The JRE basin (4° 23′ S and 37° 43′ W; 4° 36′ S and 37° 43′ W) has a surface area of about 1,350 km², located in a tropical region near the equator (Dias et al., 2009; Figure 1). The watershed of the Jaguaribe River that covers 74,327 km² with an extension of about 610 km contributing to the Western Equatorial Atlantic Ocean (Dias et al., 2009; Cavalcante and Cunha, 2012). Historically, the freshwater discharge into the estuary ranged from 0 to 7,000 m³ s^–1 (Campos et al., 2000). However, the freshwater inputs have decreased due to the construction of a series of dams at the Jaguaribe River watershed (Dias et al., 2009). Historical monthly rainfall shows a marked seasonal behavior, with a wet period between March and April (200–400 mm), and a dry period between August and November, when the precipitation can frequently be zero. Between 2012 and 2017, a persistent extended drought (<700 mm.yr^–1) dominated the river basin climate, and in 2018 annual rainfall started to return to the historical average (FUNCEME, 2021). Indeed, there is a decreasing continental runoff in the semi-arid northeastern region of Brazil as the result of decreasing annual rainfall, and this alarming scenario of decreasing freshwater supply have worsened due to river damming (Lacerda et al., 2020). The water residence time during the dry season is longer in the higher/middle estuary with average of 3 days, and up to 13 days (Lacerda et al., 2013). The estuary has a semi-diurnal and meso-tidal regime with average tidal height of 2.8 m, with peaks during spring tide reaching up to 3.4 m (Dias et al., 2009). The human population in the estuarine basin is estimated at about 100,000 inhabitants, whereas the wastewater treatment services comprise less than 40% of the households (IPECE, 2017). This has contributed to increasing levels of nutrients in the estuary, and occasional occurrence of hypoxia particularly in mangrove channels of the upper/middle estuarine regions (Eschrique et al., 2014; Marins et al., 2020). The trophic status in the estuary, using the trophic state index (Lamparelli, 2004; Silva, 2019), was classified as mesotrophic to eutrophic during the sampling period of this study. In addition to urban wastewaters, discharges from about 3,640 ha of shrimp farms significantly increases nutrient loads into the estuary (Lacerda et al., 2021). The estuary is surrounded by mangroves, which occupy about 13,000 ha (Godoy et al., 2018).

Two eulerian time series were performed, one in October-2017 (33 h of sampling), and other in September-2018 (44 h of sampling) (Figure 1). Both samplings were performed during spring tides and dry season, and at the same moored sampling station. The monthly averaged precipitation during these two surveys were null, and the water budget was negative, i.e., evaporation exceed rainfall rates (Figures 1B,C). The moored station was chosen because it represents a transitional region in the central estuary, which receives influences from continental and marine intrusion and in dry season often represents a turbidity maximum zone, where organic matter interactions and DIC retention occur (Dias et al., 2016; Cavalcante et al., 2021).

The collection campaign included continuous and discrete sampling. The continuous and real-time measurements were performed for the partial pressure of CO₂ (pCO₂), salinity, temperature, and wind velocity. The continuous measurement system is based on Pierrot et al. (2009), and well-described in Carvalho et al. (2017) and Cotovicz et al. (2020b). Briefly, a water pump was placed at a depth of ∼0.5 m and provided continuous water flow (∼2.5 L min^–1) to the boat. The water flow was directed to a thermosalinograph (SeaBird Electronics ^®) to record the surface temperature and salinity. After, passing through the thermosalinograph, this water flux was directed to two equilibrators (shower head type) and then discharged. These showerhead equilibrators promote fast equilibration between air and water phases inside the system. The gas, free from humidity, passes through a Non-dispersive InfraRed gas analyzer (NDIR) for CO₂ quantification (Licor-7000 ^®CO2/H2O gas analyzer). A data acquisition system determined the following parameters every 5 min: date and time, position of the ship, velocity of the ship, molar fraction of CO₂ in the equilibrator (xCO₂), water content in the detector, sea surface temperature (SST) and sea surface salinity (SSS). The molar fraction of CO₂ (xCO₂ ppm) measured in the equilibrator in dry gas was computed using the equation below:

where Peq is the pressure in the equilibrator (assumed to be the same of atmosphere), and P_weq is the pressure of water vapor (atm), according to Weiss and Price (1980). The temperature measured in the surface water and in the equilibrator were slightly different, and then a correction was applied to compensate such difference, according (Takahashi et al., 1993):

where pCO₂ represents the seawater pCO₂ at in situ conditions, SST is the temperature measured in situ and Teq is the temperature measured in the equilibrator.

