Groundwater dating

A small series of six samples contained ¹⁴C analysis, two of which originated from Carboniferous and Upper Devonian units at Maastricht (Glasbergen, Reference Glasbergen1985), with the others from Palaeogene units at Nieuweschans, Broekhuizenvorst, Zegge and Maarle (Table 3). Laboratory analysis of the activity of ¹⁴C yields the ‘apparent age’ of groundwater, because the activity of total inorganic carbon (TIC) is determined and the TIC concentration is the result of all hydrogeochemical processes that have taken place from the infiltration area to the monitoring point. A correction must be made to transform the analysis from apparent age into true age. This may be very complicated, depending on the hydrogeological settings, the hydrogeochemical processes that affected the carbonate chemistry along the groundwater flow path, and the related origin and age of mineralised solid organic carbon or dissolved inorganic carbon. We did not attempt to correct the apparent age to true age because these conditions are insufficiently clear for the individual samples.

Table 3.

Groundwater analyses (mg l⁻¹, except when stated otherwise) for samples from buried aquifers and the reservoirs in South-Holland that may include isotope analysis, plus the reference for groundwater close to rock salt in Gorleben, Germany (Kimpe, Reference Kimpe1963; Glasbergen, Reference Glasbergen1981, Reference Glasbergen1984, Reference Glasbergen1985; Van der Weiden, Reference Van der Weiden1983; Heederik, Reference Heederik1989; Eggenkamp, Reference Eggenkamp1994; Kloppmann et al., Reference Kloppmann, Négrel, Casanova, Klinge, Schelkes and Guerrot2001; Griffioen, Reference Griffioen2015).

^aCalculated from electroneutrality.

The laboratory analysis indicates that two samples had ¹⁴C activity below the detection limit of 2 PMC (per cent modern carbon). This yields an apparent age of at least 32,000 years according to the following formula:

$$\begin{equation*} {\rm{apparent\ age}} = \frac{{ - 5736\ln \left( {\frac{{{\rm{analysis}}}}{{100}}} \right)}}{{\ln 2}} \end{equation*}$$

where analysis refers to the ¹⁴C activity of the sample in per cent modern carbon. The other four samples had ¹⁴C activities of 0.4–9 PMC and thus apparent ages of 20,000−46,000 years. The true age will be lower, as radiogenic ¹⁴C becomes ‘diluted’ by ¹⁴C-free or ¹⁴C–poor solid reactants such as carbonate minerals or sedimentary organic matter (Clark & Fritz, Reference Clark and Fritz1997). The low ¹⁴C activities, however, make it unlikely that these groundwaters are younger than 10,000 years or thereabouts. As discussed later, few δ²H and δ¹⁸O analyses are available and two of them suggest rainwater recharge during a glacial period. Groundwater at the six sampled deep, confined aquifers is thus ‘old’, with values below 10 PMC. Comparable corrected ages were determined for groundwater in the Eocene aquifer in northwestern Flanders (Blaser et al., Reference Blaser, Coetsiers, Aeschbach-Hertig, Kipfer, Van Camp, Loosli and Walraevens2010), which is the extension of the Palaeogene aquifers in Zeeland.

Shallow, semi-confined aquifers

Figure 3 shows plots of the chloride vs sodium concentration for the shallow, semi-confined aquifers, where the seawater composition is indicated as reference. The Cl concentrations range from around 10 mg l⁻¹ to 15,000 mg l⁻¹ (except for one outlier of 24,800 mg l⁻¹ in Carboniferous rock in south Limburg): i.e. the chloride concentrations are lower than in seawater (19,000 mg l⁻¹). Fresh groundwater may thus be present when the pre-Neogene units lie within a few hundreds of metres below the surface, which can be related to infiltration of rainwater in outcrop areas or in the permeable units overlying them (Glasbergen, Reference Glasbergen1981). This is evidenced for groundwater in the Belgian Kempen by the presence of tritium (Beaucaire et al., Reference Beaucaire, Pitsch, Toulhoat, Motellier and Louvat2000). It is further illustrated in Figure 4 for groundwater in the Carboniferous unit in southern Limburg, where two groundwater types can be discerned: a shallower, fresh one and a deeper, saline one. Note that this area has been free of seawater influence since at least the Pliocene, i.e. for at least 5 Ma. The Cl concentrations for the four Belgian wells in the Palaeogene aquifer and close to the Belgian/Dutch border vary considerably − from 95 to 3357 mg l⁻¹ − which implies that the fresh/saline interface in groundwater immediately below the Rupel Clay does not coincide with this border.

