The fieldwork was designed to capture the along-channel variability of the vertical structure of current velocities, water properties, and mixing over the transition from the lagoon inlet and channel to the adjacent shelf, and into the buoyant plume beyond the jetties. The survey was conducted on 18 July 2022, coinciding with a Sentinel satellite overpass with clear-sky conditions, enabling a spatial match between in-situ data and the surface extent of the plume (Fig. 1c). During the field survey, outflow currents predominated, consistent with the conditions described in Sect. 2, where inflow and outflow at the mouth are primarily modulated by a subtidal wind-driven regime, with astronomical tides playing a minor role.

Considering the limited influence of astronomical tides in the region, the sampling strategy prioritized spatial coverage over tidal phase, and the field campaign was organized into three segments: thalweg drift (Drift), inner shelf outside the jetties (Jetties), and cross-sectional transect (Transect), as referred to hereafter (Fig 1). The Drift segment collected over ∼ 2 h and 20 min (11:21–13:40 LT) represents the main dataset of the experiment, aiming to capture the longitudinal transition from channelized flow to the early buoyant plume evolution. This segment involved a Lagrangian drift along the thalweg of the channel (with the boat moving passively with the outflow), represented by yellow dots in Fig. 1b and c, totalling 108 microstructure profiles collected over approximately 13 km. While small adjustments were required inside the channel to compensate for wind-induced drag and maintain the boat within the thalweg, no corrections to the track were made once outside the inlet. The microstructure profiles were collected with a Rockland Scientific MicroCTD, which will be elaborated upon later in this section.

Subsequently, two complementary sampling segments were performed to characterize the boundary conditions. The Jetties segment (14:30–14:50 LT) was conducted on the inner shelf adjacent to the jetties (orange dots, Fig. 1b and c) to provide baseline values of TKE dissipation outside the buoyant plume, comprising 17 microstructure profiles across three stations. Finally, the Transect segment (15:45–16:30 LT) consisted of a cross-channel section at the upstream portion of the study area (red dots, Fig. 1b and c). This segment included 26 profiles distributed across five stations (∼ 5 profiles per station), designed to capture cross-channel hydrographic variability and quantify the volume outflow discharge during the survey.

During the Drift, Transect and Jetties sampling, current velocities were recorded concurrently using a vessel-mounted Teledyne RDI Workhorse 1200 kHz Acoustic Doppler Current Profiler (ADCP) operating in bottom-tracking mode. The instrument was configured with a ping interval of 0.5 s and a vertical bin size of 0.25 m, with the first measurement bin located approximately 1.5 m below the water surface. Horizontal positioning was determined via a GPS unit directly connected to the ADCP system, with an estimated positional uncertainty of ∼ 5 m. The ADCP operated continuously throughout the field campaign, providing uninterrupted velocity measurements along the survey track. The total water flow across the transect segment was obtained from direct measurements using an ADCP, processed with the WinRiver II software provided by Teledyne RDI.

Salinity and temperature vertical profiles were recorded with a Rinko Profiler CTD (by JFE Advantech LTD) and with a Conductivity-Temperature Sensor (JAC CT) integrated into a Rockland Scientific microstructure profiler (MicroCTD). The MicroCTD recorded shear by two orthogonally mounted shear probes operating at a sampling rate of 512 Hz. The shear profiles were used to estimate turbulence in terms of TKE dissipation rate (ϵ).

The MicroCTD was deployed a total of 152 times in the downward profiling mode at the locations indicated in Fig. 1b and c, and lowered vertically using a weighted collar to ensure a controlled descent speed of approximately 0.8 m s⁻¹ and an angle of attack below 5° (Lueck, 2013; Lueck et al., 2013; Ross et al., 2019; Barros et al., 2025). The MicroCTD data were georeferenced and synchronized with the ADCP data. For each MicroCTD cast, there is a corresponding vertical profile of horizontal velocities, u and v, time-averaged over a one-minute interval.

Additional environmental data, including wind, water level, and Eulerian current measurements, were obtained from the Rio Grande Pilot Station (Fig. 1b). Wind observations were collected by a Davis Vantage Pro 2 meteorological station, recording at 5 min intervals and averaged hourly. Water level and current velocity were measured with a moored Sontek Argonaut-XR 1.5 MHz ADCP, deployed at approximately 17 m depth. This instrument recorded 5 min averaged profiles with a vertical bin size of 1.5 m.

