Accepted Manuscript

Carbonate system biogeochemistry in a subterranean estuary-Waquoit Bay,USA

Qian Liu, Matthew A. Charette, Crystaline F. Breier, Paul B. Henderson, DanielC. McCorkle, William Martin, Minhan Dai

PII: S0016-7037(17)30055-8DOI: http://dx.doi.org/10.1016/j.gca.2017.01.041Reference: GCA 10135

To appear in: Geochimica et Cosmochimica Acta

Received Date: 24 March 2016Revised Date: 20 January 2017Accepted Date: 23 January 2017

Please cite this article as: Liu, Q., Charette, M.A., Breier, C.F., Henderson, P.B., McCorkle, D.C., Martin, W., Dai,M., Carbonate system biogeochemistry in a subterranean estuary-Waquoit Bay, USA, Geochimica et CosmochimicaActa (2017), doi: http://dx.doi.org/10.1016/j.gca.2017.01.041

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Carbonate system biogeochemistry in a subterranean estuary-Waquoit Bay, 1

USA 2

3

4

Qian Liua,b

, Matthew A. Charetteb*, Crystaline F. Breier

b, Paul B. Henderson

b, Daniel C. 5

McCorklec, William Martin

b, Minhan Dai

a 6

7

a State Key Lab of Marine Environmental Science, Xiamen University, Xiamen, Fujian, China 8

9 b Department of Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, 10

Woods Hole, Massachusetts, USA 11

12 c Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, 13

Massachusetts, USA 14

15

16

17

*Corresponding author: mcharette@whoi.edu 18

Office: (508) 289-3205 19

Fax: (508) 457-2193 20

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25

26

Abbreviations: SGD, Submarine groundwater discharge; STE, Subterranean estuary;DIC, dis-27

solved inorganic carbon; TAlk, total alkalinity; WB, Waquoit Bay 28

29

30

31

32

33

34

35

36

37

38

39

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41

42

Abstract 43

Quantifying carbon fluxes associated with submarine groundwater discharge (SGD) remains 44

challenging due to the complex biogeochemistry of the carbonate system in the subterranean estu-45

ary (STE). Here we conducted time series measurements of total alkalinity (TAlk) and dissolved 46

inorganic carbon (DIC) in a well-studied coastal aquifer (Waquoit Bay, Massachusetts, USA). 47

Groundwater samples were collected monthly from May 2009 to June 2010 across the freshwater-48

saltwater mixing zone of the Waquoit Bay (WB) STE. The concentrations of both TAlk and DIC in 49

zero-salinity groundwater were variable, but were lower than those in the bay water (S ~28). DIC 50

underwent slightly non-conservative mixing between low and intermediate salinities while there 51

was an apparent additional DIC source at high salinity (> 20) in all seasons. TAlk concentrations 52

exhibited even stronger variations, with evidence of both production and consumption in high sa-53

linity zones, and consistent TAlk consumption at intermediate salinity in summer and fall (June-54

December, 2009). The increases in DIC and TAlk at high salinity were attributed to aerobic respi-55

ration and denitrification in WB sediments during bay water recharge of the STE. We infer that the 56

loss of TAlk at intermediate salinity reflects H+ production as reduced compounds (e.g. Fe2+) are 57

oxidized within the STE. In terms of impacts on surface water inorganic carbon budgets, the SGD-58

derived DIC flux was mainly controlled by seasonal changes in SGD while a combination of TAlk 59

concentration variability and SGD drove the TAlk flux. SGD-derived DIC, aqueous CO2, and H+ 60

fluxes to the bay were ~40-50% higher in summer vs. in winter, a result of enhanced marine 61

groundwater flux and significant TAlk removal (proton addition) during periods of high seawater 62

intrusion. Furthermore, the SGD-derived DIC flux was consistently greater than TAlk flux regard-63

less of season, indicating that SGD serves to reduce the CO2 buffering capacity of surface water. 64

Our results highlight the importance of seasonality and subsurface biogeochemical processes on 65

the subterranean estuary carbonate system and the resulting impact on SGD-derived TAlk, DIC, 66

aqueous CO2, and H+

fluxes to the coastal ocean. 67

68

69

70

71

1. Introduction 72

Submarine groundwater discharge (SGD) has been increasingly recognized as a significant 73

carbon source to the coastal ocean (Lee and Kim, 2015a; Liu et al., 2012; 2014; Moore, 2010; 74

2011; Reckhardt et al., 2015; Santos et al., 2011; Wang et al., 2015; Yang et al., 2015) thereby 75

playing a role in the marine carbonate system. In some coral reef lagoons, wetlands and mangrove 76

systems, the SGD flux has been shown to dominate estuarine CO2 dynamics (Cyronak et al., 2014; 77

Maher et al., 2013; Santos et al., 2012). SGD has also been reported as a main source of acidity 78

contributing to coastal acidification in both coral reef and wetland ecosystems (de Weys et al., 79

2011; Santos et al., 2011; Wang et al., 2014). In the coastal margin of a volcanic island, however, 80

seawater pH actually increased due to enhanced biological productivity supported by SGD-derived 81

nutrients (Lee and Kim, 2015a). Together, these studies illustrate the complexity and significance 82

of SGD to coastal ocean carbon dynamics. 83

Similar to river-estuary systems, terrestrial groundwater undergoes significant modifications in 84

its chemical composition before discharging into the coastal ocean due to mixing with seawater 85

and both biotic and abiotic reactions with aquifer sediments. This reaction zone has been termed 86

the ‘subterranean estuary’ (STE; Moore, 1999). In contrast with surface estuaries, however, STEs 87

generally have longer water residence times, enhanced sediment-water interactions, wider redox 88

gradients, and higher microbial activities (Moore, 1999). The CO2 system in the STE has been 89

shown to be significantly modified by diagenetic reactions such as carbonate dissolution and or-90

ganic matter remineralization (Cai et al., 2003; Dorsett et al., 2011; Lee and Kim, 2015b; Liu et al., 91

2012). 92

Given the STE’s dynamic hydrology and biogeochemistry, groundwater CO2 system meas-93

urements must be collected on fine spatial and temporal scales to better evaluate the relevant bio-94

geochemical reactions and their impacts on the coastal ocean inorganic carbon cycle. Waquoit Bay, 95

Massachusetts (WB) has a well defined salinity transition zone occurring over a narrow region (ca. 96

20 m wide, 1-2 m thick, Charette et al., 2005; Gonneea et al., 2014) that hosts a suite of biogeo-97

chemical processes (Charette and Sholkovitz, 2006; Spiteri et al., 2008). Seasonal movement of 98

the salinity transition zone in this STE responds to oscillations in the aquifer hydraulic gradient 99

and sea level (Michael et al., 2005; Gonneea et al., 2013a,b). Through year-long monthly car-100

bonate chemistry observations in the WB STE, we sought to quantify CO2 system dynamics in the 101

context of seasonal hydrogeologic changes and to determine factors that drive CO2 biogeochemis-102

try within three main aquifer zones: terrestrial, intermediate salinity, and high salinity groundwater. 103

Finally we estimated the seasonal variability in SGD-derived total alkalinity (TAlk) and dissolved 104

inorganic carbon (DIC) fluxes to the bay. 105

2. Materials and methods 106

2.1. Study site 107

Waquoit Bay is a shallow semi-enclosed estuary (average water depth 1 m) in Cape Cod (MA, 108

USA, Fig.1); it experiences diurnal tides with an average tidal range of 1.1 m (Charette et al., 109

2001). The freshwater contributed to the bay includes rainfall (11%), groundwater discharge 110

(34%), and rivers (55%) (Cambareri and Eichner, 1998). Importantly, the rivers are largely 111

groundwater fed: approximately 90% of the river flow is supplied by groundwater recharge. 112

Therefore, groundwater is the major freshwater source to the bay. 113

The upper unconfined aquifer (10 m) is mainly composed of coarse and fine sand as well as 114

gravel (Charette et al., 2005). Precipitation (114 cm yr-1

) has relatively low seasonal variation 115

(Walter and Whealan, 2005), however, the groundwater table is lower in summer compared with 116

winter and spring due to evapotranspiration (Michael et al., 2005). Inland, the maximum annual 117

water table variability can reach as high as 1 m. Multiple measurement techniques indicate that the 118

most intensive direct SGD source lies along the northern shoreline of the bay (Charette et al., 2001; 119

Mulligan and Charette, 2006). Hereafter, the salinity transition zone will be referred to as the 120