Fig. 3.

Concentration plots of the compounds Na (above) and SO₄ (below) vs Cl for groundwater in the shallow, semi-confined aquifers. Note the different kinds of scale axes. See Figure 1 for a map with the geographical names used.

Fig. 4.

Chloride concentration vs depth for groundwater in the Carboniferous aquifer in the coal-mining district of southern Limburg (data obtained from Kimpe, Reference Kimpe1963).

Many samples plot above the seawater dilution line, which implies that they are enriched in Na relative to Cl (Fig. 3). This is probably due to cation exchange under freshening conditions. Fresh or brackish samples (Cl < 1000 mg l⁻¹) are also enriched in K and Mg (not shown), which provides further evidence for freshening of the aquifers. However, the saline samples (Cl > 1000 mg l⁻¹) are depleted in Mg, and the underlying reason is not obvious. They are also depleted in SO₄, which indicates sulphate reduction. Conversely, the fresh samples are enriched in SO₄ up to 2766 mg l⁻¹, which indicates pyrite oxidation in open-air conditions and/or dissolution of gypsum or anhydrite (which are commonly found in Dutch, Cretaceous and older formations; Wong et al., Reference Wong, Batjes and De Jager2007). The fresh SO₄–rich samples were from the Limburg mining district or Twente.

Buried, confined aquifers

Generally, samples from the buried, confined aquifers have higher ranges in concentrations for the main ions than those from the shallow, semi-confined ones, but lower ranges than those from the reservoirs (except alkalinity; Table 2). Figure 5 shows chloride vs sodium concentration for the buried aquifers as well as for the reservoirs. Both the seawater composition and brine as observed close to the Permian salt diapir in Gorleben, Germany (Kloppmann et al., Reference Kloppmann, Négrel, Casanova, Klinge, Schelkes and Guerrot2001), are plotted as references. We chose the latter as reference for dissolution of halite (‘halite brine’) although it is unknown how representative this composition is for dissolution of rock salt in the Netherlands. For the deep, buried aquifers, the chemical data indicate that groundwater is usually saline, with Cl between 1000 and 32,000 mg l⁻¹ (Table 2). However, exceptions occur towards brackish or hypersaline, reflecting inflow of meteoric groundwater on the one side and contact with brine or rock salt on the other. In such cases, the highest Cl concentration exceeds the seawater concentration by less than a factor 2 and the lowest concentration no longer refers to brackish water.

Fig. 5.

Chloride vs sodium concentration plot for groundwater in deep, buried aquifers or oil and gas reservoirs (one fresh, relatively shallow sample from central Limburg lies outside the graph).

Several relatively extensive analyses including isotopes are available for groundwater in buried, deep aquifers (Table 3; Fig. 6). A mixture of rainwater and seawater is indicated for the two aquifer samples from Nieuweschans, where the δ¹⁸O and δ²H values indicate that the fraction of rainwater varied from one-third to almost half. The other samples lie on the global meteoric water line (GMWL). Based on the few δ²H and δ¹⁸O isotope analyses (Fig. 6), it has been concluded (Glasbergen, Reference Glasbergen1984, Reference Glasbergen1985) that the origin of groundwater in these samples varies from rainwater infiltrated during a glacial period (at Maastricht), recent rainwater (at Broekhuizenvorst) and a mixture of rainwater and seawater (at Nieuweschans). For the latter, hydrological contact with a salt dome − and possibly evaporation too − also plays a role. For comparison, the isotope analyses for the Belgian semi-confined Palaeogene groundwater samples indicate a rainwater origin.

Fig. 6.