Estimates of freshwater discharge into the Patos Lagoon were also incorporated, based on daily records from the primary tributaries. Discharge data were retrieved from Brazil's National Water Agency (ANA, 2023) Hidroweb database (http://www.snirh.gov.br/hidroweb, last access: 10 September 2023), using gauge stations from the Jacuí (85900000), Taquari (86720000), Caí (87170000), Sinos (87382000), and Camaquã (87905000) rivers. Additionally, discharge from the São Gonçalo Channel, which connects the Mirim Lagoon to the estuarine system, was provided by the Mirim Lagoon Agency (ALM, 2023). The methodology used to aggregate and estimate river discharge followed the approach described by Santa-Rosa and Schettini (2024).

The environmental conditions during the survey favoured the outflow of lagoon waters and the formation of a buoyant plume. To characterize this state, the river discharge of the main tributaries, along with water level, wind vectors in meteorological convention, and current velocities measured in the channel at the Pilot's Station (Fig. 1b) during the days preceding the survey, are shown in Fig. 2. On the day of the fieldwork (18 July 2022), the gauged river discharge was approximately 6900 m³ s⁻¹ (Fig. 2a), which is significantly higher than the estimated average flow of 3707 m³ s⁻¹ for the same month. Importantly, elevated discharge conditions had persisted for several days prior to the survey, indicating sustained freshwater forcing rather than a transient increase in river discharge.

The gauged river discharge represents the freshwater inflow to the lagoon measured upstream from any coastal sea level influence in fully freshwater sections of the main tributaries. However, it does not necessarily correspond to the discharge at the lagoon mouth due to storage, mixing, and exchange processes within the lagoon (Santa-Rosa and Schettini, 2024). The total discharge at the lagoon mouth, measured across the Transect segment, reached 9400 m³ s⁻¹, with an average salinity of 6 g kg⁻¹ (Fig. 3b). Using the observed coastal salinity during the survey (∼ 28 g kg⁻¹) as a reference, the freshwater discharge was estimated at 7650 m³ s⁻¹. The difference between the gauged river discharge and the actual freshwater flow at the lagoon's mouth can be attributed to wind-driven water level variations along the lagoon and the residence time of water parcels within the system.

Water level variability during the survey was minimal, fluctuating by approximately 0.20 m (Fig. 2b). Wind speeds were relatively weak compared to the preceding days (Fig. 2c), blowing from NNW to W and ranging between 1.8 and 5.4 m s⁻¹. The current flowed seaward throughout the entire water column during the fieldwork (Fig. 2d), with velocities ranging from 0.8 m s⁻¹ near the bottom to 2.2 m s⁻¹ near the surface, without reversal. These results from the moored ADCP corroborate the Transect observations (Fig. 3a), showing outflow across the entire water column, with higher velocities of approximately 2 m s⁻¹ concentrated at the surface and along the main axis of the channel.

Together, the elevated river discharge, weak water-level variability, and persistent ebb-directed flow indicate that the system was not responding to tidal oscillations but rather to sustained subtidal forcing acting over several days. Such prolonged forcing conditions are characteristic of microtidal environments and allow the estuary–plume system to evolve toward a quasi-steady state.

Nevertheless, this condition represented a significant variation compared to the previous days in the interval, during which the current had been flowing in the opposite direction due to the passage of a frontal system with up-estuary winds. Following the relaxation of those winds, the system rapidly re-established a persistent seaward-directed circulation that remained stable for several days and throughout the survey period. This high-discharge, outflow-dominated state provides the boundary conditions for the plume development analysed in the following section.

The spatial distributions along the Drift segment are presented in scatter plot diagrams, where vertical profiles feature a 0.5 m resolution and horizontal positioning corresponds to each cast's location. This graphical representation underscores the extensive sampling density achieved during the campaign and allows for the identification of spatial coherence between neighbouring profiles. Such high-resolution visualization is particularly relevant for properties with strong spatial variability, such as TKE dissipation (ϵ) and the Richardson number (Ri), ensuring that the observed structures reflect genuine physical features captured at the sampling sites.

To contextualize the magnitude of the turbulence observed within the plume, the vertical profiles conducted in the Jetties segment serve as a baseline reference for the adjacent inner shelf. Compared to the high-velocity jet observed in the Drift segment, the current velocity profiles at the Jetties segment (Fig. 11a) exhibit distinctly lower magnitudes and greater vertical uniformity, averaging to approximately 0.2 m s⁻¹. Nevertheless, the upper layer at Station 1 (green line), which is the outermost (most seaward) location, exhibited larger variations, reaching over 0.4 m s⁻¹ at the surface. This behaviour is consistent with the density profiles (Fig. 11b), which indicate a more stratified water column at Station 1 compared to moderate stratification at Station 2 (blue line). Conversely, Station 3 (yellow line), which is located closest to shore, exhibits a nearly vertical density profile, indicating a well-mixed water column with no apparent pycnocline.