“mixing zone”. The mixing zone oscillates with changes in hydraulic gradient between the aquifer 121

and sea level, which is mainly driven by the latter in this STE (Gonneea et al., 2013a). A higher 122

hydraulic gradient will lead to seaward movement of the mixing zone while a lower hydraulic gra-123

dient will result in landward movement. Our sampling period (May 2009 to June 2010) generally 124

had higher sea level compared with previous years based on monthly mean sea level (MSL) anom-125

aly data from the nearby Woods Hole tidal gauge (Fig. 2). The average MSL in summer is ~4 cm 126

higher in 2009 vs. 2005 to 2007. This pattern was likely due to the combined effects of El Niño 127

and a negative mode of the North Atlantic Oscillation (Gonneea et al., 2013a). 128

2.2. Field sampling 129

We conducted a time series sampling program for groundwater and bay water at the head of 130

the bay (Fig. 1). Monthly groundwater profiles were collected from May 2009 to June 2010 at PZ3, 131

which is a series of nested fixed-depth well points that span the mixing zone of subterranean estu-132

ary. In total there are ten well points ranging in depth from 0.5 to 5.5 m below the beach surface, 133

with intervals of between 0.15 and 1 m. Samples were collected during the same tidal cycle and 134

phase (four hours past high tide, ~3 days before the monthly spring tide) to reduce potential varia-135

bility associated with tidal fluctuations. We simultaneously collected surface waters at four sta-136

tions equally spaced along the head of the bay, ~100 m offshore, with a water depth of ~0.8 m. 137

During November 2009 and June 2010, we collected two additional piezometer profiles at PZ10 138

(13.5 m inland of PZ3) and PZ5 (3 m seaward of PZ3). Piezometer profile PZ10 included nine 139

depths from 0.9 to 5.5 m while PZ5 included sixteen depths ranging from 0.3 to 6.1 m. 140

The groundwater samples were obtained with a stainless steel drive point piezometer system 141

as described in Charette and Allen (2006). Groundwater was extracted using peristaltic pumps at 142

10-50 mL min-1

. Groundwater sampling commenced after purging the wells of ~1 L or 6-13 well 143

volumes. Ancillary water chemistry data including salinity, temperature, dissolved oxygen (DO), 144

pH, and oxidation-reduction potential (ORP, mV) were measured with a YSI 600XLM sonde in a 145

flow-through cell. ORP is a qualitative measure of the tendency of solutes to acquire electrons and 146

thereby be reduced in aquatic environments. Separate 125 mL samples were collected for salinity 147

analysis. Discrete samples for TAlk and DIC were collected in 150 mL borosilicate glass vials as 148

follows: after overflowing 1.5 times, water was removed to create a ~3 mL headspace, then poi-149

soned with 50 µL of a saturated HgCl2 solution. Laboratory tape was wrapped around the cap to 150

maintain an air tight seal. Trace metal samples were filtered into acid cleaned LDPE bottles using 151

a 0.2 µm capsule filter (Pall Acropak). Samples were acidified to pH 1-2 with 20 µL of Optima 152

ultrapure trace metal grade nitric acid (8 M). Nutrient samples were similarly filtered into acid 153

washed 20 mL polyethylene sample bottles, and stored on ice until their return to the laboratory. A 154

second nutrient vial was filled and acidified to ~pH 2 with 8 M sulfurous acid to eliminate scav-155

enging of phosphate by precipitation of dissolved iron prior to analysis. 156

2.3. Sample analysis 157

Salinity was analyzed by a Guideline AutoSal instrument. DIC and TAlk were measured 158

within two weeks of collection. DIC was determined by acidification of 0.5 mL of a water sample 159

and the subsequent quantification of CO2 with a non-dispersive IR detector (Li-Cor 6252). The 160

analytical precision is ±3 µmol kg-1. TAlk was determined using Gran titration (Metrohm 808 161

Titrando with 1 mL burette) with a precision of ±4 µmol kg-1

. Both DIC and TAlk were calibrated 162

by certified reference materials from A.G. Dickson of the Scripps Institution of Oceanography. 163

From the measured T and S, and DIC and TAlk concentrations, we used the CO2SYS.XLS 164

v.14 software (Lewis and Wallace 1998) to derive the concentrations of the DIC components 165

(aqueous CO2, bicarbonate, and carbonate), pCO2, and pH (total scale). Dissociation constants for 166

carbonic acid were those of Millero et al (2006). The CO2 solubility coefficient was taken from 167

Weiss (1974) and the sulfate dissociation constant was derived from Dickson (1990); both are the 168

default values in the CO2SYS.XLS v.14 program. The calculated calcite or aragonite saturation 169

indexes (Ω) were based on Ω= [Ca2+][CO32-]/Ksp, whereby Ω<1 indicates undersaturation for 170

CaCO3, Ω=1 represents the equilibrium CaCO3 phase, and Ω>1 is CaCO3 supersaturation (Ksp is 171

the CaCO3 solubility product; Mucci 1983). For surface water, Ca2+

was derived from an empirical 172

equation involving measured salinity (Ca2+

(mmol kg-1

) =10.28×S/35; Millero 2005). For ground-173

water samples, we estimated Ca2+

concentrations based on Waquoit Bay STE Ca measurements 174

from the summer of 2004 (Ca (mmol kg-1

) =0.28×S+0.11; Appendix Fig. S1). 175

Nutrient analyses (nitrate, phosphate, ammonium) were performed at the Woods Hole Ocean-176

ographic Institution Nutrient Analytical Facility using standard methods on a Lachat QuickChem 177

8000 Flow Injection Analyzer. Mn and Fe were analyzed on a Finnigan Element 2 high-resolution 178

inductively coupled plasma mass spectrometer at the Woods Hole Oceanographic Institution ICP-179

MS Facility. Briefly, samples were diluted 20-fold with 5% Optima nitric acid spiked with Indium 180

(In) as an internal standard to correct for instrument drift and matrix interferences of the solution. 181

Standards were prepared in the same manner as samples (In-spiked) and six-point standard curves 182

were used to calculate sample metal concentrations (Charette and Sholkovitz, 2006). 183

3. Results 184

3.1. Salinity distribution in the subterranean estuary 185

A decreasing hydraulic gradient will cause the landward movement of the mixing zone 186

(Gonneea et al., 2013a), leading to generally higher salinities within our fixed depth piezometers. 187

Conversely, lower salinities are observed with an increasing hydraulic gradient. Although the hy-188

draulic gradient from May 2009 to June 2010 was relatively invariant compared with previous ob-189

servations, we still observed a shallowing of the mixing zone (salinity 5-10) in summer and fall 190

(June to December 2009, Fig. 3) relative to winter and spring (January to May 2010, Fig. 3). Salin-191

ity anomalies in the freshwater portion of the aquifer were observed in June-July 2009 and May-192

June 2010 (Fig. 3) and were likely associated with the mixing between tide and wave induced 193

overtopping of bay water and terrestrial groundwater (Rogers, 2010; Gonneea et al., 2013b). 194

3.2. Redox processes in the subterranean estuary 195

Dissolved oxygen saturation ranged from 3 to 90% (median: 9%), indicative of persistent 196

low-oxygen conditions in the STE (Fig. 4a), while surface bay water had nearly saturated and su-197

persaturated DO ranging 70-154% (Appendix Table S1). In the STE, DO saturation was generally 198

lowest below the mixing zone. The high DO (33-90%) layer at 2.5 m was especially pronounced 199

from June to December, 2009. 200

Dissolved Mn from the monthly time series (Fig. 4b) was elevated below 4 m, ranging from 201

9.1-31.9 µmol kg-1

, but depleted at shallow depths, ranging from 0-3.6 µmol kg-1

. By comparison, 202

the bay water had low dissolved Mn (average = 0.4 µmol kg

-1; Appendix Table S1). Dissolved Fe 203

ranged from 0-542 µmol kg-1

, with the highest values observed in the shallow layer (<2.5 m) in 204

association with the freshwater plume (Fig. 4c). Like Mn, surface water Fe was generally low (av-205

erage = 0.5 µmol kg-1

; Appendix Table S1). During the STE cross-section sampling in November 206

2009 and June 2010, the highest Mn and Fe concentrations were associated with high salinity 207

groundwater at the seaward PZ5 well (Appendix Table S1). This is the most reducing region with-208

in the STE (Charette and Sholkovitz, 2006), and these elevated dissolved Mn and Fe values are the 209

result of Mn and Fe (oxy)hydroxide reduction associated with organic matter oxidation, with the 210

organic matter being supplied by seawater intrusion (Charette et al., 2005; Spiteri et al., 2008). 211