Bromide, Li, I and B concentrations vs chloride as present for a limited series of samples together with the results of isotope analysis of H₂O in some saline samples (the average for Dutch rainwater is shown for comparison, and GMWL stands for global meteoric water line). Samples with 100 µg Li l⁻¹ refer to analyses below detection limit, which must be 100 µg l⁻¹ or smaller. Seawater evaporation pathways according to data from Fontes & Matray (Reference Fontes and Matray1993a).

Few data on Br are present for the regular groundwater aquifers that do not serve as oil and gas reservoirs. The available Cl/Br data show that a series of Dutch samples lies on the seawater evaporation/dilution line (Fig. 6), including the two samples from the deep drilling in Asten that have Cl concentrations above that of seawater (and around 100 mg Br l⁻¹). This implies evaporation of seawater prior to infiltration. The samples from the Belgian Palaeogene aquifers lie on this line, too. The samples from the former Oranje Nassau 1 mine, Broekhuizenvorst, Nieuweschans and Arcen deviate markedly from the Cl/Br seawater ratio and maybe so do the two samples from Maastricht that have the lowest Cl concentrations of this subseries. This being the case, the Br concentration of 140 mg l⁻¹ vs the Cl concentration of 71,500 mg l⁻¹ for Nieuweschans is remarkable. The saline nature of this groundwater just below the Rupel Clay in the Dongen aquifer at 582 m depth has been attributed to rock salt dissolution (Glasbergen, Reference Glasbergen1981). This should normally yield a lower Br concentration, as indicated by the halite brine reference and two Dutch samples from diapirs (Fig. 6): possible explanations for this, in addition to an analytical error, are dissolution in evaporated seawater or local presence of Br-containing salt that typically precipitates in the last evaporative stages. The isotopic data of δ²H (around −25‰) and δ¹⁸O (around −3‰) suggest a mixture of seawater and rainwater. An enrichment in Br relative to Cl can be explained by mixing with a residual brine, but these isotopic data make it unlikely that mixing with a residual brine from which halite has precipitated plays a role. So does the palaeohydrological evolution of the Palaeogene aquifer from which this sample is collected. The sample from Broekhuizenvorst lies on the GMWL but not on the Cl/Br evaporation line, so, based on these similar lines of evidence, the high salinity of this sample must be attributed to rock salt (note that Broekhuizenvorst lies along the Meuse River in the southern part of the Netherlands). A similar explanation may hold for the other samples that deviate from the Cl/Br seawater ratio and are also located in the southern Netherlands.

Boron and lithium can also be useful tracers to characterise whether mixing with residual brine or dissolution of rock salt affects the groundwater composition. Figure 6 shows plots of the limited series of Li and B analyses vs Cl; Tables 2 and 3 provide specific data. The seawater evaporation line is also plotted where the evolution beyond halite precipitation is also depicted (as provided by Fontes & Matray, Reference Fontes and Matray1993a): a brine becomes enriched in these two species relative to Cl as soon as halite precipitates due to subaerial evaporation. The data for the buried aquifers show that most samples lie above the seawater evaporation/dilution line. This suggests in the first instance that mixing with residual brines plays a role. However, some of the samples are so much enriched that they lie above any mixing line irrespective of which evaporative stage is considered for the residual brine. This holds in particular for the samples from Asten that contain 42,000–50,000 µg B l⁻¹ and 2700–3000 µg Li l⁻¹ (Table 3). It is generally recognised that lithium and boron may be leached from the solid rock matrix, causing enrichment of these species above the seawater evaporation line in saline, subsurface waters (Collins, Reference Collins1969; Giblin & Dickson, Reference Giblin and Dickson1992; Kharaka & Hanor, Reference Kharaka, Hanor and Drever2004). Here, enhancement of temperature causes leaching of Li and B from rock as well as immobilisation processes for B. The samples from Asten have the largest sampling depth of 1600 m within the deep, buried aquifer group, and the high Li and B concentrations may thus be explained by the related, high temperature.