The NDIR was calibrated prior the first analyses and after every 6 h of seawater molar fraction records using air free CO₂ by passing N₂ standard followed by standard gases with nominal concentrations of 360, 1,009, and 2,009 ppmv (99.9% purity, supplied by White Martins Certified Gases).

Atmospheric pCO₂ measurements (pCO_2air) were performed every 6 h with air taken from the top of the vessel at ∼10 m high. The accuracy of the pCO₂ measurements was estimated at ± 2 μatm. One anemometer model Davis S-WCF-M003 was used to measure the wind velocity and placed at about 10 m in height.

Discrete water samples, at a depth of ∼0.5 m, were collected hourly using a 3-L Niskin bottle. The water samples were filtered in the boat using Whatman GF/F filters (diameter 0.47 mm, pore size 0.7 μm), which were used for chlorophyll a (Chl a) analysis, and the filtrate used for nutrients (phosphate) and TA analysis. All filters were pre-combusted (at 500°C for 6 h). The filters and the filtered water samples were conditioned (fixed and/or maintained in ice in the dark) for further analysis in the laboratory. At the laboratory, the Chl a was extracted in 90% acetone and determined by spectrophotometry following the procedures described by Jeffrey and Humphrey (1975). The phosphate (PO₄^3–) was determined by colorimetric method according to Hansen and Koroleff (1983). TA was determined on 60 mL of filtrate using the Gran (1952) electro-titration method with an automated titration system (Mettler Toledo model T50). The reproducibility of TA was about 3 μmol kg^–1 (n = 7). Measurements were compared to certified reference material (CRM, provided by A. G. Dickson from Scripps Institution of Oceanography) and consistent at an accuracy level of ± 5 μmol kg^–1.

The pH (at the National Bureau of Standards scale, NBS), DIC, Ω_cal, and Ω_ara were calculated from pCO₂, TA, seawater temperature, and salinity using the CO2calc 1.2.9 program (Robbins et al., 2011). The dissociation constants for carbonic acid were those proposed by Mehrbach et al. (1973) refitted by Dickson and Millero (1987), the borate acidity constant from Lee et al. (2010), the dissociation constant for the HSO₄^– ion from Dickson (1990) and the CO₂ solubility coefficient of Weiss (1974). The Ksp values for aragonite and calcite were taken from Mucci (1983) and the concentrations of calcium (Ca²⁺) were assumed proportional to the salinity variations according to Millero (1979).

According to Jiang et al. (2008), in estuaries and lagoons with low or negligible freshwater inputs, the concentrations of DIC and TA during estuarine mixing can be calculated as:

Where SSS_measured is the measured surface salinity, SSS_ocean the surface salinity of the ocean endmember, DIC_ocean and TA_ocean the DIC and TA concentrations of the ocean endmember. This model assumes that the DIC_conservative is the DIC concentration after the ocean endmember is diluted by a zero DIC freshwater. Here, we applied this same approach; however, we assumed that the ocean endmember is linearly concentrated by the evaporation considering the salt conservation. This same approach was recently applied by Cotovicz et al. (2021) studying an evaporative coastal lagoon. The salinities in the inner shelf off the Jaguaribe River during dry season are relatively constant (∼36.7); typical variabilities are < 1 in the ocean endmember (Dias et al., 2013).

The deviation from conservative mixing (ΔDIC) is defined as the DIC addition or loss relative to the theoretical DIC concentration during mixing:

Where DIC_measured is the measured DIC. In the same way, the deviation of TA from the evaporation path (ΔTA) can be calculated. As we did not measure the ocean endmember, we use the value from Cotovicz et al. (2020b), which measured TA and DIC in an adjacent coastal region with similar salinity and water temperature and during the dry season.

The air-water CO₂ fluxes (FCO₂) was calculated according:

where k is the gas transfer velocity coefficient (cm.d^–1), K₀ (mol.cm^–3.atm^–1), is the solubility coefficient of CO₂ at in situ temperature and salinity (Weiss, 1974), and pCO_2water and pCO_2air (pressure) are the partial pressures of CO₂ in equilibrium with surface water and in the overlaying air, respectively.