These microbial reaction pathways are favored upon depletion of oxygen and nitrate as the prima-212

ry electron acceptors. 213

Ammonium (Fig. 4e) and phosphate (Fig. 4f) follow a similar pattern to Mn, though there 214

were elevated concentrations of both at some shallow depths (less than 2.5 m) and in some seasons, 215

e.g. NH4+ was as high as 43 µmol L

-1 at 1.5 m in September 2009. In contrast with ammonia, ni-216

trate plus nitrite (N+N) was almost nil or below detection in the deeper groundwater (>3.5 m, Fig. 217

4d). Elevated N+N (163-245 µmol L-1

), centered at a depth of 2.5 m, corresponded with the high 218

DO layer. 219

3.3. Temporal variability of the carbonate system1 in the subterranean estuary 220

TAlk and DIC followed a consistent pattern in the STE (Fig. 5a and b): terrestrial groundwa-221

ter (salinity <1) had lower TAlk (40-873 µmol kg-1, median: 97 µmol kg-1) and DIC (512-1039 222

µmol kg-1, median: 695 µmol kg-1) as compared with marine groundwater (1<salinity ≤25.6), 223

which had a TAlk range of 84-1985 µmol kg1 (median: 1311 µmol kg

-1) and DIC range of 598-224

2046 µmol kg-1

(median: 1609 µmol kg-1

). 225

Groundwater carbonate ion concentrations were low across all salinities, less than 1% of the 226

DIC. In terrestrial groundwater, DIC was dominated by aqueous CO2, accounting for 16-93% 227

(median: 85%) of the DIC while in the marine groundwater DIC was dominated by bicarbonate 228

(11-95%, median: 83%). Both TAlk and DIC in terrestrial groundwater were lower than in bay 229

water (TAlk: 1364-2063 µmol kg-1

; DIC: 1320-1936 µmol kg-1

, appendix Table S1). The 230

TAlk/DIC ratio in all groundwater samples was below 1, which is in contrast with the bay water 231

samples, which had a TAlk/DIC of greater than 1. Seasonal variations of TAlk and DIC at fixed 232

depths were consistent with the observed salinity variations, i.e. TAlk and DIC concentrations in 233

summer and fall were greater than those in winter and spring at fixed depths, suggesting that TAlk 234

and DIC within the STE was influenced by physical mixing between terrestrial and marine 235

groundwater. 236

The entire STE had between one and two orders of magnitude higher pCO2 (~1044-22,670 237

µatm, median: 6243 µatm, Fig. 5c) relative to atmospheric CO2 (~400 ppm) and bay water pCO2 238

(138-1113 µatm, appendix Table S1). Marine groundwater had lower pCO2 than terrestrial 239

groundwater. The calculated pH values at in situ temperature were low in terrestrial groundwater, 240

ranging from 5.23-7.12; pH increased across the mixing zone reaching maximum values of 7.39 at 241

high salinity. The pH was lowest at 2.5 m in association with low TAlk (Fig. 5d). Bay water pH 242

was greater than 7.47 with a maximum of 8.36 (Appendix Table S1). The STE calcite saturation 243

index increased from 0 at the surface to 0.61 at 5.5 m (Fig. 5e and appendix Table S1), while arag-244

onite saturation ranged from 0 to 0.38 (Appendix Table S1), suggesting the system was 245

1 Four typical parameters represent carbonate system, pH, partial pressure CO2 (pCO2), total alkalinity (TAlk, [HCO3

-

]+2[CO32-

]+[B(OH)4-]+[OH

-]+[HPO4

2-]+2[PO4

3-]+[SiO(OH)3

-]+[NH3]+[HS

-]+…-[H

+]F-[HSO4

-]-[HF]-[H3PO4]-…

Dickson, 1981), and dissolved inorganic carbon (DIC, [CO2*]+[HCO3

-]+[CO3

2-]), [CO2

*] represents the concentration

of all unionized carbon dioxide including [CO2] and [H2CO3], i.e. aqueous CO2.* is the sum of [CO2] and [H2CO3].

undersaturated with respect to calcium carbonate across the entire STE during all seasons. 246

4. Discussion 247

4.1. Physical and biogeochemical control on the carbonate chemistry in the subterranean estuary 248

Terrestrial groundwater (salinity<1) displayed a wide range of TAlk and DIC concentrations 249

(Figs. 6a, 7a and 7b); elevated values were generally associated with a terrestrial groundwater 250

plume that also carried elevated concentrations of ammonia, phosphate, and iron in PZ10 (Fig. 6c 251

and d). Isotopic data such as δ56

Fe (Rouxel et al., 2008) and δ15

N-NH4+ (Kroeger and Charette, 252

2008) suggest that this plume originates from aquifer recharge through a freshwater pond to the 253

north of the study location. In addition, organic matter degradation produces DIC, ammonia, and 254

phosphate (Table 2). The observed elevated TAlk and a negative correlation between TAlk and 255

ORP suggest that TAlk and DIC are produced by organic carbon decomposition processes in the 256

terrestrial groundwater (Fig. 6a, b, and appendix Fig. S2). The positive correlation of both TAlk 257

and DIC with ammonia is also consistent with this idea (Fig. 6b, d, and appendix Fig. S2). It is un-258

likely that carbonate minerals are exerting a significant control on STE TAlk and DIC concentra-259

tions as aquifer sediments in this region are primarily composed of coarse grained sands with 260

quartz (Rouxel et al., 2008). In addition, terrestrial groundwater was observed in the shallow layer 261

(≤2.5 m) of the monthly time-series piezometer (PZ3) samples. At 2.5 m, we observed elevated 262

DO and nitrate but depleted ammonia, particularly from June to September 2009 (Fig.4), which 263

may be an indication of nitrification. 264

In order to better understand and quantify the processes that are operating on the inorganic 265

carbon cycle in the STE during mixing of terrestrial and marine groundwater, we must first define 266

the low and high salinity endmembers to be used in our analysis. On the freshwater side, ground-267

water flow is generally laminar such that much of the mixing takes place along the STE salinity 268

gradient. As such, the plume containing high DIC and TAlk that coincides with a persistent terres-269

trial groundwater nutrient and trace metal plume (2-4 m at site PZ10, Fig. 6; see also Kroeger and 270

Charette (2008) and Spiteri et al. (2008)) does not interact or mix with deep STE groundwater. 271

This plume is characterized by relatively low ORP (Fig. 6a), hence, we define our terrestrial 272

groundwater endmember as samples with an ORP>150 mV, which includes those samples that lie 273

just above the mixing zone but below the terrestrial nutrient plume (phosphate and ammonium). 274

On the high salinity side, we use average bay water concentrations since this is the water that is 275

being recharged into the STE during seawater circulation through the aquifer. In this way, any in-276

creases or decreases in DIC or TAlk across the mixing zone are assumed to be exclusively a result 277

of biogeochemical processes that occurred in the subsurface. A summary of the endmembers for 278

terrestrial groundwater and bay water are shown in Table 1. 279

We used the endmember definitions in Table 1 to construct two-endmember mixing diagrams 280

for both the time series PZ3 profiles (Appendix Fig. S3) and the high resolution profiles collected 281

in November 2009 and June 2010 (Fig. 7). We also introduce ∆ΤAlk and ∆DIC, where ∆ is the dif-282

ference between the calculated conservative two-endmember mixing and the measured value (Fig. 283

8); positive values are non-conservative addition while negative values are removal. DIC in the 284

STE displayed addition at high salinity (>20) for all seasons, while it exhibited approximately con-285

servative mixing through intermediate salinities (1-20). TAlk appeared to be added at high salinity 286

in some seasons, but this TAlk excess is smaller and less consistent than the DIC excess. At inter-287

mediate salinity, TAlk displayed a large seasonal variation from nearly conservative mixing (De-288

cember 28, 2009, January, February, April, and June 2010) to non-conservative removal (June to 289

August 2009, November, December 1, 2009, and May 2010) and slightly non-conservative addi-290

tion in March 2010 (Fig. 8 and appendix Fig. S3). These DIC and TAlk patterns are evidence for 291

the existence of temporally-variable physical and biogeochemical controls on their concentration 292

during transport through the STE. To examine in more detail the potential driving forces on DIC 293

and TAlk behavior in the STE, we will separately discuss the transformations occurring at mid sa-294

linity (1-20) and high salinity (>20) with a focus on the more detailed Nov. 2009 and Jun. 2010 295

datasets. 296

4.1.1. Processes driving DIC and TAlk concentrations at mid salinity 297

DIC displayed only slight curvature (consumption) in the mid salinity zone (1-20) from No-298

vember 2009 and June 2010 (Fig. 7b and d). This indicates that DIC was predominately controlled 299

by physical mixing between terrestrial groundwater and bay water in these two seasons. However, 300

while TAlk was approximately conservative in the same zone during June 2010, it was significant-301

ly depleted in November 2009 (Fig. 7a and c). The TAlk removal was located primarily between 302

