage ecological processes like photosynthesis to remove atmospheric CO₂, generally with lower energy demands and a range of ecological co-benefits. Beyond their role in climate mitigation, biologically driven strategies also support biodiversity conservation, water quality improvement, and ecosystem resilience [1]. Nevertheless, despite the progress achieved through existing initiatives, the total volume of CO₂ sequestered to date remains far below the scale needed to significantly alter the atmospheric carbon trajectory, highlighting the magnitude of the challenge that lies ahead.

1.1. Intensification of Harmful Blooms

One of the consequences of global warming is the intensification of HCBs during the last few decades and their spreading to relatively cold regions where such bloom events were previously unheard of. The blooms are defined as a rapid rise in phytoplankton abundance reaching values higher than 2 × 10⁴ cells/mL or 10 µg chlorophyll a/L. For harmful blooms, readers may refer to the World Health Organization guidelines for relevant toxin thresholds (https://www.who.int/publications/i/item/9789241549950, accessed on 20 July 2025). The blooms capitalize on the elevated dissolved CO₂ concentrations, a rising nutrient supply due to leaching from agricultural fields or poorly treated wastewater and the warming lakes temperatures [2,3,4,5,6] altogether enhancing their photosynthetic activity and enabling them to produce massive, and frequently toxic blooms [4]. These blooms substantially alter the optical and biogeochemical characteristics of the waterbody and their biodiversity, severely affecting the water quality in thousands of waterbodies in both developed and underdeveloped countries and hence the wellbeing of people and wildlife. It is estimated that 4.4 billion people currently lack access to clean drinking water worldwide [7], intensification of the HCBs severely aggravate the problem by adding toxins to the infested water.

In my view (and that of many others often discussed in scientific meetings) it is time we begin to recognize the intensification and global spread of HCBs as a man-made pandemic, much like the way our colleagues in the health sector classify widespread threats to public well-being. As long as these blooms continue to be regarded and hence dismissed as “natural phenomena,” policymakers and the public will remain reluctant to allocate the resources urgently needed for their control or mitigation. The consequence is that communities will go on drinking water laced with cyanotoxins or turning to bottled water, thereby compounding the problem with plastic pollution. Reframing HCBs as a preventable, human-driven crisis is essential if we are to galvanize the political will and societal commitment necessary to restore water safety, protect ecosystems, and safeguard human health.

This raises a fundamental question: if we term HCBs “natural phenomena,” how do we define “natural processes”? This distinction is critical, as many policymakers, environmental agencies, and public service leaders remain reluctant to confront HCBs despite their severe global impacts. While HCBs do occur naturally, extensive laboratory and field evidence demonstrates that their proliferation and intensification, particularly over the past three decades, are largely driven by human activities, including nutrient enrichment and climate change. These pressures have enabled HCBs to expand into previously unaffected, cooler regions such as Scandinavia, Canada, southern Australia, and the southern parts of Argentina and Chile.

Given this context, it is difficult to view today’s intense blooms as entirely “natural”. With the development of effective, environmentally friendly mitigation protocols, it may be possible to partially “cure” affected lakes, restoring ecological stability, biodiversity, and toxin-free water supplies, much like returning them to a pre-bloom state. In doing so, we are not merely intervening, but actively healing these ecosystems, enabling them to regain ecological stability and biodiversity. A mitigation treatment, when implemented, should selectively target the dominant bloom-forming harmful species, particularly Microcystis spp. which are responsible for most ecological disruptions in eutrophic freshwater systems. As discussed below, the impact of the applied mitigation protocol on both the rate and magnitude of biomass sedimentation strongly influences the spatial and temporal patterns of GHGs emissions, primarily through its effects on oxygen availability and the redox potential along the water column. Without detailing specific protocols, one possible approach exploits the high sensitivity of cyanobacteria to oxidative stress. As the earliest oxygen-producing organisms, cyanobacteria are less equipped to cope with oxidative damage than eukaryotes in the phytoplankton assemblage [5]. Mild oxidative stress can trigger programmed cell death-like processes in cyanobacteria without significantly harming other phytoplankton groups [5].

As mentioned, one of the consequences of global warming and HCB proliferation is a dramatic decline in the biodiversity of waterbodies where toxic cyanobacteria may represent over 90% of the photosynthetic organisms. It is widely recognized that, in addition to nutrient availability and various abiotic conditions, both inter- and intra-species interactions, particularly allelopathy, play a significant role in shaping the phytoplankton community composition [8,9,10,11,12]. One example is growth experiments that demonstrated interspecies competition between the dinoflagellate Peridinium gatunense and the cyanobacterium Microcystis aeruginosa which strongly inhibit each other, with the final cell abundance and community composition determined by their initial inoculum ratios and the temperature [13,14]. Since, in most cases, the optimal temperature for dinoflagellate, green algae and diatoms growth is lower than that of toxic cyanobacteria strains, the warming winters plays a major role determining the seasonal succession in freshwater systems and in the intensification of HCBs [15]. This is further aggravated by the larger fraction of cyanobacterial cells that overwinter under warmer conditions [16,17,18] serving as an inoculum for the development of HCB populations in the spring. In our view, the competitive advantage provided by this larger overwintering inoculum is one of the principal drivers behind the intensification of HCBs and their spread into cooler regions over recent decades.

1.2. Shallow Lakes Are Hotspots for HCB

Shallow eutrophic lakes are more vulnerable to massive phytoplankton blooms compared to deeper lakes. Examples include the (many) lakes in the drainage basin of the Yangtze River in China, Lake Okeechobee in Florida, parts of Lake Victoria in Africa and many others. The main reasons being that the nutrients in the sediments are more accessible to the phytoplankton, closer to the photic zone compared with deeper lakes where they may sink to the hypolimnion and become unavailable for the rest of the summer, due to stratification, awaiting for the forthcoming water mixing in the winter. In shallow waterbodies, where the smaller volumes warm up more rapidly in the spring [17], cyanobacteria which thrive in warm conditions gain an early-season advantage, allowing them to establish dominance and outcompete other algae.

Given the strong link between rising atmospheric CO₂ and the growth of HCBs, we examine how mitigation treatments that promote biomass sinking may enhance both organic- and inorganic C sequestration. We also explore strategies to reduce other GHGs emissions from freshwater bodies, focusing on key processes that govern GHG formation and exchange with the atmosphere. Rather than a comprehensive literature review, we highlight mechanisms affected by algal bloom mitigation, as illustrated in the Figure 1.