The gas transfer velocity was parameterized as a function of wind speed taking account two parameterizations available for estuaries (Raymond and Cole, 2001, RC01; Jiang et al., 2008, J08), and two specific parameterizations for oceanic waters (McGillis et al., 2001, M01; Wanninkhof, 2014, W14). The parameterizations of RC01, M01, J08, and W14 can be calculated as follows:

Where U represents the measured wind velocity (m s^–1), Sc is the Schmidt number for CO₂, and 660 is the Schmidt number of CO₂ in seawater at 20°C.

The NCP was calculated considering the changes in dissolved inorganic carbon (DIC) with time, comparing consecutive peaks of maximal and minimal tidal heights. For highest tidal heights, the NCP_HT was calculated according:

Where NCP_HT is in (mmol m^–2 h^–1), ρ is the seawater density (kg m^–3), d is the average depth (m) of the area, t represents the time interval (12 h), and FCO₂ is the averaged carbon dioxide flux (mmol m^–2 h^–1) across the water–atmosphere interface during this period. DIC_1HT and DIC_2HT represent the salinity-normalized concentration of DIC (mmol kg^–1) during two consecutive measurements at the highest tide. In this manner, we can compare consecutive DIC measurements at the maximal tidal level and in different diel period (daytime and night-time). Furthermore, we calculated the NCP for the lowest tidal heights (NCP_LT) according:

Where DIC_1LT and DIC_2LT represent the salinity-normalized concentration of DIC (mmol kg^–1) during two consecutive measurements at the lowest tide. The averaged values of NCP_LT and NCP_HT provide the NCP of the estuary.

The Shapiro-Wilk test was applied to check whether a given variable follows a parametric or a non-parametric distribution. All investigated variables in this study were not normally distributed; therefore, we only applied non-parametric statistics. To compare two unpaired groups, we calculated the Mann-Whitney test. To compare three or more unmatched groups, we calculated the Kruskal-Wallis test. Spearman rank coefficient was calculated to assess the statistical correlation between the rankings of two variables. All statistical analysis were based on significance level of 0.05. We used the GraphPad Prism 8 program (GraphPad Software, Inc., La Jolla, California) to perform graphs and statistical tests.

Changes in carbonate chemistry in semi-arid, hypersaline estuaries are still overlooked (McCutcheon et al., 2019; Yao et al., 2020). Hypersalinity occurs frequently in the JRE during dry months because evaporation exceeds water inputs (Dias et al., 2013). This favors water retention by tidal forcing which increases water residence time in the estuary and enables sediment trapping in the middle estuary (Dias et al., 2013). Water in the middle estuary is blocked for long periods before reaching the sea (Lacerda et al., 2020). Negative water balance is reflected in higher salinities in the estuary (averaging 39.5) compared to the adjacent coastal ocean (<36.5; Cotovicz et al., 2020b; 36.45 ± 0.24 along the equatorial coast; Carvalho et al., 2017). This corroborates previous results in this estuary that exhibits negative estuarine circulation (Dias et al., 2009). Indeed, the JRE behaves as a retainer of DIC during the dry season (Cavalcante et al., 2021). The evaporation of seawater increases the salinity in the remaining seawater (evaporation-concentrated seawater) in the same proportion to the amount of seawater that is evaporated, considering that precipitation and evaporation create no salt flux across the air-sea interface. Thus, it is possible to calculate the effect of evaporation in the concentrations of TA and DIC in the estuary. The values of ΔTA and ΔDIC were mostly positive (above the evaporation line), evidencing that estuary is producing TA and DIC in significant amounts by oxic and sub-oxic processes (Figure 5). The estimated net gains of TA and DIC are about 80.8 and 179.2 mmolC m^–2 d^–1, respectively. This gain of DIC is higher than the average of CO₂ emissions in the world’s estuaries (45.2 mmolC m^–2 d^–1; Chen et al., 2013) and higher than the NCP calculated for this estuary (section “pCO2 Variability Amplified by Eutrophication” of this manuscript).