Close to 60 analyses that contain results on major ions on groundwater in buried aquifers were included in our dataset. Figure 5 shows that for this series, fresh groundwater is enriched in Na relative to Cl, and the molar ratio is about 1 to 1 when Cl is between 2000 and 70,000 mg l⁻¹ and above this concentration may be depleted in Na relative to Cl for the samples from the reservoirs. Figure 7 shows concentration plots for four other main compounds vs Cl. Near-identical patterns are observed for Ca, Mg and K in the buried aquifer samples: the data points lie above the seawater dilution line for fresh groundwater, and a near-linear pattern is observed up to Cl concentration of 70,000 mg l⁻¹ that lies above (Ca) or between (K, Mg) the seawater evaporation line and the halite dissolution lines. The two most saline samples of the buried aquifer data group refer to samples from salt diapirs (Zuurdeeg et al., Reference Zuurdeeg, Coenegracht, Mebius and Van der Weiden1984): they lie below the seawater evaporation and halite dissolution lines, where the latter has the halite brine from Gorleben (Germany) as reference composition (Kloppmann et al., Reference Kloppmann, Négrel, Casanova, Klinge, Schelkes and Guerrot2001).

Fig. 7.

Concentration plots of the main compounds SO₄, Ca, Mg and K vs Cl for samples from the deep, buried aquifers and the reservoirs at linear (left) and logarithmic scale (right). Seawater evaporation pathways according to data from Fontes & Matray (Reference Fontes and Matray1993a). The mixing line represents mixing between a freshwater type and halite brine.

A mixing line between fresh water and halite brine is also plotted in the log graphs for K and Mg (Fig. 7), where the fresh end-member contains 10 mg l⁻¹ as reference concentrations for Cl, K and Mg. Most samples lie above this mixing line, too. When the brackish and saline samples are the result of mixing between fresh water and seawater (evaporated or not) or halite brine, these samples are thus enriched in Na and also in Mg and K when halite brine is assumed as end-member, but are mostly depleted in Mg and K when seawater is assumed as end-member. Enrichments of these cations can be attributed to cation-exchange or weathering processes; depletions are solely attributable to cation exchange, as dolomite precipitation and precipitation of K-bearing minerals is unlikely at the in situ temperatures of most buried aquifers sampled (mostly below 40°C). The results for alkalinity and Ca suggest that cation exchange is an active process for these brackish and saline samples: Figure 8 illustrates that the alkalinity to Ca ratio on an equivalent basis is frequently substantially higher than 1, i.e. the ratio cannot be explained by congruent dissolution of Ca carbonate by which similar equivalent concentrations of Ca²⁺ and HCO₃ ⁻ become produced. It is worth pointing out that the alkalinity is frequently several hundreds of mg l⁻¹ in this data group and also in the shallow, semi-confined aquifer group (data not shown). As groundwater is usually saturated or supersaturated for calcite (Fig. 9) and the Lower Carboniferous, Cretaceous and Palaeogene aquifers are mostly marine, calcareous deposits, it is unrealistic to suppose that the concentrations of Na, K, Mg and Ca are controlled by silicate weathering where Ca silicates are barely present. Dissolution of gypsum or anhydrite, if present, is also unlikely as a control because SO₄ is mostly diminished compared to the seawater ratio (Fig. 7) and additional SO₄ reduction would not enhance the Ca concentration under Ca carbonate equilibrium. A combination of cation exchange and Ca carbonate precipitation/dissolution is more likely where the overall status is freshening. This may be less self-evident for the lacustrine and fluvial Late Carboniferous rocks that also contain coal but that may be less carbonate-rich. However, groundwater in the Carboniferous aquifers in south Limburg is mostly saturated in Ca carbonates and frequently has alkalinity to Ca equivalent ratios above 1.5. This groundwater thus has no general characteristics that differ from those of the other aquifers. Except for the small series of samples discussed earlier, there is no evidence on whether mixing with seawater (evaporated or not) or dissolution of rock salt is the major process that determines the brackish or saline nature of many samples in the buried, confined aquifers. The Cl to Br ratio for the small series suggests that both processes may be operational in deep, pre-Neogene aquifers in the Netherlands.

Fig. 8.

Alkalinity to Ca ratio vs the Cl concentration for samples from the three data groups.

Fig. 9.

Saturation indices of halite, anhydrite, calcite and siderite vs the Cl concentration of groundwater for samples from the three data groups. Note the different scales of the x/y-axes.