Under these conditions, the TKE dissipation rates (ϵ) at these stations (Fig. 11c) are significantly lower than those recorded in the main plume axis, ranging between 10⁻⁸ and 10⁻⁶ W kg⁻¹, with the minimum values occurring mid-water column, just below the pycnocline at each station. This is even the case for station 3, which has a nearly uniform density profile, although the values of ϵ are up to a magnitude smaller than the other two stations. Furthermore, Az fluctuates between 10⁻⁵ and 10⁻³ m² s⁻¹ without a specific vertical pattern (Fig. 11d). The reduced velocities and relative uniformity, combined with the markedly lower stratification and dissipation at these reference points, confirm that the intense turbulent activity described in the Drift segment is an intrinsic process of the plume dynamics rather than a general characteristic of the inner shelf. This contrast highlights the unique turbulent signature of the plume, where internal structures are governed by localized mixing processes at the estuary–shelf interface.

The field campaign spanned approximately 5 h, yet the observations can be interpreted as a quasi-synoptic snapshot of the hydrodynamic conditions. In estuarine environments where astronomical tides are dominant, sampling designs must strictly adhere to hourly variations to capture tidal phases (Dyer, 1979; Kjerfve, 1990). However, in the Rio Grande Channel, astronomical tides play a secondary role, accounting for less than 20 % of the current variance (Santa-Rosa and Schettini, 2024). As established by the spectral characterization for July 2022 (Barros et al., 2025), the system is dominated by low-frequency energy (periods > 25 h) associated with synoptic frontal systems.

This meteorological dominance ensured a quasi-steady state during the survey, characterized by minimal water level fluctuations (< 0.20 m; Fig. 2b) and a persistent, uniform ebb-directed flow (Fig. 2d). Such stationarity validates the comparison of data across different sampling moments, allowing the analysis to move beyond temporal variability and focus on the spatial processes and internal plume structure that define the estuary-shelf transition. To further assess the representativeness of the sampled conditions, the field observations were compared with the annual wind–flow regime space derived from 2022 observations (Fig. 12). Together, these analyses indicate that the field campaign sampled a representative quasi-steady outflow state, providing a robust basis for interpreting the spatial structure and mixing dynamics across the estuary–plume transition.

The spatial distribution of the TKE dissipation rate (ϵ) along the drift segment highlights a clear maximum near the mouth and within the buoyant plume's surface layer. These high values align with the plume's “lift-off” zone (MacDonald and Geyer, 2004) and exhibit a decay proportional to the distance from the inlet (O'Donnell et al., 2008; McPherson et al., 2019). This maximum dissipation zone, associated with the plume's evolution over the quiescent coastal water, follows expected theoretical patterns for near-field river plumes (Horner-Devine et al., 2015).

Within the channel, the region of minimum ϵ between approximately km 4 and 7 coincides with an expansion of the cross-sectional area, suggesting the influence of the geometric widening. However, morphology alone is unlikely to explain the observed decrease in ϵ. The structure and magnitude of the along channel current does not change much in this sector, however the variability of ϵ may be related with the secondary circulation. Previous observations in the Rio Grande Channel demonstrated that secondary circulation could generate localized shear and enhance turbulent mixing through the interaction of baroclinic pressure gradients, channel curvature, and bathymetric forcing (Barros et al., 2025). Although the present survey captured a largely flushed and weakly stratified state within the inner channel, coherent lateral circulation patterns remain evident (Fig. 7b). Notably, the region of reduced ϵ between km 4 and 7 also corresponds to weaker cross-channel velocity gradients. This correspondence suggests that the observed decrease in turbulence may reflect not only the local channel expansion but also reduced shear production associated with secondary circulation.

As the flow approaches the final constriction near the mouth (dashed line, Fig. 13), the cross-sectional area decreases (approximately km 8–8.5), followed by an increase in ϵ values. However, the maximum dissipation rates (∼ 10⁻³ W kg⁻¹) occur immediately downstream of the inlet mouth, suggesting that, in addition to channel morphology, plume lift-off and the rapid transition from a laterally confined jet to an expanding buoyant plume contribute substantially to turbulence production. The combined action of these processes appears sufficient to overcome the stabilizing effect of the vertical density stratification, resulting in the intense mixing observed at the estuary–shelf interface.