The reactions of reduction and oxidation are coupled to proton production and consumption leading to changes in TA and DIC, particularly in estuaries and coastal regions enriched in organic matter (Abril and Frankignoulle, 2001; Hu and Cai, 2011). The stoichiometry of diagenetic reactions alters TA and DIC in specific ways (Rassmann et al., 2020). The highest deviations of TA and DIC concentrations from the expected evaporation line (Figures 5A,B) occurred in the samples with highest salinities from JRE, indicating that there are net gains of TA and DIC from middle/upper estuarine zones. The data from the JRE are between the slopes of sulfate reduction (vector v), denitrification (vector vi), and aerobic respiration (vector vii). Indeed, the biogeochemical processes of aerobic respiration, ammonification, denitrification, and sulfate reduction explained the gains of TA and DIC in other tropical mangrove-dominated estuaries (Dutta et al., 2019; Akhand et al., 2021). A previous study has suggested the influence of sulfate reduction in the JRE (Lacerda et al., 2013), which is particularly pronounced during the dry period due to the longer water residence time and greater influence of the waters trapped inside the mangrove area and when bidirectional flux can be observed at the middle of the estuary (Dias et al., 2016). Sulfate reduction is the main digenetic pathway of organic matter mineralization in mangroves (Borges et al., 2003; Bouillon et al., 2007). Furthermore, high levels of ammonium (NH₄⁺) and low levels of dissolved oxygen were reported at mangrove tidal creeks in the JRE, indicating processes of ammonification and denitrification (Eschrique et al., 2014; Marins et al., 2020). Increasing levels of NH₄⁺ and PO₄^3– in the middle estuarine portion suggest influences of wastes from shrimp aquaculture ponds and domestic sewage (Eschrique et al., 2014). The net heterotrophy explains the higher production of DIC compared to TA.

In classical river-dominated estuaries, there is a clear gradient of salinity in the main estuarine channel, with salinities decreasing from the sea to the higher estuary (freshwater domain) (Frankignoulle et al., 1998; Chen et al., 2013). Tropical rivers present lower TA concentrations than temperate/boreal rivers overall (Cai et al., 2008, 2013; Cotovicz et al., 2020a). Consequently, tropical river-dominated estuaries show lower values of pH, DIC, CO₃^2–, Ω_cal, and Ω_ara in the freshwater domain (Salisbury et al., 2008; Hu and Cai, 2011), and these carbonate chemistry parameters tend to increase seaward or during flooding tides. However, the JRE did not present this classical salinity gradient frequently, and principally during dry conditions; to the contrary, the estuary presents increasing salinity, pH, DIC, TA, CO₃^2–, Ω_cal, and Ω_ara during ebb tides (Figure 3). The high residence time of water in the estuary contributes to the high rates of evaporation, resulting in hypersalinity and creating this atypical pattern of carbonate chemistry at semi-diurnal (tidal-driven) time scale. Evaporation would increase solute concentrations (TA, DIC) and salinity simultaneously without changing the TA/salinity and DIC/salinity ratios (Hu et al., 2015; McCutcheon et al., 2019). Therefore, the residual water (after evaporation) shows higher values of DIC and TA concentrations than the initial seawater entering the estuary. Since evaporation also exerts control on major cations (like Ca²⁺) and consequently changes the solubility of calcium carbonate minerals, it may have the potential to change the CO₃^2–, Ω_cal, and Ω_ara (Millero, 2007; McCutcheon et al., 2019). Indeed, changing only salinity and keeping the other parameters constant, the CO₃^2–, Ω_cal, and Ω_ara will be lower in waters with salinity of 40 than in waters with salinity of 35, due to the changes in DIC speciation (Middelburg et al., 2020). This means that the proportion of dissolved CO₂ to the DIC pool increases, whereas the proportion of CO₃^2– decreases with increasing salinity (maintaining TA and DIC constant).