This approach is especially relevant for eutrophic waterbodies burdened by HCBs, offering a potential pathway to reduce their environmental impact while contributing to climate change mitigation.

2. Sequestration of Organic C

In aquatic systems, organic carbon (OC) is produced as part of the “biological CO₂ pump” Figure 1 and [19,20,21], whereby atmospheric CO₂, dissolved in the sunlit upper layers of the water column, is converted into organic matter through photosynthesis and associated cellular processes. In addition to this autochthonous production, allochthonous organic matter can also enter the waterbody from the surrounding terrestrial environment. A portion of the OC sinks to the sediments, contributing to long-term carbon sequestration, while another fraction undergoes microbial decomposition and respiration, releasing CO₂ and methane (see below) and thereby completing the carbon cycle (Figure 1). One of the central challenges is to better understand the drivers of these fluxes—what determines the balance between them—and whether it is possible to increase the proportion of OC that remains buried over long timescales.

2.1. Are Lakes a Source or a Sink of CO₂?

As mentioned earlier, photosynthetic CO₂ fixation promotes the dissolution of atmospheric CO₂ into waterbodies. Within aquatic environments, CO₂ arises primarily from two processes: (1) the microbial decomposition of OC and (2) calcification, in which two bicarbonate ions combine to form carbonate, releasing CO₂ in the process (Figure 1). Numerous studies have shown that lakes play a dual and complex role in the global carbon cycle, acting alternately as sources and sinks of CO₂ depending on their geographical location, season, time of day, organic carbon inputs, and other environmental factors [3,22,23,24,25].

Eutrophic lakes generally emit much larger quantities of GHGs per unit surface area than oceans (oligotrophic systems), although likely less per cell because of self-shading effects on redox balance. Model-based estimates suggest that, globally, lakes may contribute up to 20% of fossil fuel–equivalent CO₂ emissions [22]. However, because a portion of the photosynthetically produced OC becomes buried in sediments rather than fully mineralized, lakes should, in principle, act as sinks for atmospheric CO₂. Consequently, the observation of a net CO₂ efflux may appear to contradict the second law of thermodynamics, since part of the fixed carbon remains stored instead of being respired. This apparent paradox is resolved by recognizing that the balance between OC burial and its decomposition ultimately determines whether a lake functions as a net CO₂ (and O₂) source or sink. A key factor in this balance is the input of allochthonous OC—terrestrial carbon fixed on land and transported into lakes via runoff [26,27,28]. Once introduced, this external carbon is subject to microbial degradation, contributing additional CO₂ (and methane) emissions. As a result, even systems showing strong in-lake CO₂ drawdown through photosynthetic alkalization and elevated dissolved oxygen (DO) concentrations (see Figure 2) may, on a whole lake scale, exhibit net GHG release.

Figure 2. Vertical profiles of pH (A) and dissolved oxygen (DO) (B) with water depth in three irrigation reservoirs. Nitzanim and Galon are filled with nutrient-rich, treated recycled water, while Masu’ot Itzhak receives floodwater collected from a nearby stream. Aerial images of the reservoirs (Figure 3A–C) show an extensive bloom in Nitzanim, predominantly Microcystis sp., in contrast to Galon (where routine mitigation treatments effectively limit bloom development) and Masu’ot Itzhak where algal growth is constrained by nutrients availability. The pH and DO measurements were performed using a YSI ProDSS Multi-Parameter Water Quality Meter on 4 June 2025.

Figure 2. Vertical profiles of pH (A) and dissolved oxygen (DO) (B) with water depth in three irrigation reservoirs. Nitzanim and Galon are filled with nutrient-rich, treated recycled water, while Masu’ot Itzhak receives floodwater collected from a nearby stream. Aerial images of the reservoirs (Figure 3A–C) show an extensive bloom in Nitzanim, predominantly Microcystis sp., in contrast to Galon (where routine mitigation treatments effectively limit bloom development) and Masu’ot Itzhak where algal growth is constrained by nutrients availability. The pH and DO measurements were performed using a YSI ProDSS Multi-Parameter Water Quality Meter on 4 June 2025.

Figure 3. Aerial photographs of the Nitzanim (A), Galon (B), and Masu’ot Itzhak (C) reservoirs, showing the relative phytoplankton (HCB) intensity, along with corresponding methane emission fluxes (D). Methane emissions were quantified using an ABB GLA131 laser-based gas analyzer connected to a custom-built floating flux chamber of known volume and surface area, measured in the middle of the waterbody.

Figure 3. Aerial photographs of the Nitzanim (A), Galon (B), and Masu’ot Itzhak (C) reservoirs, showing the relative phytoplankton (HCB) intensity, along with corresponding methane emission fluxes (D). Methane emissions were quantified using an ABB GLA131 laser-based gas analyzer connected to a custom-built floating flux chamber of known volume and surface area, measured in the middle of the waterbody.

2.2. Burial Efficiency (BE) of Organic Carbon

Notably, cyanobacteria produce large quantities and a wide variety of extracellular polysaccharides (EPS), which exert significant influence on their environment, including in biological desert soil crusts, rice paddies, and aquatic systems [29,30,31,32,33]. During intensive blooms, EPS facilitate cell-to-cell adhesion and colony formation, promoting the aggregation of phytoplankton into particulate organic carbon (POC) that sinks toward the sediments as “marine snow” in the oceans or “flocs” in freshwater bodies. The size of these particles impacts both their sedimentation rate [34] and internal oxygen dynamics; larger flocs exhibit slower oxygen diffusion, potentially leading to anoxic microenvironments within them. Consequently, anaerobic conditions may develop within the flocs delaying biomass degradation by bacterial respiration. It is not known, and worth exploring, whether anaerobic conditions stimulate methane formation in the flocs since they typically originate in the surface layer where the abundance of methanogens is likely much lower than in the sediments. In addition, the EPS also affect anaerobic decomposition, condensation of the sediment and mineralization [35,36] by aiding sediment compaction and maintaining structure. Their decomposition produces extracellular organic compounds that inhibits degrading bacterial communities delaying the breakdown of settled cyanobacterial biomass and methane formation, at least in rice fields [37,38,39].