0.3 to 0.9 m in PZ5 (Fig. 9b). Previous studies have observed oxidation of reduced Mn and Fe in 303

terrestrial groundwater upon transiting the STE, and Fe/Mn oxide coated subsurface sands have 304

been observed in the mixing zone down to the depth of at least 6 m around PZ3 and PZ5, the so-305

called “iron curtain” (Charette and Sholkovitz, 2002; Charette et al., 2005; Gonneea et al., 2008). 306

There is also a Fe and Mn redox front along the high salinity boundary of STE, which accounts for 307

the large vertical extent of the “iron curtain” (Charette and Sholkovitz, 2002; Charette et al., 2005; 308

Charette and Sholkovitz, 2006). 309

Since Fe2+

and Mn2+

oxidation reactions produce H+ but no DIC, these processes would serve 310

to reduce TAlk and pH yet have no effect on DIC. Spiteri et al (2006) reported that the pH gradient 311

between PZ3 and PZ5, driven by seawater intrusion with a relatively high pH, caused a 7-fold in-312

crease in the rate of Fe2+

oxidation relative to fresh groundwater. We observed shallowing of the 313

salinity mixing zone in summer and fall (June to December 2009, Fig. 3) relative to winter and 314

spring (January to May 2010, Fig. 3), suggesting enhanced seawater intrusion occurred in summer 315

and fall concurrent with our observation of TAlk removal during June through early December 316

2009 (Appendix Fig. S3 and Fig. 8). The pH-salinity relationship for November 2009, which 317

showed a reduction of up to 1 pH unit across the mixing zone, also supports our observation (Fig. 318

10). Conversely, pH was largely conservative through the mid salinity mixing zone in June 2010, 319

consistent with the observed conservative distributions of both TAlk and DIC.

320

Of the entire time-series dataset (PZ3 and PZ5; Figs. 8 and 9), the maximum observed TAlk 321

depletion was 780 µmol kg-1 at 0.5 m in PZ3 (salinity 12.0, Fig. 8a, May-2010) and 723 µmol kg-1 322

at 0.9 m in PZ5 (salinity 20.4, Fig. 9b, November-2009), which would require a corresponding 323

proton addition of 780 and 723 µmol kg-1

, with the actual H+ increase determined by the carbonate 324

chemistry and hence buffering capacity. Based on the reaction formulas (Burdige 2006): 325

Fe2+ + 0.25O2 + 2.5H2O → Fe(OH)3+ 2H+ (1) 326

Mn2+ + 0.5O2 + H2O → MnO2+ 2H+ (2) 327

where Fe2+

:H+ and Mn

2+:H

+ equal 1:2, the total amount of reduced Fe and Mn should be ~390 and 328

361 µmol kg-1

assuming that these redox processes are the sole source of acidity driving the TAlk 329

decrease. Manganese concentrations reached a maximum of only ~4 µmol kg-1

in the terrestrial 330

groundwater, while peak dissolved iron concentrations were well in excess of 400 µmol kg-1

. Ele-331

vated ferrous iron in terrestrial groundwater was also observed in prior studies (Charette and 332

Sholkovitz, 2006), therefore, a sufficient level of reduced Fe exists to support the proton produc-333

tion needed to balance the TAlk consumption in this shallow zone. Further, this level of Fe2+

oxi-334

dation would have consumed 97 and 90 µmol kg-1

of DO, well within the amount available within 335

the TAlk depletion zone (up to 225 µmol L-1). 336

While Fe-associated redox reactions can largely explain the observed changes in TAlk, other 337

contributing reactions cannot be entirely ruled out. For example, a slight addition of dissolved Ca 338

in the mixing zone hints that calcium carbonate dissolution may play a small role (Appendix Fig. 339

S1). Also, nitrification could serve to reduce TAlk with no effect on DIC. For the large TAlk deple-340

tion at PZ5 (0.9 m), low nitrate (0.6 µmol L-1

) and elevated ammonia (19.5 µmol L-1

) preclude a 341

significant contribution from nitrification. However, this process may have contributed slightly to 342

the TAlk removal in PZ3 at 2.5m (Fig. 4), though this is well above the zone of significant TAlk 343

removal (3.5 m, Fig. 8). The “iron curtain” is within this depth range (Gonneea et al., 2008), sug-344

gesting that Fe2+

oxidation is likely the major controlling factor even at 3.5 m. We note, however, 345

that these estimates of STE TAlk removal depend on the choice in TAlk endmembers. If slightly 346

shallower terrestrial groundwater with higher TAlk were used, then these calculated values would 347

represent lower limit estimates. 348

4.1.2. Processes driving DIC and TAlk concentrations at high salinity 349

In June 2010, non-conservative addition of DIC and TAlk was observed in the high salinity 350

groundwater that was characterized by elevated ammonia, phosphate, and dissolved Mn (Fig. 11). 351

Because changes in DIC relative to TAlk would be differentially modified by various organic car-352

bon decomposition reactions (Table 2), changes in their ratio can be combined with changes in the 353

products of organic matter decomposition (e.g. NH4, Mn, and Fe) to quantify the relative im-354

portance of the geochemical processes occurring within this zone of the STE (Froelich et al., 1979; 355

Bender and Heggie, 1984; Cai et al., 2003). During our study, we did not measure SO42-, H2S or 356

CH4 though previous studies indicated that SO42- is conservative and H2S has never been detected 357

(Gonneea and Charette, 2014). We therefore only considered aerobic respiration, denitrification, 358

manganese reduction, and iron reduction as potential driving mechanisms for the observed TAlk 359

and DIC distributions in the high salinity zone (Table 2). Further, Ca2+

data measured in 2004 for 360

these same piezometer locations (Appendix Fig. S1) indicate conservative mixing through the high 361

salinity region. Hence, we also assume that no calcium carbonate precipitation/dissolution oc-362

curred in this zone. The details of this approach are presented in Appendix. 363

In June 2010, both excess TAlk and DIC were observed in the high salinity region (Figs. 12b, 364

c, and 13); these excesses at 2.4-6.1 m of PZ5 and deep in PZ3 (4.0-5.5 m, Fig. 12) are within the 365

range of reactions from aerobic respiration to iron reduction, but are closest to aerobic respiration 366

and denitrification (Fig. 13). Our calculations suggest that aerobic respiration was responsible for 367

24-100% (average 55±19%; Fig. 12d) of the organic carbon remineralization, while denitrification 368

was in the range of 0-79% (average 42±18%; Fig. 12e). Both Mn and Fe reduction were much 369

smaller (Mn reduction; 0-11%, average 3±3%; Fe reduction; 0-1%, average 0.1±0%; Fig. 12f). If 370

we assume that bay water in the STE recharge zone had an initial O2 concentration of 240-448 371

µmol L-1

(the range in our surface water measurements) and the STE O2 in the high salinity zone 372

was in the range of 52-136 µmol L-1, then the oxygen utilization should be 188-312 µmol L-1, 373

which is within the range of our modeled DO consumption (40-457 µmol L-1). Therefore, aerobic 374

respiration and denitrification are the dominant reactions behind organic matter oxidation at high 375

salinity in June. In addition, aside from denitrification, elevated NH4+ might result from ammonifi-376

cation within the organic carbon rich sediments of the outer bay (Gonneea and Charette, 2014) 377

where saline groundwater recharge originates (Michael et al., 2005). However, previous work sug-378

gests that denitrification does occur within this high salinity zone (Korner and Zumft, 1989). 379

The ∆TAlk and ∆DIC distributions for November 2009 at high salinity are more complex. 380

While TAlk and DIC additions were observed at the base of PZ3 (4.3 to 5.5 m) and PZ5 (4.9 to 6.1 381

m), above these depths there was DIC addition in the presence of TAlk depletion (Fig. 9b and c). 382

Restricting our model calculations to these greater depths, we obtained an aerobic respiration frac-383

tion of 27-93% (average 67±26%), a denitrification fraction of 1 to 70% (average 31±26%), Mn 384

reduction from 0 to 4% (average 1±2%), and Fe reduction of 0 to 5% (average 1±2%). Similar to 385