In addition to this atypical semi-diurnal pattern, the carbonate chemistry in JRE also exhibited a clear light-driven diel cycle due to the influence of biological metabolism. Comparing consecutive periods of lowest tidal height occurring at daytime and night-time, there is a marked difference of CO₃^2– concentrations, and values of pH, Ω_cal, and Ω_ara. The values of these parameters are always lower at night-time because of respiratory processes exceed photosynthetic production of organic matter. There is a variety of organic matter sources available in the estuary, with autochthonous (estuarine phytoplankton) and allochthonous (predominantly mangrove detritus) origins (Mounier et al., 2018; Cavalcante et al., 2021). The respiratory processes lead to a production of CO₂, reducing the CO₃^2– concentrations, and values of pH, Ω_cal, and Ω_ara. On the other hand, the concentrations of CO₃^2– and values of pH, Ω_cal, and Ω_ara were higher at daytime because of primary production, with uptake of CO₂ (Cotovicz et al., 2018; Cyronak et al., 2018). The phytoplankton biomass is highest at daytime and during ebb tide conditions (Chl a concentrations reach up to 2.5 μg L^–1), suggesting that the phytoplankton biomass is concentrated in middle/upper estuarine sections.

The DIC/TA ratio gives information about the buffering capacity of seawater; the waters have lowest buffering capacity when DIC/TA ratio are close to 1 (Egleston et al., 2010). Indeed, the CO₃^2–, pH, Ω_cal, and Ω_ara were lowest when the DIC/TA ratio was at highest levels approaching to 1 (Figure 7). As sulfate reduction and denitrification produce TA and DIC in a very close proportion (1:1), aerobic respiration can be considered the main driver of reduced buffering capacity in the estuary, increasing the DIC/TA. Indeed, the prevalence of respiratory processes over primary production reduces buffering capacity (Anthony et al., 2011; Cotovicz et al., 2018; Cyronak et al., 2018). Enhanced respiratory processes caused by coastal eutrophication has recently been described in estuaries, contributing to coastal acidification (Cai et al., 2011; Wallace et al., 2014). In the JRE, the buffering capacity was lower at night-time than at daytime. This diel pattern was verified in other productive coastal ecosystems, following the diel balance between photosynthesis and respiration. The concentrations of CO₃^2–, and values of pH, Ω_cal, and Ω_ara in the JRE were, in average, slightly lower than those found in adjacent shelf waters (Cotovicz et al., 2020b) suggesting a slight acidification driven by eutrophication in the estuary compared to shelf waters. Calcite minerals were observed in JRE sediments at 30 cm depth by X-ray microanalysis (SEM/EDS), suggesting that the trophic state that affects pH, Ωcal, and Ωara has changed through time (Miguens et al., 2010, 2011).

Overall, estuarine waters are oversaturated in CO₂ with respect to atmospheric equilibrium because estuaries are net heterotrophic ecosystems (Borges and Abril, 2011). The JRE follows this pattern, with pCO₂ values always above the atmospheric pCO₂ (∼ 410 uatm). The variability of pCO₂ shows strong influences from semi-diurnal and diel time scales, as described for other carbonate chemistry parameters. The values of pCO₂ increases during ebb tides, and decreases during flood tides, a common pattern found in diverse estuaries worldwide (Borges and Abril, 2011). However, in the JRE the higher pCO₂ values are coincident with higher salinities, which is not commonly described in estuaries. The hypersalinity in the JRE was developed in more confined waters, coinciding the highest organic matter concentrations and nutrients, which comes from diverse sources, both natural and anthropogenic (Lacerda et al., 2013; Eschrique et al., 2014; Mounier et al., 2018; Marins et al., 2020; Cavalcante et al., 2021). These sources of organic matter contribute to maintain high rates of microbial respiration, sustaining a net heterotrophic metabolism. Indeed, there is a strong positive correlation between pCO₂ values and PO₄^3– concentrations, evidencing that microbial production of CO₂ and remineralization of nutrients by respiratory processes are coupled in the estuary. Other probable explanation is that part of the CO₂ and part of the PO₄^3– in the estuary are coming from same allochthonous sources (shrimp farm; domestic effluents) (Marins et al., 2011, 2020), as well in other eutrophic tropical estuaries (Cotovicz et al., 2015, 2018).