2.3. Distinguishing Factors in Carbon BE: Lakes vs. Oceans

One of the long-standing dogmas in the field has been that dispersal, grazing by fish and zooplankton and bacterial respiration consumes a large portion of the aggregates, releasing most of the fixed CO₂, while the particles sink to the bottom and during further degradation in the sediments [3,40]. This is indeed the case in the marine environment where only about 1% of the OC produced in the photic zones reach the sediments at >1 km depth [34,41,42] where it remains buried for geological timescale [34,43]. In contrast, detailed studies performed on fresh waterbodies, lakes and reservoirs, spanning various environmental condition, suggested a much larger BE than in the marine environment [26,44,45]. One example is a study performed by the US Geological Survey on 697 waterbodies in continental USA that concluded that the average BE in water reservoirs was 58% [46].

We must conclude that the long-standing dogma that “lakes’ OC is much less permanent than in oceans” doesn’t withstand empirical evaluation. The results obtained, particularly from eutrophic freshwater bodies show just the opposite. Naturally, the question arises—what distinguishes marine systems from freshwater environments as reflected by the much higher BE in the latter. To clarify this important point, it is essential to reveal the processes driving OC degradation within the waterbody which ultimately determine the OC fraction that reach and then buried in sediments. Briefly, the primary factors influencing these processes are the duration of exposure to oxygen, temperature, and the biomass density.

2.3.1. Exposure to Oxygen and the Temperature Impact

Seminal studies by Sebastian Sobek and colleagues [26,40,45] demonstrated that the BE of POC is strongly influenced by duration of exposure to oxygen during descent to the sediments and by temperature. Prolonged exposure to O₂ and higher temperatures result in lower BE, primarily due to bacterial respiration which consumes OC and nutrients released from lysing cells. Since the oceans are much deeper than lakes the aggregates sinking to the sediments are subjected to grazing and bacterial respiration for a considerably longer duration than in lakes. Further, in shallow lakes, OC particles often reach anaerobic conditions in the sediments within a few meters from the water surface much faster than in the ocean [40,47]. In addition, the massive biomass often observed in eutrophic freshwater lakes, with substantial amount of EPS, forms larger flocs than marine snow. As mentioned, the larger the flocs are the slower their bacterial degradation is.

Furthermore, the strong currents characteristic of oceanic continental shelves where nutrient inputs fuel intensive algal blooms transport substantial amounts of oxygen to deeper coastal layers, thereby enhancing OC degradation. Ultimately, in addition to the effects of temperature and oxygen [40,41], large spatial variability was observed in the amount of OC reaching specific sediment sites [45] likely reflecting the impact of currents and changing wind directions on local algal abundance. In densely populated waterbodies and/or high turbidity, the depth of the photic zone (where net oxygen production occurs) is attenuated, particularly in shallow lakes where turbulence by wind or currents may raise sediment particles [48]. Consequently, anoxic condition may develop within a short distance from the water surface [17].

2.3.2. Cell-Density and Phytoplankton Biomass

As indicated, the biomass density is an important parameter affecting the oxygen gradient profile and hence OC degradation. Although blooms intensity increased significantly in the open, oligotrophic, ocean during the last century [49], the overall cell density of phytoplankton biomass is rather low, equivalent to approximately 0.1–0.3 mg chlorophyll a/m³. It may reach 5–10 mg chlorophyll a/m³ during spring-summer blooms mostly near coastal nutrient-rich areas. In contrast, the cell density in eutrophic freshwater bodies can easily exceed 300–500 mg chlorophyll/m³, particularly during HCB that form scums on the water surface [50,51] and https://portal.edirepository.org/nis/mapbrowse?scope=edi&identifier=1756&revision=2 (accessed on 20 July 2025). Chlorophyll extraction methodologies, along with recent advances in remote sensing technologies, have made it easier and more accessible to assess its concentrations across large areas. Where evaluations of carbon sequestration are required, conversion factors for estimating dry matter from chlorophyll concentrations [52,53] can be employed. However, remote sensing primarily assesses surface phytoplankton populations, while their spatial (including three-dimensional) and temporal distribution is, in most cases, highly variable. Therefore, ground-truthing is essential to accurately calibrate remote sensing measurements. It is timely to develop robust models that integrate remote sensing data with in situ measurements to capture the spatial and 3D distribution of phytoplankton, enabling comprehensive, whole lake biomass assessments. Alternatively, biomass can be estimated by systematically collecting and analyzing samples from multiple locations and depths. However, this process is labor-intensive, costly, and subject to inherent logistical limitations.

2.4. Sequestration of OC in the Marine and Freshwater Bodies

Geologically speaking, it is widely recognized that sediments derived from phytoplankton biomass have played a significant role in the formation of kerogen, a key intermediate in the transformation of OC into natural gas and other fossil fuels. This highlights their crucial contribution to the global carbon cycle over geological timescales. For instance, a recent review [54] proposed a model describing cyanobacterial proliferation and the preservation of organic matter, emphasizing their critical role in OC burial throughout Earth’s history.

This perspective raises a fundamental question: to what extent can the permanence of OC in a given waterbody be predicted based on its environmental conditions and biological activity. The concept of “burial permanence”—the length of time that OC remains sequestered in sediments is inherently complex as it involves understanding and measuring the dynamic processes that govern the stability and longevity of OC [55]. Radiometric dating of sediment cores using isotopes such as ¹⁴C, ¹³⁷Cs, and ²¹⁰Pb is commonly employed to assess the permanence. These analyses provide a detailed reconstruction of the waterbody’s depositional history and have been published in numerous peer-reviewed studies [44,56,57,58,59,60,61], governmental agencies reports (e.g., USGS SIR 2004-5184) and large-scale research initiatives like New Zealand’s Lake380 project (https://ourlakesourfuture.co.nz/, accessed on 20 July 2025). Naturally, the preservation of OC in the sediment is strongly influenced by local environmental conditions. In all cases we are aware of, sediment cores show a pattern of rapid OC degradation during the first few years, followed by a much slower breakdown phase, highlighting a stabilization process in the deeper sediment layers.