June 2010, the modeled DO consumption (181-441 µmol L-1) is within the range of the observed 386

DO loss (48-430 µmol L-1). In general, the high salinity zone in PZ5 is predominately controlled 387

by aerobic oxidation both in Nov. 2009 and Jun. 2010 (Figs. 9 and 12). 388

Since our approach is based on net changes in water chemistry, it should be noted that anaer-389

obic decomposition processes and subsequent re-oxidation of the reduced compounds could be 390

interpreted in our model as being due to aerobic decomposition. Potential evidence for this is the 391

negative ∆TAlk/∆DIC ratios (-3.1 to -29.3) observed for November 2009 at PZ3 from 3.4 to 3.6 m 392

and PZ5 from 1.5 to 4.2 m (Fig. 9a). These ratios are beyond those predicted for various OM de-393

composition processes (-0.2 to 8, for aerobic respiration to iron reduction, Fig. 13). These could 394

include oxidation of Fe2+ and Mn2+ (Fig. 13), which are believed to be responsible for the large 395

TAlk sink in the mid salinity zone. Other possibilities include Mn2+ mediated-oxidation of ammo-396

nia as suggested by Kroeger and Charette (2008) for this same STE. 397

4.2. Seasonal variability of SGD-derived TAlk and DIC fluxes to surface water 398

To assess seasonal variability of SGD-associated TAlk and DIC fluxes into the bay, we com-399

bined a previously published hydrological flow model for the Waquoit Bay STE (Gonneea et al., 400

2013a, b) with our time-series TAlk and DIC concentrations. The model simulated SGD flux was 401

derived from a two-dimensional cross sectional groundwater flow and salt transport model 402

(SEAWAT) (Mulligan et al., 2011). The model was bounded by an upland margin (CCC1 in Fig.1) 403

and extended 125 m seaward of CCC1. Since the groundwater head at CCC1 was not measured 404

during our sampling period, we used the average SGD water flux for summer and winter (2005-405

2007) from Gonneea and Charette (2014). As with Gonneea et al. (2013a, b), we divided SGD into 406

five salinity groups: 0-5, 5-10, 10-15, 15-20, and >20. Groundwater-derived TAlk and DIC fluxes 407

were then calculated by multiplying the average groundwater TAlk and DIC concentrations by the 408

water flux for the same salinity grouping (Table 3). The same approach was utilized to estimate 409

fluxes of aqueous CO2 and H+ discharged into the bay from groundwater (Table 3). Since the 410

simulated groundwater flux was from 2005 to 2007, while our carbonate system parameters were 411

sampled during 2009-2010, we limit our discussion to assessing the dominant controlling factors 412

for seasonal variability in SGD-derived carbonate system fluxes from a typical unconfined coastal 413

aquifer. 414

Gonneea et al. (2013a) found that the fresh groundwater flux (salinity from 0 to 5) was rela-415

tively constant between summer and winter (2.2±0.6 vs. 3.2±0.4 m3 m-1 day-1; Table 3). They also 416

showed that landward movement of the mixing zone in summer was due to the relatively low hy-417

draulic gradient compared with winter (Fig. 3). As such, the saline groundwater flux was ca. 5 418

times greater in summer than winter (3.9±1.3 vs. 0.8±0.2 m3 m

-1 day

-1). This contrast is largely 419

responsible for the ~50% higher estimated SGD-derived net DIC and aqueous CO2 fluxes to the 420

bay in summer (3.8±2.5 vs. 2.7±0.4 mol m-1

day-1

for DIC flux, 2.5±0.6 vs. 1.7±0.2 mol m-1

day-1

421

for aqueous CO2 flux, Table 3). The estimated SGD-associated H+ flux followed the same trend 422

(9.4±2.5 vs. 6.7±1.1 mmol m-1 day-1, Table 3). Hence, SGD could be an important contributor to 423

coastal ocean acidification in addition to other factors that are driving ocean pH lower. 424

Despite the 5x higher marine SGD flux in summer, the net bay water SGD-TAlk flux was 425

comparable between the two seasons (0.9±2.3 vs. 0.9±0.3 mol m-1 day-1, Table 3), a result of low 426

summertime STE TAlk concentrations. Here the multi-salinity zone model we employed was cru-427

cial: because of the strong mixing zone gradient in TAlk, an approach that divided SGD water 428

fluxes into just two components (fresh and marine groundwater), would have resulted in a signifi-429

cant overestimate in the SGD-TAlk flux (nearly 5x for the summer period). 430

Regardless of season, the SGD-derived DIC flux is always greater than TAlk flux, indicating 431

that SGD would serve to increase the surface water DIC/TAlk ratio. In general, our bay water DIC 432

concentrations were lower than TAlk (Appendix Table S1). Since the CO2 buffering capacity of 433

seawater is weakened as the DIC/TAlk ratio approaches 1 (Egleston et al., 2010), our results for 434

this system indicate that SGD serves to reduce the CO2 buffering capacity of surface water, which 435

is consistent with other recent studies suggesting that this may be a common effect of SGD on the 436

global carbon cycle (Liu et al., 2014; Sadat-Noori et al., 2016; Wang et al., 2015). 437

4.3. Potential SGD effects on the surface water CO2 system on annual time scales 438

To examine the net annual effect of SGD on the coastal ocean carbon cycle, we compared the 439

yearly inputs and outputs of DIC and TAlk to the Waquoit Bay STE. The input terms include ter-440

restrial groundwater and the bay water that enters the STE as marine groundwater recharge, while 441

the output is the groundwater discharged into the bay. Using the model average water flux for 442

2005, the DIC output was greater than the combined inputs which translated into a net STE DIC 443

production rate of 420 mol m-1

yr-1

, which is ~19 % of the DIC flux to the bay via groundwater 444

discharge (Table 4). This is consistent with the DIC excess that was observed in groundwater dur-445

ing almost all seasons. In contrast, the TAlk input was roughly equivalent to the output (5% differ-446

ence, Table 4), suggesting that TAlk removal in the mid salinity zone was offset by TAlk produc-447

tion at high salinity on an annual basis. We would reach the same conclusions if we used the mod-448

eled mean water fluxes for 2006. 449

In Waquoit Bay, mixing zone dynamics appear to be largely controlled by seasonal variations 450

in mean sea level. Gonneea et al. (2013a) noted that the period of our study was characterized by a 451

positive sea level anomaly driven by regional climate oscillations. Their model results indicate that 452

such an anomaly would result in enhanced STE mixing and marine groundwater re-453

charge/discharge in summer. Such an enhancement would serve to increase the DIC flux via en-454

hanced groundwater flux and organic matter oxidation. Further, since nutrients are regenerated in 455

this process, they might serve to reduce the net effect of the SGD-delivered DIC flux through bio-456

logical uptake in surface waters. However, these two processes may not perfectly offset each other 457

as some nitrogen is removed in the high salinity zone of the STE (Kroeger and Charette, 2008), 458

and phosphate may be retained in the subsurface through sorption to Fe oxides (Charette and 459

Sholkovitz, 2002, 2006; Kroeger and Charette, 2008). Therefore, seasonal oscillations in MSL 460

may still result in a net DIC flux to surface water by SGD. 461

We estimated that TAlk removal was offset by TAlk excess in the STE on an annual basis in a 462

“normal” year (2005-2007). However, our data suggest that net STE TAlk removal could occur in 463

association with positive sea level anomaly as was observed from 2009-2010. In this scenario, the 464

SGD-derived DIC flux would be even greater than the SGD-TAlk flux, resulting in (1) enhanced 465

surface water aqueous CO2 inputs thereby enhancing air-sea exchange and reducing the seawater 466

buffering capacity and (2) enhanced proton fluxes that would exacerbate coastal ocean acidifica-467

tion. Our results highlight the need for in-depth studies on estuarine acidification in systems like 468

Waquoit Bay, where organisms including many commercially valuable shellfish may already be 469

experiencing carbonate chemistry conditions at least as severe as those predicted for the open 470

ocean by the end of the 21th century (Doney et al., 2009). 471

472

5. Conclusions 473

This study demonstrates the complexity of the inorganic carbon cycle in subterranean estuar-474

ies due to time-varying hydrologic and biogeochemical processes. Our main conclusions are as 475

follows: 476

(1) The major component of DIC in the terrestrial groundwater was aqueous CO2 (83% of 477