The highest levels of pCO₂ occurred at night-time due to the organic matter respiration and absence of photosynthesis (light-driven diel cycle) (Figures 4A,B). Most of the calculated NCP showed negative values (Table 2) indicating that on average the ecosystem is net heterotrophic, confirming an overall metabolism found in estuaries (Borges and Abril, 2011). Even taking into account the net heterotrophy, the estuarine phytoplankton consumes some CO₂ from the water column considering that pCO₂ values were always lower at daytime compared to night-time considering similar tidal heights. The uptake of CO₂ by estuarine phytoplankton is particularly exacerbated in shallow, stratified, and eutrophic ecosystems (Cotovicz et al., 2015). Reduction in pCO₂ over daylight corresponds to a diel increase in water temperature and solar radiation (Reiman and Xu, 2019). The net heterotrophy in the JRE seems to amplify the diel variability of pCO₂. The diel amplitude of pCO₂ values in the JRE was between 219 and 231 μatm. It is highly above than diel variations found in offshore and oligotrophic sites (Dai et al., 2009) but similar to productive coastal bays, nearshore macrophyte meadows, and hypersaline mangrove-dominated wetlands (Borges et al., 2003; Dai et al., 2009; Delille et al., 2009; Saderne et al., 2013), and lower than those found in coral reef ecosystems and mangrove creeks (Dai et al., 2009; Cotovicz et al., 2020a).

In addition to NCP, the contribution of DIC (and pCO₂) from other sources can be relevant in the main estuarine channel, particularly from mangrove ecosystems and effluent discharges. Carbon export from mangrove soils to the ocean occurs due to soil’s semi-diurnal inundation by tide, a process called “tidal pumping” (Becherer et al., 2016; Santos et al., 2021). Waters flowing from mangroves during ebb tide are enriched in DIC and pCO₂ compared to flood tides (Borges et al., 2003). This outwelling of mangrove-derived carbon can partially sustain the high pCO₂ values found in the main estuarine channel of the JRE, as the estuary has about 13,000 ha covered by mangrove forests (Godoy et al., 2018). Furthermore, there is increasing evidence of direct inputs of effluents from shrimp aquaculture ponds and urban discharges into the estuary (Eschrique et al., 2014; Marins et al., 2020; Lacerda et al., 2021). High pCO₂ values associate with the mineralization of domestic organic matter was described in urbanized estuaries (Frankignoulle et al., 1998; Borges and Abril, 2011). The relationship between E-DIC and AOU shows that JRE present values above the 1:1 line (Figure 8), which represents the quotient between photosynthesis and respiration (Borges and Abril, 2011). The deviation above this line confirms additional sources of sources of DIC (and pCO₂) other than NCP. Moreover, there is thermodynamic influences supporting high pCO₂ values. The CO₂ solubility in hypersaline and warm waters is low (Millero, 2007). Thermodynamic calculations using CO2Calc (Robbins et al., 2011) show that the increase of 1°C in water temperature and 1 unit in water salinity increase the pCO₂ in 25 and 16 μatm, respectively. Hypersaline and warm water holds less dissolved CO₂ increasing the pCO₂, a process described in another evaporative estuary (Yao and Hu, 2017; McCutcheon et al., 2019). However, at the diel time scale, the changes of pCO₂ driven by thermodynamics are minor compared to influences driven by ecosystem metabolism and allochthonous sources.

Estuaries show large range of variability in the values of gas transfer velocity (k₆₆₀) as a function of wind speed compared to lakes, streams, rivers, and open ocean environments (Borges and Abril, 2011). The gas transfer velocity is highly site-specific and one of the major uncertainties in estimating flux rates for CO₂ (Borges and Abril, 2011; Wanninkhof, 2014; Jeffrey et al., 2018). For that reason, when the gas transfer velocity is not directly measured, it is recommended to quantify the air-water CO₂ fluxes using different parameterizations to provide a wider range of air-water CO₂ flux estimates (Jeffrey et al., 2018; Akhand et al., 2021). That is the reason we have applied four commonly used k₆₆₀ parameterizations (McGillis et al., 2001; Raymond and Cole, 2001; Jiang et al., 2008; Wanninkhof, 2014). The FCO₂ in the JRE showed a strong diel influence driven by wind velocity and pCO₂ disequilibria. Indeed, some estuaries show daytime decrease in pCO₂ in coincidence with increase in average wind speed (Reiman and Xu, 2019). In others, the increased evasion during the day is explained by usually higher wind velocities during this period (Cotovicz et al., 2015; Maher et al., 2015). Summarizing, the highest CO₂ effluxes in the JRE are verified when the gradient of pCO₂ at the air-water interface and wind velocities are greatest.