The reader is encouraged to explore the Lake380 project available on their website and emerging publications thereof. The data represents one of the most extensive and detailed waterbody analyses, over a range of ambient condition, conducted globally. The sediment cores collected as part of this effort offer invaluable data on both contemporary and historical water quality, encompassing a broad range of biotic and abiotic parameters. In particular, the use of pigment extraction and DNA sequencing along sediment cores serves as a powerful approach for distinguishing between, and reconstructing of algal, shrub and tree-originated biomass over time. Interestingly, the algal originated OC remains fairly constant over tens of thousands of years. This is probably due to various processes taking place in the sediment, where the buried OC undergoes physiochemical processes such as condensation, mineralization, and compaction, ultimately transforming into stable hydrocarbon structures, commonly known as kerogen [37,38,39,54]. Altogether, offering strong evidence for the capacity of freshwater systems to sequester OC over long periods.

2.5. Can Mitigation of HCBs Be Used for Substantial Carbon Burial?

Phytoplankton blooms in eutrophic freshwater systems hold considerable potential for OC sequestration [62]. However, their long-term effectiveness and sustainability remain subjects of active debate. Under business-as-usual conditions, a substantial fraction of bloom-derived OC is mineralized in the water column (Figure 1), thereby limiting BE. In contrast, prolonged exposure to mild oxidative stress has been shown to trigger a programmed cell death (PCD)-like response in cyanobacteria, with mortality rates reaching up to 99% [5,63,64,65,66]. The resulting biomass rapidly aggregates into large flocs that sink to the sediment. As reported by Sobek and colleagues [26,40,45] and discussed in Section 2.2, larger flocs and faster sedimentation rates are strongly correlated with higher BE because they reduce OC degradation by bacterial respiration. A key hypothesis worth testing is that mitigation treatments inducing massive bloom collapse may also create localized anoxia in bottom waters, further limiting microbial respiration. Interestingly, PCD-induced rapid collapse appears to minimize cyanotoxin release into the surrounding waters [63,67,68], though the mechanism is not yet understood. One possible explanation is that many intact cells are captured within the sinking flocs, carrying intracellular toxins and nutrients with them to the sediment bed—a possibility that requires experimental examination.

Over longer timescales, the removal of toxic cyanobacteria facilitates phytoplankton succession, restoring biodiversity and supporting renewed oxygenic photosynthesis even after dramatic declines in cyanobacterial abundance. Another clear benefit of mitigation-induced bloom collapse is the marked improvement in water clarity, which enhances light penetration and thus photosynthetic oxygen production throughout the water column. Taken together, these processes suggest that BE under treated conditions may significantly exceed that under untreated, baseline scenarios.

Researchers involved in carbon harvesting from phytoplankton biomass may utilize commercial systems that measure Total Organic Carbon (TOC) via combustion and CO₂ detection, followed by acidification to quantify Total Inorganic Carbon (TIC) enabling full TIC/TOC profiling across depths and locations. Remote sensing offers a complementary approach by estimating chlorophyll-a as a proxy for biomass. However, assessing carbon flux to sediments post-mitigation remains challenging due to the spatial and temporal variability in HCB biomass, influenced by factors like wind mixing, stratification, light, and nutrients. Addressing this requires an integrated strategy combining remote sensing, in situ measurements, and modeling as mentioned in Section 2.3.2, above. While chlorophyll-a accounts for 0.6–1.6% of dry biomass [69] under laboratory conditions, field values may vary.

3. The Fate of Inorganic Carbon in the Marine- and Fresh-Waterbodies

3.1. The CO₂-Concentrating Mechanisms (CCMs)

Cyanobacteria and algae have evolved sophisticated CCMs that enable them to accumulate Dissolved Inorganic Carbon (DIC) within their cells, particularly under CO₂-limiting conditions [70,71,72,73,74]. They can thus be regarded as DIC containing bags, intracellular levels as high as 50 mM DIC were recorded [72]. These mechanisms allow the organisms to elevate the internal CO₂ concentration in close proximity to their carboxylating enzyme, ribulose-1,5-bisphosphate carboxylase/oxygenase (Rubisco), which is predominantly localized within carboxysomes in cyanobacteria and pyrenoids in algae. This adaptation enables them to overcome the large gap between Rubisco’s relatively low affinity for CO₂ and the low concentration of dissolved CO₂ in equilibrium with the atmosphere, particularly in cyanobacteria, where the Km(CO₂) of Rubisco is approximately tenfold higher than ambient levels. The size and dynamics of the internal DIC pool are strongly influenced by environmental factors, especially the CO₂ concentration during growth and light intensity [70,72,75], and they vary among species and even among strains, including Microcystis sp. [76,77].

Seminal studies by Elke Dittmann, Martin Hagemann, and colleagues [78,79,80] revealed a non-canonical CCM in toxic Microcystis sp., where a significant part of Rubisco is located outside the carboxysomes. This distribution results in a less efficient CCM compared to the classical model cyanobacteria studied previously. Additionally, microcystin, a toxic secondary metabolite produced by Microcystis sp., has been shown to bind Rubisco [81]. Fluctuations in extracellular microcystin levels such as those occurring during cell lysis under stress affect the cytoplasmic localization of Rubisco [79]. These findings support earlier hypotheses [82] and provide a mechanistic framework for how cells might sense extracellular microcystin levels and adjust their physiology to cope with environmental stress, thereby enhancing ecological fitness.

A recent study [83] further demonstrated that a similar non-canonical CCM may also operate in the filamentous cyanobacterium Nostoc punctiforme. Here, a poorly functioning CCM necessitates a reliance on heterotrophic bacteria for CO₂ supply under limiting DIC conditions, a surprising dependency that could explain the difficulty in isolating axenic cultures of various cyanobacteria and suggests that elevated ambient CO₂ levels may be critical for their isolation and cultivation. Finally, a quantitative analysis of the internal DIC (and calcite) pool is essential to fully understand the operation and regulation of non-canonical CCMs and their effect on intracellular DIC accumulation [72]. In addition to enhancing the apparent photosynthetic affinity for CO₂, the substantial intracellular accumulation of DIC plays a pivotal role in calcification processes across various algal groups.

3.2. Calcification Processes

The example in Figure 2 shows that during HCBs, intense photosynthetic activity in the upper water layers, where light is most available, is often accompanied by pronounced alkalization. As CO₂ is rapidly depleted and fixed through photosynthesis, pH levels can rise above 10, indicating a substantial departure of carbonate chemistry from equilibrium with atmospheric CO₂ (Figure 2A). Elevated DO concentrations in surface waters (Figure 2B), far exceeding equilibrium values with air, further confirm that this alkalization is driven by vigorous photosynthesis. Such conditions promote calcium carbonate precipitation both in the water column and on the surfaces of algal species with mucilaginous sheaths, which create localized supersaturation and facilitate calcite deposition [84,85,86,87,88,89,90].