DIC); the entire STE had one to two orders of magnitude higher pCO2 levels relative to 478

atmospheric CO2. The STE had lower pH relative to surface water and it was 479

undersaturated with respect to calcite and aragonite during all seasons. 480

(2) DIC concentrations in brackish groundwater (salinity 0-20) were lower than in surface wa-481

ter. DIC was slightly consumed at low and mid salinities while there was an apparent DIC 482

source at high salinity (> 20) in all seasons. In November 2009 (fall) and June 2010 483

(summer), TAlk and DIC addition occurred in high salinity groundwater. Stoichiometric 484

calculations suggest that these additions were mainly driven by aerobic oxidation and 485

denitrification. 486

(3) Within brackish groundwater during summer and fall, TAlk generally decreased due to 487

proton production from Fe2+

oxidation reactions, in part due to nitrification; TAlk largely 488

displayed conservative mixing in other seasons. Hence, oscillations in the location of the 489

mixing zone, which is also a redox/pH boundary, are driving Fe-based oxidation-reduction 490

cycles that in turn are driving the STE TAlk distribution. 491

(4) The net SGD-derived DIC flux to the bay exceeded the TAlk flux by a factor of ~3-5; the 492

former was ~50% higher in the summer vs. winter while the latter was constant year-round. 493

The lower TAlk flux was due to lower TAlk vs. DIC concentrations in the STE. The mag-494

nitude of the groundwater flux controlled SGD-DIC flux, while, depending on the season, 495

groundwater flux and TAlk concentration were co-drivers of the SGD-TAlk flux. 496

(5) On an annual cycle, the subterranean estuary was a net producer of DIC while the TAlk 497

cycle was in balance, i.e. consumption and production of TAlk was offset in the mid and 498

high salinity zones, respectively. 499

Future studies aimed at quantifying SGD-carbonate chemistry fluxes should take into ac-500

count potential non-conservative processes in the STE as well as seasonal variability. In particular, 501

our finding of significant TAlk removal (H+ addition) during high seawater intrusion periods im-502

plies that SGD’s role in the carbon cycle for receiving water bodies may increase during seasonal 503

periods of high sea level. In Waquoit Bay and the greater northeastern U.S., Gonneea et al. (2013) 504

postulated that sea level is a major controlling factor for SGD; recent studies support this idea for 505

other areas of the world’s coastline (Lee et al., 2013; Wood and Harrington, 2015). Enhanced sea-506

sonal or annual variability in sea level due to climate oscillations like the El Nino-Southern Oscil-507

lation (ENSO) and North Atlantic Oscillation (NAO) could disrupt the year to year balance in net 508

inorganic C fluxes from the STE, which may lead to further enhancements in the DIC and aqueous 509

CO2 discharge from groundwater and the associated impacts on seawater’s CO2 buffering capacity. 510

511

512

513

514

515

Acknowledgements 516

The research was financially supported by the National Science Foundation Chemical Oceanogra-517

phy program (OCE- 0425061 and OCE-0751525 to M.C.) and a China Scholarship Council to Q.L. 518

We thank Dr. Weijun Cai and Dr. Christophe Rabouille for their thoughtful comments on early ver-519

sions of the manuscript. We are also grateful to Dr. Meagan Eagle Gonneea for helpful suggestions 520

related to the groundwater flux model and to Dr. Caroline Slomp for the calcium measurements. 521

We would also like to acknowledge three anonymous reviewers for their critical and constructive 522

comments, which resulted in substantial improvements to the manuscript. 523

524

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regime of a coastal groundwater-Fed Wetland. Groundwater 53, 90-98. 705

706

Yang L., Chen C.-T. A., Hong H., Chang Y.-C., Lui H.-K. (2015) Mixing behavior and bioavaila-707

bility of dissolved organic matter in two contrasting subterranean estuaries as revealed by fluores-708

cence spectroscopy and parallel factor analysis. Estuar. Coast. Shelf Sci. 166, 161-169. 709

710

711

712

713

714

715

716

717

718

719

720

721

722

723

724

725

Figure Captions 726

Fig. 1. Map of Waquoit Bay off Cap Cod in Massachusetts with a schematic figure showing the 727

sampling stations in the surface bay water and subterranean estuary. The STZ is the salinity transi-728

tion zone where mixing between terrestrial groundwater and marine groundwater occurs. Monthly 729

sampling was conducted from May 2009 to June 2010 both in the surface water and groundwater 730

at PZ3, which is a series of nested, fixed-depth wells installed at 10 depths ranging from 0.5-5.5 m 731

centered on the STZ of subterranean estuary. In addition, we sampled two more piezometer wells 732

at PZ10 (inland, 13.5 m away from PZ3 at depths ranging from 0.9 to 5.5 m) and PZ5 (seaward, 3 733

m away from PZ3, at depths ranging from 0.3 to 6.1 m) in November 2009 and June 2010. 734

735

Fig. 2. Monthly mean sea level (MSL) anomaly from Woods Hole tidal gauge (NOAA, 736

#8447930). The shaded region indicates our sampling period from May 2009 to June 2010. 737

738

Fig. 3. Time-series measurements of pore water salinity at PZ3 from May 2009 to June 2010. The 739

x-axis indicates time and the y-axis indicates depth below the beach surface (m). 740

741

Fig. 4. Time-series contours of pore water chemistry for (a) dissolved oxygen (DO), (b) total dis-742

solved Mn, (c) total dissolved Fe, (d) nitrate and nitrite (NO3-+NO2

-), (e) ammonium (NH4

+), and 743

(f) phosphate (PO43-

) at PZ3 from May 2009 to June 2010. The dashed line represents salinity con-744

tour line of 1.The x-axis indicates time and the y-axis indicates depth below the surface. 745

746

Fig. 5. Time-series contours of pore water chemistry for (a) TAlk, (b) DIC, (c) pCO2 (in situ tem-747

perature), (d) pH (total scale, in situ temperature ), and (e) calcite saturation index at PZ3 from 748

May 2009 to June 2010. The dashed line represents salinity contour line of 1. The x-axis indicates 749

time and the y-axis indicates depth below the surface. 750

751

Fig. 6. Depth profiles showing (a) salinity, oxidation reduction potential (ORP), pH, (b) TAlk, 752

DIC, (c) dissolved Mn, Fe, (d) nitrate, ammonium, and phosphate for the inland well (PZ10) sam-753

pled in November 2009. The general overlap in the distribution of elevated TAlk, DIC, Fe, Mn, 754

and ammonia at 2-4 m with low ORP and high pH is indicated by the gray shading. 755

756

Fig. 7. Plots showing TAlk versus salinity (a, c), and DIC versus salinity (b, d) in the Waquoit Bay 757

subterranean estuary during November 2009 and June 2010 in piezometer wells PZ10, PZ3, and 758

PZ5. TAlk and DIC for terrestrial groundwater in Nov are also shown as inset graphs with an ex-759

panded salinity scale (0-0.5). The solid lines represent the conservative mixing lines for TAlk and 760

DIC in deep STE, which was used to quantify the variations of the carbonate system in high salini-761

ty zone. 762

763

Fig. 8. Time-series contours of pore water chemistry for (a) ∆TAlk and (b) ∆DIC (∆ means the 764

difference between the calculated conservative mixing and the measured value) at PZ3 from May 765

2009 to June 2010. Dashed lines are the salinity contour lines of 1, 5 and 20. The x-axis indicates 766

time and the y-axis indicates depth below the surface. 767

768

Fig. 9. Distributions of (a) salinity, (b) ∆TAlk, (c) ∆DIC, fraction of (d) aerobic oxidation, (e) 769

denitrification, and (f) Mn & Fe reduction in the piezometer wells sampled in November 2009. 770

Triangle shows the site of terrestrial groundwater endmember. In panel (a), triangles show the sites 771

of terrestrial groundwater endmember, rectangles denote areas of high salinity (>20) but with neg-772

ative ∆TAlk/∆DIC ratios, which are beyond those predicted for various OM degradation reactions. 773

The box in panel (b) represents area with maximum negative ∆TAlk. 774

775

Fig. 10. The pH versus salinity within subterranean estuary samples collected in November 2009 776

and June 2010 from piezometer wells PZ10, PZ3, and PZ5. The conservative mixing lines for pH 777

were calculated as follows: conservative mixing concentrations for TAlk and DIC were estimated 778

as a function of salinity using the terrestrial groundwater and bay water endmembers from Table 1. 779