Once the flow exits the jetties and enters the shelf, a reduction in ϵ is observed (Fig. 13). Although the buoyant plume remains vertically confined to the upper layer (Fig. 4a, c), the loss of lateral confinement permits rapid horizontal spreading. This transition from a topographically constrained jet to an expanding buoyant plume is accompanied by a marked decrease in ϵ, suggesting that the redistribution of momentum over a broader cross-sectional area contributes to the decay of turbulence and the transition away from the high-energy near-field regime.

The observations presented here represent the peak of a discharge event triggered by the relaxation of a meteorological sea-level set-up. As shown in Fig. 2d, the preceding landward wind stress promoted an inward water accumulation within the lagoon; once this forcing ceased, the resulting pressure gradient drove seaward flushing. This mechanism is a recurrent feature of the system's sub-tidal dynamics, where the passage of frontal systems every 5–7 d dictates the alternation between extreme outflow and salt-wedge intrusion (Möller et al., 2001; Fernandes et al., 2005). The comparison with Barros et al. (2025), conducted in the same area only 24 h later, highlights the remarkable temporal variability of the Patos Lagoon inlet. Whereas the previous study documented a salt-wedge configuration with a strong pycnocline that decoupled the surface flow from the bottom topography and suppressed turbulence, the present observations captured a fully flushed, inertia-dominated state in which peak discharge conditions exhibited a much stronger response to channel morphology.

These results suggest that, in microtidal environments, the temporal variability of mixing and turbulence can be governed primarily by synoptic forcing rather than by the predictable tidal cycle. The rapid shift observed here, from an inertia-dominated “flushed” state to a stratified “salt-wedge” intrusion state within only one day, highlights the importance of high-resolution spatial surveys to capture these transient regimes. In this context, the reduced tidal influence represents a key advantage, as it allows the isolation of wind-, discharge-, and morphology-driven processes that are often masked in strongly tidal systems. Therefore, these observations provide a useful framework for understanding the turbulent dynamics of estuarine–shelf systems, not only in microtidal environments but also in more complex settings where multiple forcings interact.

Considering the hydrodynamic transition along the Drift segment the evolution of hydraulic conditions along the estuary–plume transition was investigated. Based on the vertical density distribution and layer-averaged velocities (Fig. 14a–c), the internal Froude number (Fr) was calculated to characterize the flow's criticality (Fig. 14d). Such phenomena are common in estuarine plumes where the flow transitions from a supercritical (Fr > 1) to a subcritical (Fr_i < 1) regime. These transitions are often interpreted through idealized models of internal hydraulic jumps, though their application to natural, high-energy flows requires careful consideration of mixing and layer entrainment (Thorpe et al., 2017). The value of Fr is highly sensitive to the adopted parameters (interface depth, upper layer velocity and g′).

In Fig. 14d, the blue line indicates the internal Froude number calculated using the parameters obtained for every vertical profile. The result is unexpected, to say the least, as it indicates the flow is supercritical over the entire section. Alternatively, the internal Froude number was calculated using a unique reference g′, obtained based on the density difference between the channel's upper layer (∼ 1008 kg m⁻³) and the bottom coastal water (∼ 1022 kg m⁻³), following an approach suggested by Rock Geyer (personal communication, 2026), in similar fashion by MacDonald and Geyer (2005). This resulted in more expected values for the internal Froude number, and in reasonable agreement with MacDonald and Geyer (2005) values found in the uplift region of the Fraser River plume. A supercritical flow means the mean flow velocity is higher than the wave propagation velocity. As a consequence, the water surface becomes very agitated. The condition of a very agitated water surface was observed from the jetties' inlet outward for some distance. In the channel and offshore, away from the jetties' inlet, the sea surface was very calm, as the surveys were taken under breeze conditions.

The results using variable g′ seem hard to interpret physically, which means the present case is out of the scope for its application. On the other hand, the use of a constant reference g′ provides a nice fit to what is expected for the transition from channel flow to the early formation of the buoyant plume (Horner-Devine et al., 2015). From Figs. 7 and 8, it is possible to diagnose intense changes in plume velocity and thickness, and interfacial mixing, when the flow is supercritical. High surface velocities (∼ 2 m s⁻¹) resulted in elevated Fr_i values, as the flow's momentum far outweighed the stabilizing effect of buoyancy, even under sharp stratification. This supercriticality along the main axis is consistent with flow dynamics observed in energetic channel-to-plume transition zones, where high inertia and high-velocity discharge prevent immediate hydraulic adjustment to subcritical conditions (Dorrell et al., 2016).