Intracellular CaCO₃ accumulation in bloom-forming Microcystis spp. has been documented [87,91,92], though quantification of calcite content under different environmental conditions remains incomplete [89,93,94,95,96,97,98]. Given the remarkable ability of cyanobacterial cells to accumulate DIC through their CCM [72], they can effectively be regarded as DIC-bearing entities. This highlights an unexplored potential for employing molecular tools to further enhance DIC accumulation and, consequently, carbon sequestration [99,100]. Promising genetic targets for such interventions include bicarbonate transporters (SbtA, BCT1, and BicA) as well as the NDH-1 complexes that facilitate CO₂ recapture from leaking pools [71,101]. Noticeably, the abundance of calcium ions may be limiting calcite formation in freshwater bodies. Nevertheless, while the release of genetically modified organisms into natural environments remains subject to strict regulatory constraints, testing their capacity for enhanced DIC accumulation and cycling [102] under controlled or simulated bloom conditions could provide valuable insights into their potential utility for large-scale carbon capture applications.

Calcification in aquatic systems involves the reaction of Ca²⁺ with HCO₃⁻ to form calcium carbonate, releasing CO₂ (and water, Figure 1). As mentioned, this process can lead to substantial CO₂ emissions from both freshwater and marine environments [89,90,96,103,104]. However, the accumulation of calcium carbonate in coral or coastal reefs and freshwater sediments suggests a net uptake of DIC over time, as atmospheric CO₂ dissolves in water to replenish bicarbonate consumed during calcification. In freshwater bodies, significant calcification often occurs in areas with high alkalinity, hard water or during substantial phytoplankton blooms where intensive CO₂ uptake raises the pH to values where most of the DIC is in the form of carbonate, promoting calcite precipitation [105]. The Sea of Galilee (Lake Kinneret), a stratified hard-water lake, provides a notable example. It was estimated that over 52,000 tons of calcite settles into the sediment during the spring blooms of the dinoflagellate Peridinium gatunense [93,106]. Mineralization processes within the sediment also lead to carbonate formation particularly during intensive algal blooms [107]. Here, calcification is driven by the high level of dissolved CO₂ consequent on intensive bacterial respiration. It is estimated that carbonate formation through mineralization at the sediment accounts for 3–8% of the deposited OC. However, distinguishing this carbonate from that settling onto the sediments from the upper water layer remains challenging [35]. Attempts to resolve this issue using δ¹³C would be complicated by the disequilibrium between CO₂ and bicarbonate species at both the sediment surface and in the photic zone. This disequilibrium, driven by a substantial pH gradient, significantly shifts the δ¹³C values of these species away from the equilibrium value of approximately 7‰ [108].

Naturally, the amount of carbon buried through calcification is strongly influenced by various factors, with calcium, magnesium and proton concentrations playing a particularly significant role. These concentrations vary widely between different lakes and depth in the water column (Figure 2A) and affect both calcification and dissolution processes. Unfortunately, detailed analyses of the amount of carbonate reaching the sediments, whether due to calcification in the epilimnion or mineralization of OC, and its subsequent fate, including the extent of dissolution [104], remain scarce. Direct measurements of descending carbonates using sediment traps and cores provide valuable insights, though these techniques are labor intensive and costly with substantial variation observed among replicates. Alternatively, an assessment of calcium and magnesium cycles within the waterbody through measurements of their levels in the epilimnion, hypolimnion, as well as inflowing and outgoing streams throughout the year offers a robust approach to evaluate calcification and dissolution dynamics. In the case of Lake Kinneret, analysis [109] revealed that approximately one-third of the calcite formed during the spring bloom undergoes dissolution, while the remaining portion remains buried (Nishri, personal communication). Such assessments are, in our view, crucial for evaluating carbon fluxes in aquatic environments, understanding their role in the global carbon cycle, and exploring potential pathways for effective carbon sequestration (Figure 1).

3.3. A Challenge: Can We Use Changing pH to Assess the Amount of Atmospheric CO₂ Dissolution in Freshwater Bodies?

Dissolution of atmospheric CO₂ to replace that consumed during OC formation is commonly referred to as the “biological pump” (see Section 2). In a similar vein, the accumulation of calcite and other carbonate minerals in sediments, driven by water alkalization as described above, may be regarded as a complementary abiotic process. The extent of pH increase during intensive CO₂ uptake by photosynthesis (Figure 2A) is strongly influenced by the alkalinity of the waterbody and the chemical composition of its buffering system. Higher alkalinity results in a smaller pH rise for a given amount of CO₂ removed. In principle, alkalinity can therefore be used to estimate the amount of CO₂ required to return the pH across a given range, including to pre-bloom conditions, assuming equilibrium among DIC species throughout the process. However, this equilibrium-based assumption can be misleading. Unlike oceanic systems, eutrophic lakes often exhibit significant disequilibrium among DIC species, driven by dynamic biological and chemical processes, most notably the intensive removal of CO₂ during photosynthesis and the precipitation of carbonates at elevated pH (Figure 1). The latter process is especially pronounced when calcium and magnesium ions are sufficiently available leading to effective removal of DIC from the water column further complicating efforts to model or quantify DIC speciation accurately (Figure 1). To address this complexity, we propose a time-resolved modeling approach based on observed pH dynamics, incorporating non-equilibrium processes that may provide a more accurate framework to estimate CO₂ fluxes.

When HCB is abruptly terminated (e.g., through a mitigation treatment) the cessation of photosynthetic CO₂ uptake permits dissolved CO₂ in the water to re-equilibrate with the atmosphere, typically resulting in a gradual decline in pH toward pre-bloom levels. The rate of atmospheric CO₂ dissolution is influenced by abiotic drivers such as temperature, wind, turbulence, and vertical mixing. Additional contributors to the rising CO₂ concentration include the balance between respiratory CO₂ production (B) and the photosynthetic uptake (C, notably, C is expected to decrease after a mitigation treatment, leading to alterations in the other fluxes), and shifts in the distribution of DIC species as the system rebalances with changing pH (D). The relationships between these components can be described as:

A = (B − C) + D + E

where (A) is the amount of CO₂ required to reduce the pH over a given range and E is the net amount of atmospheric CO₂ dissolved in the water towards the new steady state reached following the mitigation treatment.