Oil and gas reservoirs

Table 2 shows that the oil and gas reservoirs have the highest ranges in formation water concentrations for the main ions except for alkalinity. The chloride concentrations for the reservoirs are presented in Figure 5, plotted against the sodium concentrations. The reservoirs in South-Holland, Broad Fourteens Basin (which lies off the coast of North-Holland and is oriented southeast−northwest) and northeastern Netherlands have Cl concentrations above 40,000 mg l⁻¹, i.e. more than twice as high as in seawater. Water containing such high Cl concentrations will henceforth be referred to as hypersaline. The highest concentrations observed equal the Cl concentrations close to the salt dome in Gorleben. Groundwater having more than 85,000 mg Cl l⁻¹ has only been encountered in the reservoirs in the northeastern Netherlands and deep (>2800 m) reservoirs in Broad Fourteens Basin and Central Offshore Saddle. Exceptional samples from the Carboniferous Limestone Group at 1700 m depth in Merksplas and 2250 m depth in Turnhout (B), have 60,000 and 80,300 mg Cl l⁻¹, respectively. No such hypersaline water has been observed elsewhere in the southern part of the Netherlands, not even in the reservoirs at about 900−1800 m depth in South-Holland. These reservoirs have a typical Cl concentration of 42,000−62,000 mg l⁻¹. Verweij (Reference Verweij2003) attributed this to a combination of two reasons: (1) absence of major halite deposits in the southern part of the Netherlands and (2) meteoric flushing during Tertiary and Quaternary times. Permian Zechstein salt is the most widespread and thickest halite deposit in the Netherlands. It is found in the northern half of the Netherlands (Fig. 1; Geluk, Reference Geluk, Wong, Batjes and de Jager2007a). Halite deposits are also known for the Triassic Röt, Muschelkalk and Keuper Formations (Upper Germanic Trias Group) and Late Jurassic Weiteveen Formation (Schieland Group) (Geluk, Reference Geluk, Wong, Batjes and de Jager2007b). The distribution is most widespread for the Röt Formation and overlaps considerably with that of Zechstein salt. The Muschelkalk Evaporite Member also contains halite more southward than the Zechstein: local occurrences are found in the West Netherlands Basin and in the Roer Valley Graben. In Broad Fourteens Basin, hypersaline groundwater is observed in both the Permian Rotliegend deposits below and the Lower Triassic and Lower Cretaceous deposits above the Permian Zechstein evaporites. All wells in the Central Offshore Platform refer to the Rotliegend deposits below: one well has 45,000 mg Cl l⁻¹ and three others approximately 150,000 mg Cl l⁻¹.

Sulphate ranges from below detection limit to 2000 mg l⁻¹, with one outlier. This outlier having 11,350 mg SO₄ l⁻¹ may be meaningful, as the electroneutrality condition for this sample is better when this concentration is considered. Up to 70,000 mg Cl l⁻¹, groundwater is unsaturated for both anhydrite and gypsum (Fig. 9). Thus calcium sulphate precipitation does not control the SO₄ concentration for these waters; the major influence is most likely reduction of SO₄, combined with mineralisation of organic matter.

There are several tracers that may provide insight into the hydrogeological mechanism responsible for the brine compositions (e.g. Fontes & Matray, Reference Fontes and Matray1993a, Reference Fontes and Matray1993b; Eastoe et al., Reference Eastoe, Long and Knauth1999; Kharaka & Hanor, Reference Kharaka, Hanor and Drever2004). The Cl to Br ratio may give a clue, because as a result of evaporation this is equal to the seawater weight ratio of 292 for brines, is lower for interstitial water when there has been precipitation of rock salt, and exceeds 1000 when natural rock salt has dissolved into flowing groundwater (Kloppmann et al., Reference Kloppmann, Négrel, Casanova, Klinge, Schelkes and Guerrot2001). Kloppmann et al. (Reference Kloppmann, Négrel, Casanova, Klinge, Schelkes and Guerrot2001) observed Br concentrations of 55−65 mg Br l⁻¹ for brines having 182,000−192,000 mg Cl l⁻¹ in the vicinity of the Gorleben salt dome and around 20−25 mg Br l⁻¹ for brines having 40,000−65,000 mg Cl l⁻¹. Two hypersaline Dutch samples from ‘a salt diapir’ fit into this picture, having Cl concentrations of 148,000 and 209,000 mg l⁻¹ (Fig. 6). Unfortunately, there are no data on Br for the brines in the oil and gas reservoirs except for wells Akkrum 13 in Friesland and Bergermeer 4 in North-Holland (Fig. 6). Akkrum 13 is hypersaline, with a Br concentration between that of evaporated seawater and halite brine, and the Bergermeer 4 data suggest a straightforward mixture of seawater and halite brine. The Cl to Br ratio from these two samples gives no clues about the origin of the brines.