A global estuarine compilation of air-water CO₂ emissions showed that lower estuaries (with salinities higher than 25) are weak sources of CO₂ (23.0 ± 38.3 mmol C m^–2 d^–1) (Chen et al., 2013). With respect to latitude, lower latitude estuaries exhibited moderate emissions (23.5–0°S: 44 ± 29 mmol C m^–2 d^–1; 0–23.5°N: 39 ± 55 mmol C m^–2 d^–1) (Chen et al., 2013). The average of CO₂ emissions in the middle JRE estuarine portion is 51.9 ± 26.7 mmol C m^–2 d^–1 and higher than values verified in other estuaries with similar characteristics. Our values are close to those found in tropical waters surrounding mangrove forests, with air-water CO₂ fluxes of about 50 mmolC m^–2 d^–1 (Borges et al., 2003; Maher et al., 2015). Overall, the hypersaline estuarine waters hold less dissolved CO₂ contributing to enhanced CO₂ efflux, a pattern found in estuaries located at the northwestern Gulf of Mexico (Yao and Hu, 2017; Yao et al., 2020). Our averaged values of CO₂ outgassing exceeded those found in hypersaline waters in the northwestern Gulf of Mexico estuaries, suggesting that eutrophication and allochthonous contributions from mangrove and effluent discharges in the JRE can amplify the CO₂ efflux in these hypersaline waters.

Figure 9 shows a conceptual model with interrelations between semi-arid climate, river damming, eutrophication, and carbonate chemistry (with emphasis on CO₂ emissions) in the JRE. The presence of organic matter retained by the negative estuary enhances the respiratory processes, decreasing the pH values and DIC/TA ratio (buffering capacity), but increasing the pCO₂ and CO₂ emissions. Consequently, CO₂ emissions by the estuarine waters are about 10 times higher compared to nearshore waters (Carvalho et al., 2017; Cotovicz et al., 2020b). Observations and modeling presume the increase of evaporation compared to precipitation in arid/semi-arid regions as the result of climate change (Chou et al., 2009; Huang et al., 2017) and in JRE this process is potentiated by river damming (Dias et al., 2009; Lacerda et al., 2020). Currently, semi-arid regions cover ∼15% of the Earth’s continental surface (Safriel and Adeel, 2005), and the global area of drylands is estimated to expand ∼10% by 2,100 (Feng and Fu, 2013). Indeed, Godoy et al. (2018) reported a 5.6 mm yr^–1 decrease in annual rainfall over the Jaguaribe River basin from the late 1960s that has accelerated in the past 50 years. The consequences of this decrease in rainfall were potentialized by the construction of five large dams. Concomitantly, the eutrophication of estuaries is a widespread environmental problem and particularly enhanced in densely populated regions (Cloern et al., 2014). Despite the recent efforts to implement wastewater treatment plans in developing countries (Tong et al., 2020), the delivery of nitrogen and phosphorus is continuing to grow with urbanization and the associated population increase (Larsen et al., 2016). The water quality of the Jaguaribe River reflects this tropical scenario of environmental degradation, sustaining an increasing process of eutrophication (Eschrique et al., 2014; Marins et al., 2020). The current levels of Ω_ara and pCO₂ are still higher than 2 (supersaturation) and less than 1,000 μatm (hypercapnia), respectively, and similar to other tropical evaporative estuaries (McCutcheon et al., 2019). However, eutrophication and warming tend to increase the values of pCO₂ and CO₂ efflux in the estuary, which could be critical in the coming years considering the actual scenario of climate change and environmental degradation. The enhanced respiratory processes driven by eutrophication tend to decrease the pH, concentrations of CO₃^2–, and vales of Ω_ara and Ω_cal. The critical levels of Ω_cal and Ω_ara (<1) will probably appear early in the JRE compared to adjacent coastal waters due to the ongoing process of eutrophication. Indeed, the coastal eutrophication could increase the susceptibility of coastal waters to ocean acidification (Cai et al., 2011). The diel variability of pH exceeded 0.1 in the JRE, which is the magnitude comparable to the change in the mean ocean pH during the industrial era (Orr et al., 2005). Finally, the amplitude of diel variations is expected to increase with increasing aquatic pCO₂ values in the future, because of simultaneously decreasing buffer capacity driven by global (ocean acidification) (Schulz and Riebesell, 2013), and local (enhanced respiration) processes (Cai et al., 2011).