A key challenge is whether we can quantitatively distinguish between the contributions of each of these processes, in order to retrieve the net amount of CO₂ dissolved from the atmosphere following bloom collapse. Potential approaches include measuring the initial rate of pH decline in sealed water containers (excluding air exchange) may provide insight into non-atmospheric CO₂ contributions (e.g., from respiration and internal redistribution of DIC species). Quick time points measurements are critical since prolonged incubation can result in substantial pH reduction by microbial respiration alone, especially where oxygen and organic carbon are available. A more robust method involves quantifying DIC directly by acidifying collected samples, which converts all carbonate species to CO₂ for measurement. This bypasses the reliance on pH-alkalinity modeling alone. We acknowledge that spatial heterogeneity, both horizontal and vertical, and temporal fluctuations in pH and alkalinity (see Figure 2A) are inherent to the dynamics of bloom development and collapse. These variations are also reflected in local DO levels (Figure 2B), which integrate the combined effects of photosynthesis, respiration, and gas exchange with the atmosphere. To account for these variabilities, representative sampling across depth profiles, followed by volumetric integration over the entire waterbody, is recommended.

4. The Methane Enigma

The escalating methane (CH₄) level in the atmosphere is a matter of growing concern due to its projected impact on global warming, surpassing CO₂ (per mol) by approximately 30-fold. It is now firmly established that waterbodies constitute significant contributors to atmospheric CH₄ (Figure 1 and [110,111,112,113,114,115,116,117,118]). Large spatial and temporal variability in methane (CH₄) emissions has been reported in studies assessing fluxes from freshwater systems. For example, Hounshell and colleuges [119] and Waldo et al. [120], using the eddy covariance approach, showed that under business-as-usual conditions, eutrophic lakes release between 1.3 and 51.4 g CH₄ m⁻² yr⁻¹. This technique, which measures vertical gas fluxes (such as CO₂ and CH₄) by monitoring high-frequency fluctuations in gas concentration, wind speed, and direction above the ecosystem, provides valuable insights into real-time emission dynamics. However, its application is currently limited to relatively small and homogeneous waterbodies, where the spatial variability in turbulence and gas exchange can be adequately captured. Buoyant flow chambers can also be used to measure CH₄ emissions through detection of its infrared absorption bands or by gas chromatography [121]. However, due to their small surface area, a major challenge is capturing the fraction of CH₄ that escapes via ebullition directly into the atmosphere. Addressing this requires multiple spatial measurements and subsequent modeling to estimate the total CH₄ flux across the water surface. For a more detailed mechanistic understanding of CH₄ production and consumption dynamics (see below), it is advisable to conduct simultaneous measurements of dissolved CH₄ and O₂ concentrations, ideally complemented by redox potential profiling, along the water column.

4.1. Methane Biosynthesis

The mechanisms of CH₄ biosynthesis and degradation are being resolved with recent findings challenging longstanding paradigms and potentially unveiling new strategies to mitigate CH₄ emissions. Biogenic CH₄ is primarily produced under anaerobic conditions by several archaea groups known as methanogens. They can combine CO₂ and hydrogen to form CH₄ in a reaction that consumes the OC produced within the lake, as well as from allochthonous inputs from its surroundings. Additional important CH₄ producing pathways include methylotrophic bacteria that can convert methanol to methane and CH₄ production by phytoplankton under oxic condition (discussed below). In addition to O₂ level, CH₄ formation is also affected by temperature though in some casees less than expected due to the presence of various methanogens whose optimal activities vary across a wide temperature range.

Biogenic CH₄ has the lowest isotopic carbon ratio compared to other natural sources because it is extremely depleted in ¹³C [122]. Its δ¹³C value, typically ranges from −50‰ to −110‰ (relative to the Vienna Pee Dee Belemnite standard, VPDB), is the main diagnostic tool that distinguish it from other CH₄ sources. This highly depleted isotopic signature is due to the preferential use of the lighter carbon isotope (¹²C) by methanogenic archaea during microbial methanogenesis, particularly in the reduction of CO₂ or acetate to CH₄ via the two main pathways involved. However, studies utilizing CH₄ to assess the efficacy of CH₄-derived carbon as a substrate for phytoplankton [123,124] should bear in mind that ¹³C discrimination also occurs during CO₂ fixation, and that the CO₂-concentrating mechanism in phytoplankton raises the utilization of ¹³C by the carboxylating enzyme [72].

The Methane Paradox

Until recently, the prevailing belief was that under business-as-usual conditions, most CH₄ produced in anoxic sediments by archaea (see Section 6, below) is consumed by methanotrophs in the overlying oxic layers. This paradigm was supported by laboratory experiments and studies in small-scale aquatic systems, but recent research challenges this view. A growing body of evidence indicates that various phytoplankton inhabiting the photic, oxygenated water layers, including cyanobacteria, can produce significant amounts of CH₄ during photosynthesis [112,125,126,127,128,129,130,131,132,133]. This aligns with earlier studies on aerobic CH₄ biosynthesis by bacteria [134].

Collectively, these findings contribute to the recognition of the “methane paradox,” wherein CH₄ concentrations in the oxic surface layers of lakes and coastal ocean regions rise toward the top water layer and can surpass levels expected at saturation with air [124,132,135]. The specific organisms involved in freshwater and oceanic methane production are being identified [131,135,136], and the exact metabolic pathways are under intensive investigation. Recent studies have demonstrated that CH₄ formation is light- and photosynthesis-dependent, though the precise mechanisms remain poorly understood [131]. We hypothesize that CH₄ production by photosynthetic organisms under oxic conditions may serve as a “safety valve” for dissipating excess reducing equivalents generated in the light energy harvesting machinery under high illumination or other stress conditions. If so, the inactivation of genes involved in CH₄ biosynthesis would be expected to increase cellular sensitivity to high illumination and to alter both the redox balance and the turnover dynamics of photosynthetic components.