δ²H and δ¹⁸O isotopic analyses of formation water are present for only four samples from South-Holland (Fig. 6; Van der Weiden, Reference Van der Weiden1983): the data vary around −3 for δ¹⁸O and −22 for δ²H, which suggests a mixture of seawater and rainwater. These four samples have a Cl concentration close to 60,000 mg l⁻¹. A third tracer is the isotopic ratio of stable Cl: the δ³⁷Cl of rock salt at evaporation of 0.0‰ seawater varies between −0.5 and +0.5‰, with the highest values associated with halite and the lower ones with other K and Mg salts (Eggenkamp et al., Reference Eggenkamp, Kreulen and Koster van Gross1995; Eastoe et al., Reference Eastoe, Long, Land and Kyle2001). The related evaporated brine has δ³⁷Cl values down to −0.9‰, where such a value is associated with the last stage of evaporation and precipitating K/Mg salts (Eggenkamp et al., Reference Eggenkamp, Kreulen and Koster van Gross1995; Eastoe et al., Reference Eastoe, Long and Knauth1999). Eggenkamp (Reference Eggenkamp1994) analysed eight samples that were studied by Van der Weiden (Reference Van der Weiden1983) too. Some samples were analysed for both H₂O and Cl isotopes (cf. Table 3). The δ³⁷Cl values vary between −1.77 and −0.46‰, with a good positive correlation (r² = 0.89) with the Cl concentration for six samples from sandstones, while two samples from sand/shale beds lie below this regression line. Eggenkamp (Reference Eggenkamp1994) concluded that the samples formed by mixing of interstitial water from a source rock and aquifer water. This explanation does not hold for the two sand/shale samples, for which ultrafiltration or diffusion with related isotopic fractionation provides a better explanation. Many mixing scenarios can be established between seawater (evaporated or not), fresh meteoric water, dissolution of rock salt and diffusion, to explain the observations of the δ²H, δ¹⁸O and δ³⁷Cl tracers for these reservoir samples from South-Holland. Note that this group of samples lies below the halite dissolution line for K and between this line and the seawater evaporation line for Mg. This might indicate that neither KCl salts nor evaporative brines play a role in their last stages, although K may also disappear in secondary reactions.

As discussed for the deep, buried aquifers, Li and B may also be useful tracers on the genesis of brines. Figure 6 shows their concentrations for ten samples from the South-Holland reservoirs and Tables 2 and 3 provide additional data. Unfortunately, no Br analysis was available for these samples. It is striking that, for B, the samples lie above the seawater evaporation line and, for Li, they lie below (where the samples with 100 µg Li l⁻¹ actually refer to analyses below unknown detection limit, which must be 100 µg l⁻¹ or smaller based on other samples). This different behaviour implies that water–rock interactions play a role for at least one of these solutes. The highest B concentrations around 40,000 µg l⁻¹ cannot be explained by mixing with any residual brine as B is too high for this compared to Cl. Conversely, the lower Li concentration can be explained by mixing with a brine from halite dissolution.