Such momentum implies that the flow is too rapid for internal disturbances to propagate upstream (Nash and Moum, 2005). This observed vertical structure, which is characterized by a thin, high-velocity surface layer over a saltier ambient water mass, follows the conceptual model of strongly stratified river plumes documented in large-scale systems (Nash et al., 2009). While the supercritical flow blocks the upstream migration of internal waves, the intense velocity gradient it creates is precisely what triggers localized mixing. The strong shear at the plume's base may promote Kelvin-Helmholtz (KH) instabilities, which drive the recorded peaks in ϵ, a process consistent with observations in other highly stratified discharge regions (e.g., Spicer, 2022). The spatial correspondence between elevated ϵ (Fig. 13) and supercritical flow conditions (Fig. 14) indicates that substantial turbulent mixing can occur within the near-field plume before any hydraulic adjustment to subcritical conditions takes place.

The results suggest that hydraulic adjustment occurs ∼ 2 km offshore from the jetties' inlet, when Fr_i reduces to < 1. This is supported by the progressive decrease in velocity, the deepening of the surface layer interface (McPherson et al., 2020). The persistence of supercritical state over a few kilometres seaward of the mouth is a known feature of river plumes where the buoyant layer spreads laterally upon entering the ocean (Hetland, 2010; Horner-Devine et al., 2015). Under these conditions, the lateral straining of the shear layer acts as a persistent mechanism that promotes turbulent mixing. This process sustains elevated ϵ values across the region (Fig. 8) and delays the hydrodynamic transition to a subcritical state until substantial interfacial entrainment and lateral plume expansion are achieved (Geyer et al., 2017). The transition to the mid-field region, associated with subcritical conditions and reduced inertial dominance, is observed beyond km 11.

To evaluate the magnitude of the mixing processes observed, the maximum and minimum values of the primary hydrodynamic parameters were compared across the transition (Table 1) and contextualized with other estuarine systems globally (Table 2). The high-energy nature of the plume is initially evidenced by a peak in stratification (N2 ∼ 10⁻¹ s⁻²) and vertical shear (S2 ∼ 10^−0.5 s⁻²) that concentrate in the outer portion of the channel, marking the region of intense interfacial interaction (Fig. 9a, b). These conditions promote elevated TKE dissipation (ϵ), with peaks on the order of 10⁻³ W kg⁻¹ recorded immediately adjacent to the mouth (Fig. 8).

As synthesized in Table 2, the maximum dissipation rates recorded at the Patos Lagoon mouth and its adjacent plume place this microtidal system among the most energetic estuarine environments documented in the literature. Furthermore, the extreme inertial dominance observed during this survey leads to dissipation rates that match or exceed those recorded in larger meso- and macrotidal systems, such as Gironde (Ross et al., 2019), Merrimack (MacDonald et al., 2007), Changjiang (Wang et al., 2020), Fraser (MacDonald and Geyer, 2004), and Hudson (Peters and Bokhorst, 2000) rivers. This energy level also significantly surpasses other topographically constrained environments in the Southern Hemisphere, including the Patagonian fjords, where tidal-driven turbulence at sills typically reaches values two orders of magnitude lower than those observed in the present study (Pérez-Santos et al., 2018).

Although direct comparisons should be interpreted cautiously due to differences in forcing conditions, sampling strategies, and methodologies among studies, the observed ϵ values place the Patos Lagoon Estuary within the upper range of turbulent conditions reported for energetic estuarine and plume environments. Rather than providing a strict ranking among systems, Table 2 is intended to provide a broader context for interpreting the magnitude of the observed mixing.

These results suggest that the primary distinction between microtidal and meso-macrotidal systems may lie in the forcing mechanisms responsible for generating turbulence rather than in the resulting mixing intensity. In the Patos Lagoon Estuary, turbulence levels comparable to those reported in energetic tidal systems developed under negligible tidal forcing, driven instead by the combined effects of sustained river discharge, persistent outflow, and strong topographic constriction at the inlet. These observations indicate that microtidal plumes can achieve similarly energetic states through fundamentally different forcing pathways.