Stable carbon and hydrogen isotope analyses suggest that methylated compounds such as methylphosphonate (MP), methylamine, and methionine serve as precursors for oxic CH₄ production. This process involves a gene cluster (phn) that is widespread in filamentous cyanobacteria and also present in certain strains of toxic Microcystis sp., such as those collected during a large bloom in Dianchi Lake, China [137], emphasizing the role of cyanobacteria in oxic methane formation. Complicating matters further, studies suggest that MP released by lysing cyanobacteria is utilized by bacteria as a phosphate source, thereby releasing CH₄ in the photic zone [137]. Mutations in the C-P lyase phosphonate degradation pathway of the marine bacterium Pseudomonas stutzeri disabled its ability to produce methane [135]. Currently, accurate assessments of the relative contributions of cyanobacteria and algal species to CH₄ formation and utilization in the upper aerobic zone are lacking, highlighting the need for genome screenings, identification of potentially involved genes [130] and physiological flux analyses.

4.2. Methane Consumption

Methane is consumed by methanotrophic microorganisms, which utilize it as both a carbon and a reducing power source, thereby mitigating its emission to the atmosphere. The activation of the stable C-H bond in CH₄ is thermodynamically unfavorable, requiring a high redox potential and the transfer of eight electrons [138]. Methane oxidation proceeds via two main pathways: aerobic and anaerobic. These pathways differ significantly in their enzymatic mechanisms, the organisms involved, and their sensitivity to environmental factors such as temperature and redox potential. In both systems, CH₄ is oxidized by forms of methane monooxygenase (MMO) or analogous enzymes, each exhibiting distinct kinetic properties. In aerobic methanotrophs, MMOs utilize molecular oxygen as the terminal oxidant. Here, the eight-electron transfer is mediated by specific components of the MMO complex, which catalyzes the conversion of CH₄ to methanol, a key step in the metabolism of CH₄ by aerobic methanotrophs. During anaerobic methane oxidation, CH₄ is ultimately converted to CO₂ and water through the action of microbial consortia, typically involving anaerobic methanotrophic archaea in partnership with sulfate-reducing bacteria. Here, alternative oxidants such as Fe³⁺ or copper, and in marine systems, sulfate or nitrate, often serve as terminal electron acceptors for methanotrophic respiration. Where sulfate availability is limited, such as in freshwater lakes, alternative electron acceptors like nitrate, iron or manganese can be utilized. Theoretically, an ORP above approximately +400 mV is likely required to support meaningful anaerobic methanotrophic activity. For reference, the redox potentials of key electron acceptors used in anaerobic processes, Fe³⁺/Fe²⁺ and NO₃⁻/NO₂⁻ are +770 mV and +940 mV, respectively. These alternative pathways enable energy generation that supports methane oxidation under oxygen-depleted conditions, contributing to CH₄ removal even in anoxic zones. Hypothetically, an addition of coarse-grained iron ores (or nitrate), designed to sink to anoxic sediment layers, could enhance CH₄ consumption by stimulating Fe³⁺—dependent methanotrophic activity and thereby reduce overall methane emissions. Repeated applications may be necessary to sustain elevated levels of iron-reducing methanotroph populations and maintain their metabolic associations that enable the process. It is important to note that in most aquatic systems studied to date, nitrate is often depleted in the oxic (photic) zone but remains abundant in deeper, anoxic layers of the water column. Therefore, measuring the redox potential at the sediment-water interface is recommended to assess the potential for anaerobic CH₄ consumption and of the impact of iron fertilization on anaerobic methane oxidation. In contrast, the redox potential required for aerobic CH₄ oxidation is provided directly by oxygen, acting as the terminal electron acceptor during bacterial respiration. Thus, where possible, simultaneous measurements of ORP and DO are highly recommended to better understand the methanotrophic potential and to evaluate strategies for enhancing methane removal in aquatic systems.

As previously noted, cyanobacteria are increasingly recognized as significant contributors to CH₄ production. However, despite CH₄ being a potential source of both carbon and reducing power and despite cyanobacteria’s capacity to support eight electron transfer processes such as nitrogen fixation, they appear to lack the MMO enzyme necessary for cleaving the stable C-H bond in CH₄. To the best of our knowledge, no published studies have demonstrated the growth of axenic cyanobacterial cultures using CH₄ as the sole carbon source. Nevertheless, growing interest in cyanobacteria-methanotroph interactions has led to co-culture studies that reveal enhanced CH₄ oxidation in the presence of cyanobacteria. This enhancement is likely attributable to increased local oxygen availability provided by cyanobacterial photosynthesis. Conversely, methanotroph presence has been associated with accelerated cyanobacterial growth [123,139]. The precise metabolic contribution of methanotrophs to cyanobacterial physiology remains uncertain. In our opinion, given the highly efficient CCM characteristic of cyanobacteria, it seems unlikely that the observed benefits arise solely from respiratory CO₂ release by methanotrophs, as has been previously suggested. This hypothesis could be further evaluated by experimentally increasing the ambient CO₂ concentration.

In the absence of extensive data from freshwater systems, insights may be drawn from studies on cyanobacteria–methanotroph interactions in rice paddy soil microcosms. In such systems, inoculation with cyanobacteria has been shown to reduce CH₄ emissions by up to 90%, with outcomes strongly dependent on the cyanobacterial species used. For instance, Nostoc sp. and Calothrix sp. were found to influence different methanotrophic communities, leading to variable outcomes in CH₄ flux [37], These emerging studies reveal a more complex and nuanced picture of CH₄ dynamics within microbial consortia, underscoring the importance of species-specific interactions, possibly allelopathic mediated by secondary metabolites, in shaping these processes. Continued investigation of these relationships is essential for elucidating their ecological significance and for identifying potential applications in CH₄ mitigation strategies.

4.3. Can HCB Mitigation Reduce Methane Emissions?

Naturally, the net flux of CH₄ emissions from waterbodies is governed by the balance between its production and consumption. The significant variability observed in CH₄ emission fluxes across different eutrophic lakes likely reflects differences in local conditions and various parameters affecting this balance. Studies have shown that cyanobacterial blooms can enhance methane emissions through multiple pathways. For instance, research on Lake Tai, China, showed that regions with dense cyanobacterial populations emitted far more CH₄ compared to low-density areas, 867 vs. 3.4 μg CH₄ m⁻² min⁻¹, respectively, indicating a strong correlation between bloom intensity and methane production [139]. This is an important observation, as it may point to a potential route for decreasing global methane emissions from eutrophic lakes through proper management and control of HCBs.