These data were entered in CO2SYS, which was used to obtain the corresponding mixing-derived 780

pH. The observed pH in Nov. 2009 from intermediate and high salinity are much lower than the 781

predicted values based on conservative mixing. In Jun. 2010, the observed pH is nearly consistent 782

with the theoretical mixing line at mid salinity (1-20), but is slightly smaller than the calculated 783

conservative pH value at high salinity (>20). 784

785

Fig. 11. Relationship between (a) ammonia, (b) phosphate, (c) dissolved Mn, (d) dissolved Fe and 786

salinity for PZ10, PZ3, and PZ5 wells sampled in November 2009 and June 2010. 787

788

Fig. 12. Distributions of (a) salinity, (b) ∆TAlk, (c) ∆DIC, fraction of (d) aerobic oxidation, (e) 789

denitrification, and (f) Mn & Fe reduction in the piezometer wells sampled in June 2010. The tri-790

angles mark the locations used in generating the terrestrial groundwater endmember. 791

792

Fig. 13. The ∆TAlk versus ∆DIC within subterranean estuary in mid salinity (1-20) and high salin-793

ity (>20) zone samples collected in November 2009 and June 2010 from piezometer wells PZ3 and 794

PZ5. The arrows indicate how ∆TAlk and ∆DIC will change in response to different biogeochemi-795

cal processes including aerobic respiration, denitrification, manganese reduction, iron reduction, 796

Fe2+

or Mn2+

oxidation, and CaCO3 production/dissolution. Numbers inside parentheses represent 797

stoichiometric ratios of changes in TAlk and DIC per mole organic matter degradation. The de-798

tailed biogeochemical reactions are presented in Table 2. 799

800

801

Table 1. Salinity, TAlk, DIC, and pH in the endmembers of terrestrial groundwater and bay water 802

collected in November 2009 and June 2010. 803

Sampling date Endmember Salinity TAlk DIC pH

(µmol kg-1

)

Nov-09 *Terrestrial

groundwater 0.1 67 747 5.4

Baywater 28.8 1888 1754 8.0

Jun-10 #Terrestrial

groundwater 0.4 75 666 5.4

Baywater 27.6 1849 1747 7.7

*Average from stations at 4.9 and 5.5 m in piezometer well PZ10; #Average from stations at 4.3, 4.9, and 804

5.5 m in piezometer well PZ10, 2.4 m in PZ3. The locations of these stations were marked in Figs. 9 and 13. 805

The pH is calculated from the measured TAlk and DIC with in situ temperature. 806

807

808

809

810

811

812

813

814

815

816

817

818

819

820

821

822

823

824

825

826

827

828

829

830

831

832

833

Table 2. Organic matter degradation reactions (Froelichetal., 1979; Bender and Heggie, 1984; Cai 834

et al., 2003) 835

836

Process Geochemical reactions

Aerobic respiration OM+138O2 → 106CO2+16HNO3+H3PO4+122H2O (1)

Denitrification OM+0.8×106NO3

-+0.8×106H

+ →

106CO2+0.4×106N2+16NH3+H3PO4+1.4×106H2O

(2)

Manganese reduction OM+2×106MnO2+4×106H

+ →

106CO2+2×106Mn2+

+16NH3+H3PO4+318H2O

(3)

Iron reduction OM+4×106FeOOH+8×106H

+ →

106CO2+4×106Fe2++16NH3+H3PO4+583H2O

(4)

The stoichiometric ratios of changes in TAlk and DIC per mole organic matter degradation

(1)dδTA/δDIC = (-16-12)/(106)=-0.160

(2)dδTA/δDIC = (0.8×106+16-12)/(106)=0.932

(3)dδTA/δDIC = (4×106+16-12)/(106)=4.132

(4)dδTA/δDIC = (8×106+16-12)/(106)=8.132

OM=(CH2O)106(NH3)16(H3PO4) 837

δ represents the change in TAlk and DIC associated with individual organic matter degradation as it pro-838

ceeds from left to right. 839

840

841

842

843

844

845

Table 3. A comparison of carbonate chemistry and SGD-related TAlk and DIC fluxes in summer and winter in the Waquoit Bay sub-846

terranean estuary 847

Season Salinity Zone SGR Net flux

0-5 5-10 10-15 15-20 >20

Summer (July (July1 and July 22) and August, 2009)

Water Flux＊

(m3 m

-1 day

-1)

2.2±0.6 0.7±0.7 0.3±0.3 0.5±0.4 2.4±1.0 -2.4±0.8

TAlk (mmol m

-3)

170±141 230±55 411±72 732±272 1648±62 -1748±135

DIC (mmol m

-3)

724±63 1019±68 1187±56 1255±202 1826±26 -1670±55

Aqueous CO2 (mmol m

-3)

551±77 552 475 524±143 205±34 -17±3

H+

(mmol m-3

) 3.47±0.52 1.03 0.59 0.77±0.36 0.15±0.03 -0.01±0.004

TAlk flux

(mol m-1

day-1

) 0.4±0.3 0.2±0.1 0.1±0.1 0.4±0.3 4.0±1.6 -4.2±1.5 0.9±2.3

DIC flux

(mol m-1

day-1

) 1.6±0.4 0.8±0.7 0.3±0.4 0.7±0.5 4.4±1.8 -4.0±1.4 3.8±2.5

Aqueous CO2 flux (mol m

-1 day

-1)

1.2±0.4 0.4±0.4 0.1±0.1 0.3±0.2 0.5±0.2 -0.04±0.02 2.5±0.6

H+ flux

(mmol m-1

day-1

) 7.8±2.3 0.8±0.7 0.2±0.2 0.4±0.4 0.4±0.2 -0.03±0.01 9.4±2.5

Winter (December(December 1 and December 28), 2009; January and February, 2010)

Water Flux＊

(m3 m

-1 day

-1)

3.2±0.4 0.4±0.1 0.2±0.1 0.1±0.1 0.1±0.0 -0.4±0.06

TAlk (mmol m

-3)

284±43 512±141 701±197 1116±183 1719±74 -1843±256

DIC (mmol m

-3)

750±23 875±77 1062±167 1424±100 1863±43 -1753±214

Aqueous CO2

(mmol m-3

) 465±38 344±156 335±141 313±92 167±63 -22±6

H+

(mmol m-3

) 2.08±0.25 0.4±0.19 0.31±0.1 0.26±0.14 0.1±0.05 -0.01±0.003

TAlk flux (mol m

-1 day

-1)

0.9±0.2 0.2±0.1 0.2±0.1 0.2±0.1 0.1±0.02 -0.7±0.2 0.9±0.3

DIC flux (mol m

-1 day

-1)

2.4±0.3 0.4±0.1 0.2±0.1 0.2±0.1 0.1±0.02 -0.7±0.1 2.7±0.4

Aqueous CO2 flux

(mol m-1

day-1

) 1.5±0.2 0.1±0.1 0.08±0.04 0.04±0.02 0.01±0.005 (-9±3)×10

-3 1.7±0.2

H+ flux

(mmol m-1

day-1

) 6.6±1.1 0.2±0.1 0.07±0.03 0.04±0.02 0.01±0.004 (-4±1)×10

-3 6.7±1.1

Error bar is standard deviation, representing monthly variation in one season within one year or three years; all fluxes are normalized to per meter 848 of shoreline. 849 * Water flux was modeled from 2005 to 2007 (Gonneea and Charette, 2014). 850 Negative values denote water and material transport from bay water to the subterranean estuary; 851 SGR means submarine groundwater recharge, here we used the average surface samples as the endmember of SGR. 852 853

854

30

Table 4. Annual mass Balance of water, total alkalinity (TAlk) and dissolved inorganic carbon 855

(DIC) in the Waquoit Bay subterranean estuary 856

857

Water Flux# TAlk Flux DIC Flux

m3 m

-1 year

-1 mol m

-1 year

-1 mol m

-1 year

-1

STE Inputs

Terrestrial groundwater* 1200 276 509

Submarine groundwater re-

charge (from Waquoit Bay) 730 1305 1216

STE Output

Submarine groundwater dis-

charge (into Waquoit Bay) 1930 1468 2177

Net STE Production - 68 (5%) 419 (19%) # Water flux was modeled from Jan. 2005 to Dec.2005, and normalized to per meter of shoreline; 858 * Terrestrial groundwater TAlk is 0.19 mol m

-3 and DIC is 0.72 mol m

-3. 859

860

861

Figure 1 1

2

3

4

5

6

7

8

9

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11

12

13

14

15

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17

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20

21

22

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26

27

A

A’

11

A

Vineyard Sound

Waquoit Bay

A’