Consistent with these high dissipation rates, the vertical eddy viscosity (Az) reached maximum values of approximately 10⁻¹ m² s⁻¹ within the plume's interface (Table 1). This magnitude surpasses the maximum value of 10⁻² m² s⁻¹ previously observed in the Patos Lagoon inlet (Barros et al., 2025), reinforcing that under high outflow and specific morphological constraints, the Patos Lagoon inlet functions as a high-energy “nozzle” that maximizes turbulent transport to the inner shelf. Moreover, Az spanned approximately six orders of magnitude across the surveyed domain over less than 14 km, revealing extreme spatial heterogeneity in mixing conditions. This variability highlights the difficulty of representing estuarine mixing using a single representative eddy-viscosity value and suggests that caution is warranted when adopting characteristic mixing parameters to describe estuary–plume systems, whether in classification frameworks or numerical modelling applications.

The buoyancy Reynolds number (Re_b) was applied to map the various stages of turbulence along the transect section (Fig. 10b). Dominant values of Re_b > 200 (blue points) are observed throughout most of the water column, indicating that vigorous, fully developed, and isotropic turbulence prevails in the system. In these regions, buoyancy forces are unable to inhibit the vertical transport of mass and momentum, a state that coincides with the high internal Froude number (Fr_i) values recorded (Fig. 14). This competition between shear and buoyancy was previously identified as a key driver of turbulence within the Patos Lagoon inlet (Barros et al., 2025), where Re_b often serves as a more sensitive indicator of mixing states than traditional stability metrics, as it directly scales the turbulence levels relative to the buoyancy-controlled Ozmidov scale (Stacey et al., 2012).

However, turbulence is locally suppressed within thin layers directly associated with the pycnocline interface (dashed line in Fig. 10b). In these specific pockets, particularly between km 4 and 8, Re_b < 15 occurs (red points), indicating a regime where stratification inhibits turbulent fluctuations and vertical mixing is restricted toward molecular levels (Shih et al., 2005). Surrounding these suppressed zones, a transitional regime (15 < Re_b < 200, yellow points) marks the areas where turbulence is active yet anisotropic. In this regime, despite the sharp stratification, the interfacial shear at the plume front promotes pycnocline erosion and subsequent saltwater entrainment. Such behaviour demonstrates that the presence of stratification does not necessarily imply the complete suppression of turbulent mixing, even under established stratification (Huguenard et al., 2015).

This spatial distribution suggests that while the plume's massive inertia maintains a state of generalized high-energy mixing, the density interface still acts as a localized, albeit intermittent, barrier to vertical transport. In this context, metrics based on the Richardson number (Ri) provide a consistent view of the potential for stratification-driven stability, particularly near the mouth and along the offshore interface, where it indicates a robust buoyancy barrier (Fig. 9c).

Nevertheless, the presence of these stable Ri layers does not necessarily imply the absence of mixing in such high-energy systems. While Ri ≈ 0.25 is commonly interpreted as a classical threshold for the onset of shear instability, recent studies have emphasized that this criterion is strictly applicable only when Richardson number is evaluated over sufficiently small vertical scales (MacDonald and Goodman, 2026). In field observations, Richardson number estimates are often evaluated over observational scales that are larger than those associated with the onset of individual shear instabilities. Under these conditions, turbulence may persist for Ri values substantially greater than 0.25 because unresolved sublayers can locally satisfy instability criteria and sustain turbulent production. Therefore, the classical threshold alone is not sufficient to fully describe mixing in natural environments, as turbulence may be sustained for Ri values approaching unity under energetic forcing conditions (Simpson and Sharples, 2012). This perspective is particularly relevant in the present study, where elevated ϵ and Re_b values frequently coexist with Ri values above the classical instability threshold. As observed between km 9 and 10, even where Ri suggests stability, the buoyancy Reynolds number (Re_b) and elevated ϵ confirm that turbulence remains vigorous enough to dynamically erode the interface. In estuarine environments, this corresponds to an intermediate regime (0.25 < Ri < 1), where mixing is not suppressed but instead modulated by the interplay of shear, straining, and advection (Giddings et al., 2011).

These results suggest that the classical Ri threshold alone may not fully capture the complexity of mixing dynamics in this system. Such limitations of the gradient Richardson number as a universal mixing proxy have been recently documented in other diverse high-energy, wind-dominated systems (Arevalo et al., 2022; Álvarez‐Silva et al., 2026), where vigorous mixing and asymmetric hydrodynamic structures persist even under seemingly stabilizing density gradients. While the transitional criteria Reb=15 and 200 inherently involve statistical uncertainty and scale dependencies, the buoyancy Reynolds number (Re_b) provides a more robust description of turbulent efficiency in this energetically active plume, capturing the full spectrum of mixing states that traditional stability thresholds may overlook.