Three water reservoirs located in the northwestern Negev region of Israel were selected to examine this possibility—Nitzanim (5 ha, Figure 3A) and Galon (6 ha, Figure 3B) are filled with tertiary treated recycled nutrient-rich water, whereas Masu’ot Itzhak (8 ha, Figure 3C) is supplied with water collected from occasional winter floods in a nearby stream. Aerial photographs taken by drone clearly show a dense phytoplankton population in Nitzanim, while phytoplankton presence is barely visible in Galon and Masu’ot Itzhak. Routinely mitigation treatment applied in Galon Reservoir effectively suppressed the HCB that are typically observed in Nitzanim. This difference is further supported by the higher surface pH (Figure 2A) and DO concentrations (Figure 2B) in Nitzanim, likely resulting from intensive photosynthetic activity associated with the denser HCB population. Additionally, the steeper vertical decline in both pH and DO in Nitzanim suggests stronger light attenuation, and thus reduced oxygen production at depth, caused by the bloom’s density.

Measurements of methane flux (Figure 3D) revealed substantially higher CH₄ emissions in Nitzanim compared to the other two reservoirs. It remains to be determined whether the lower emissions observed in Galon and Masu’ot Itzhak are primarily due to reduced CH₄ production by cyanobacteria or to a lower availability of organic carbon in the sediments, which serves as a substrate for methanogenesis. Nevertheless, these preliminary findings are encouraging and underscore the need for further research to elucidate how local environmental factors shape the balance between methane production and consumption, ultimately influencing net emissions. Further, the emerging correlation between eutrophic conditions, bloom intensity, and methane release suggests that bloom-targeted mitigation strategies, such as those employed in Galon Reservoir, may offer an effective approach to reduce CH₄ (and N₂O?, see below) emissions from eutrophic waterbodies.

5. Nitrous Oxide in Freshwater Bodies

In addition to its detrimental effects on the ozone layer, nitrous oxide (N₂O) is a potent GHG, with a global warming potential approximately 270 times greater than that of CO₂. In recent years, freshwater lakes have been recognized as significant sources of atmospheric N₂O emissions [22,140,141,142]. Cyanobacteria and microalgae have been implicated in N₂O biosynthesis in aquatic environments [143,144,145,146].

The primary contributors to N₂O production in waterbodies are microbial processes, particularly ammonia oxidation and denitrification. In oxic waters, ammonia-oxidizing Archaea and Bacteria release N₂O as a byproduct during the oxidation of ammonia to nitrite. Under oxygen-depleted conditions, N₂O also serves as a key intermediate in the denitrification pathway, in which nitrate and nitrite are sequentially reduced to molecular nitrogen (N₂), often coupled with the oxidation of organic matter.

Phytoplankton blooms, and mitigation treatments thereof, can significantly influence N₂O dynamics by altering the structure and extent of oxic-anoxic transition zones (Figure 2D), such as at the sediment-water interface or within stratified water columns. These zones are critical sites for microbial nitrogen transformations, and shifts in redox gradients can determine whether the dominant end-product of nitrogen metabolism is benign N₂ or climate-active N₂O.

Despite its environmental significance, mechanistic understanding of N₂O synthesis and degradation in aquatic systems is poorly understood. Recent research is unfolding the complex interplay of biotic and abiotic pathways that govern N₂O fluxes [147]. Gaining a deeper understanding of these mechanisms is essential for improving nitrogen cycle models in eutrophic waterbodies and developing strategies to mitigate N₂O emissions from inland waters. Interestingly, recent studies have shown that sunlight can drive abiotic N₂O production in both freshwater and marine systems. This occurs primarily through photochemical reduction of nitrate to nitrite, followed by subsequent reactions leading to N₂O formation [148]. However, the precise mechanisms and global relevance of this pathway are still under investigation.

Of note, some methanotrophs are capable of anaerobic methane oxidation using nitrate or nitrite as electron acceptors, effectively reducing one GHG (CH₄) while potentially producing another, far more harmful (N₂O) through a synergistic metabolic process. As with methane, it remains to be experimentally determined whether mitigation treatments targeting algal blooms which can significantly reduce organic matter and alter redox dynamics, will also lead to measurable reductions in N₂O emissions to the atmosphere, as expected.

6. Concluding Remarks

While mitigation strategies targeting HCBs hold considerable promise for reducing GHGs emissions, a more comprehensive understanding of their impacts, underlying mechanisms, and environmental constraints remains essential. Beyond the specific biogeochemical pathways involved, three overarching mechanisms appear to link HCB management to the mitigation of GHG emissions from lakes:

1.

GHG reduction: Elevated cyanobacterial biomass is linked to excess GHG emissions; thus, effective bloom control that decreases overall blooms is likely to reduce these fluxes.

2.

Biomass sinking and carbon burial: Some mitigation approaches promote rapid sinking of cyanobacterial biomass, enhancing burial of organic and inorganic carbon in sediments and effectively removing carbon from the atmosphere.

3.

Community composition shift: bloom-forming cyanobacteria possess physiological traits that favor GHG production. Shifting community composition toward non-cyanobacterial dominance may reduce emissions, regardless of the specific mitigation technique applied.

As noted above, both organic and inorganic carbon sequestration, as well as CO₂ and CH₄ emissions, vary substantially among water bodies, while information on N₂O emissions remains scarce. Ongoing research aims to study GHG fluxes under diverse environmental conditions to better quantify freshwater contributions to global warming and to identify key controlling factors. Seminal studies by Tan, Zhu, and Grasset and colleagues [149,150,151] have provided robust models for simulating CH₄ and carbon dynamics, including temperature-dependent variations in their relative fluxes. These tools enable quantitative assessment of the environmental and biological factors shaping carbon cycling and will help refine mitigation strategies and guide the development of targeted, science-based interventions across aquatic ecosystems.

Finally, recent insights into the modification of methyl–coenzyme M reductase, the key enzyme in archaeal methanogenesis [152], provide a valuable framework linking environmental drivers to methanogen physiology and isotopic signatures of biogenic methane. In this framework, factors such as temperature, pressure, substrate availability, and isotopic gradients interact through partitioning of the available ΔG among enzymatic steps. These interactions collectively determine enzymatic reversibility, growth rates, and isotopic outcomes. It remains to be seen whether such mechanistic understanding, and potential manipulation of ΔG partitioning, can be harnessed to more effectively control GHG emissions under natural conditions.