12 14 13

Berm

Bay CCC1 PZ10

32.9 m 13.5m

Terrestrial groundwater

Marine groundwater

time series

fixed-depth

well

PZ5 3m

Piezometer

Profiles

PZ3

STZ

5.5

m

6 m

5.5

m

A

Figure 2 28

29

30

31

32

33

34

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66

67

2005/1/1 2006/1/1 2007/1/1 2008/1/1 2009/1/1 2010/1/1 2011/1/1

MS

L a

nom

aly

(m

)

-0.15

-0.10

-0.05

0.00

0.05

0.10

0.15

0.20

Figure 3 68

69

70

71

72

73

74

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110

0.0

2.5

5.0

7.5

10.0

12.5

15.0

17.5

20.0

22.5

25.0

4/2010 5/2010 6/2010

Dep

th (

m)

5/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

Salinity

0

5

10

15

20

25

30

35

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

Dep

th (

m)

5/20104/2010 6/20105/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

0

10

20

30

40

50

60

70

80

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

4/2010

Dep

th (

m)

5/2010 6/20105/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

0

50

100

150

200

250

300

350

400

450

Dep

th (

m)

5/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 4/2010 5/2010 6/201012b/2009 1/2010 2/2010 3/2010

111

112

113

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146

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-4.5

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-1.5

-0.5

0

50

100

150

200

4/2010

Dep

th (

m)

5/2010 6/20105/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

4/2010

Dep

th (

m)

5/2010 6/20105/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

0

5

10

15

20

25

30

35

40

45

50

55

60

65

(a) DO (%)

(e) NH4

+

(μmol L-1

)

(d) N+N

(μmol L-1

)

(b) Mn

(μmol kg-1

)

Figure 4

(c) Fe

(μmol kg-1

)

147

148

149

150

151

152

153

154

155

156

157

158

159

160

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162

163

164

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-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

0

1

2

3

4

5

6

7

8

9

10

11

Dep

th (

m)

5/20104/2010 6/20105/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

(f) PO4 (μmol L-1

)

pCO2 (utam)

1000

2500

4000

5500

7000

8500

10000

11500

13000

14500

16000

17500

19000

20500

22000

4/2010 5/2010 6/2010

Dep

th (

m)

5/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

0

150

300

450

600

750

900

1050

1200

1350

1500

1650

1800

1950

Dep

th (

m)

4/2010 5/2010 6/20105/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

Figure 5 179

180

181

182

183

184

185

186

187

188

189

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191

192

193

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205

206

207

208

500

650

800

950

1100

1250

1400

1550

1700

1850

2000

4/2010 5/2010 6/2010

Dep

th (

m)

5/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

4/2010 5/2010 6/2010

Dep

th (

m)

5/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

5.2

5.4

5.6

5.8

6.0

6.2

6.4

6.6

6.8

7.0

7.2

7.4

(a) TAlk (μmol kg-1

)

(c) pCO2

(μatm)

(d) pH-Cal

(b) DIC (μmol kg-1

)

209

210

211

212

213

214

215

216

217

218

219

220

221

222

223

224

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226

227

228

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238

239

240

241

0.00

0.08

0.16

0.24

0.32

0.40

0.48

0.56

0.64

4/2010 5/2010 6/2010

Dep

th (

m)

5/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

(e) Omega Ca

Figure 6 242

243

244

245

246

247

248

249

250

251

TAlk ( mol kg-1

)

0 400 800 1200

DIC ( mol kg-1

)

400 800 1200 1600

TAlk

DIC

Mn ( mol kg-1

)

0 1 2 3 4

Fe ( mol kg-1

)

0 100 200 300 400 500

Fe

Mn

PO4

3- ( mol L

-1)

0.0 0.2 0.4 0.6 0.8 1.0

N+N ( mol L-1

)

0 5 10 15 20 25

NH4

+ ( mol L

-1)

0 50 100 150 200 250

PO43-

NH4+

N+N

Salinity

0.0 0.2 0.4 0.6 0.8 1.0

Dep

th (

m)

1

2

3

4

5

6

ORP (mv)

0 100 200 300

pH

5.2 5.6 6.0 6.4 6.8

Salinity

ORP

pH

(a) (b) (c) (d)

Figure 7 252

253

254

255

256

257

258

259

260

261

0 10 20 30

DIC

(m

ol

kg

-1)

0

500

1000

1500

2000

2500

3000

Salinity

0 10 20 30

DIC

(m

ol

kg

-1)

0

500

1000

1500

2000

2500

Salinity

0 10 20 30

TA

lk (

mol

kg

-1)

0

500

1000

1500

2000

2500

STE-PZ10

STE-PZ3

STE-PZ5

Baywater

0 10 20 30

TA

lk (

mol

kg

-1)

0

500

1000

1500

2000

2500

3000

STE-PZ5

STE-PZ10

STE-PZ3

Baywater

0.0 .1 .2 .3 .4 .5

0

200

400

600

800

1000 (a) (b)

(c) (d)

0.0 .1 .2 .3 .4 .5

600

800

1000

1200

1400

1600

Nov. 2009 Nov. 2009

Jun. 2010 Jun. 2010

TA

lk (

mol

kg

-1)

DIC

(

mol

kg

-1)

DIC

(

mol

kg

-1)

TA

lk (

mol

kg

-1)

4/2010 5/2010 6/2010

Dep

th (

m)

5/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

-800

-700

-600

-500

-400

-300

-200

-100

0

100

200

300

400

500

600

700

800

4/2010 5/2010 6/2010

Dep

th (

m)

5/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

-800

-700

-600

-500

-400

-300

-200

-100

0

100

200

300

400

500

600

700

800

Figure 8 262

263

264

265

266

267

268

269

270

271

272

273

274

275

276

277

278

279

280

281

282

283

284

285

286

bDIC (μmol kg-1

)

aTAlk (μmol kg-1

)

Figure 9 287

288

289

290

291

292

293

294

295

296

297

298

299

300

301

302

303

304

305

306

307

308

309

310

311

312

313

314

315

316

(c) DIC (mol kg-1

)

(a) Salinity (b) TAlk (mol kg-1

)

(d) Fraction of aerobic

respiration (%)

(e) Fraction of

denitrification (%) (f) Fraction of Mn and

Fe reduction (%)

Figure 10 317

318

319

320

321

322

323

324

325

326

327

328

329

330

331

332

333

334

335

336

337

338

339

340

5.0

5.5

6.0

6.5

7.0

7.5

8.0

0 10 20 30

pH

Salinity

observation-Nov.09

observation-Jun.10

conservative

mixing-Nov.09

conservative

mixing-Jun.10

Figure 11 341

342

343

344

345

346

347

348

349

350

351

Salinity

0 5 10 15 20 25 30 35

Fe (

mo

l k

g-1

)

0

50

100

150

Salinity

0 5 10 15 20 25 30 35

Mn

(m

ol

kg

-1)

0

20

40

60

80

100

0 5 10 15 20 25 30 35

NH

4

+ (

mo

l k

g-1

)

0

20

40

60

80

100

STE-Nov. 2009

Baywater-Nov. 2009

STE-Jun.2010

Baywater-Jun.2010

0 5 10 15 20 25 30 35P

O4

3- (

mo

l k

g-1

)

0

2

4

6

8

10

12

14(a) (b)

(c) (d)

243

439...

Mn

(

mol

kg

-1)

Fe

(m

ol

kg

-1)

NH

4+ (

mo

l k

g-1

)

PO

43- (

mol

kg

-1)

Figure 12 352

353

354

355

356

357

358

359

360

361

362

363

364

365

366

367

368

369

370

371

372

373

374

375

376

377

378

379

380

381

(a) Salinity

(c) DIC (mol kg-1

)

(e) Fraction of

denitrification (%)

(f) Fraction of Mn and

Fe reduction (%)

(d) Fraction of aerobic

respiration (%)

(b) TAlk (mol kg-1

)

Figure 13 382

-400 -200 0 200 400 600 800

-800

-600

-400

-200

0

200

400

600

Mid salinity-Jun.10

High salinity-Jun.10

Mid salinity-Nov.09

High salinity-Nov.09

TA

lk (

μm

ol

kg

-1)

DIC (μmol kg-1

)

Aerobic respiration (-0.2)

Denitrification (0.9)

Manganese reduction (4)

Iron reduction (8)

Fe/Mn oxidation

CaC

O 3 d

isso

lution

(2)

CaC

O 3 p

rodu

ctio

n (-

2)

(-∞)

DIC (mol kg-1

)

T

Alk

(

mo

l k

g-1

)