Important types of palaeogroundwater flow in the reservoirs are induced by sedimentary loading and topography-driven flow (Verweij et al., Reference Verweij, Simmelink, Van Balen and David2003). Sedimentary loading will probably expel residual brines from the evaporite deposits, and topography-driven flow may introduce fresh groundwater in the reservoir's aquifer under palaeoenvironmental conditions. The palaeohydrological and diagenetic evolution of the reservoirs in the southern North Sea Basin is very complicated and varies among the individual sub-basins (Verweij et al., Reference Verweij, Simmelink, Van Balen and David2003, Reference Verweij, Simmelink and Underschultz2011; Gaupp & Okkerman, Reference Gaupp and Okkerman2011). Diagenetic evidence exists that meteoric water entered the Permian Rotliegend reservoir in the northern part of the Netherlands during the Early Cretaceous Period, giving rise to flushing of this aquifer (Lee et al., Reference Lee, Aronson and Savin1989). Palaeohydrological evidence also exists that topography-driven flow could reach great depths in the Triassic and Lower Cretaceous reservoirs above the Zechstein evaporites during the Late Cretaceous Period in the northern Broad Fourteens Basin (Bouw & Oude Essink, Reference Bouw and Oude Essink2003). There is also evidence that the Permian Rotliegend reservoir is overpressurised under present-day conditions, with the degree of overpressurisation increasing from southeast to northwest under the southern North Sea (Verweij et al., Reference Verweij, Simmelink and Underschultz2011). This suggests the presence of compaction-driven flow and absence of topography-driven flow under current conditions for this reservoir. The palaeo-fluid movements are difficult to reconstruct: the situation is further complicated by the influx of groundwater and hydrocarbons from the deeper Carboniferous formations during the palaeoenvironmental evolution of the Rotliegend reservoir (Gaupp & Okkerman, Reference Gaupp and Okkerman2011). Whether residual brine is still present in this reservoir under the Zechstein evaporites cannot be ascertained from the available hydrochemical data. Dissolution of rock salt together with fractionating diffusion, however, seems a plausible phenomenon for the South-Holland reservoirs above the Zechstein evaporites.

Figure 7 shows that the Ca concentrations outrange the SO₄ concentrations by far for the most saline reservoir brines, when dissolution of gypsum or anhydrite as evaporites is assumed to be the principal source. The brines in the Broad Fourteens Basin usually have 43–2620 mg SO₄ l⁻¹ (except for one outlier), whereas the brines in South-Holland repeatedly have SO₄ below detection limit. It is also known that the natural gas may contain up to 7 ppm H₂S in the reservoirs in South-Holland (Core Lab, Reference Lab1987). This means that the gas is not classified as sour (criterion >1% H₂S; Lokhorst, Reference Lokhorst1998). With one exception, all SO₄ concentrations in saline or hypersaline groundwater lie below those of the seawater and the halite brine: the groundwater is always depleted in SO₄ by comparison with two end-member solutions, which is probably due to sulphate reduction and possibly anhydrite equilibrium for the reservoir formation waters having Cl above 100,000 mg l⁻¹ (Fig. 9). The first process is mediated microbially below 80°C and thermochemically above 100°C or even 160°C (Machel, Reference Machel2001). To reach 100°C, the corresponding depth must be about 3 km, assuming a thermal gradient of 30°C km⁻¹ (Verweij, Reference Verweij2003; Bonté et al., Reference Bonté, Van Wees and Verweij2012). The gas and oil fields in South-Holland lie at a depth of 900−1800 m, with corresponding temperatures around 30−60°C. One may thus wonder whether thermochemical SO₄ reduction could have played a role for these reservoirs, which would imply that the reduction of SO₄ was microbially controlled or the pore water originates from deeper. Conversely, biodegraded oils are present in the Lower Cretaceous reservoirs of South-Holland due to tectonic inversion which brought these reservoirs close to the surface at the onset of the Palaeogene (De Jager & Geluk, Reference De Jager, Geluk, Wong, Batjes and de Jager2007), so microbial reduction of SO₄ is plausible too.

Figure 7 also shows that the Ca concentrations are substantially higher than the K or Mg concentrations for the most saline groundwater in the reservoirs (Cl > 100,000 mg l⁻¹) and also exceed by far the seawater and halite brine concentrations (except for one outlier). They generally increase with the Cl concentration. For K, two clusters are found: one saline, which lies below the halite brine dissolution line and one hypersaline, which mostly lies between this line and the seawater evaporation line. With the exception of one outlier, the Mg concentrations also lie between these two lines. Most samples from the first cluster originate from the South-Holland reservoirs and those from the second originate mostly from the Broad Fourteens Basin.

Two mechanisms may give rise to brines in the oil and gas reservoirs of the North Sea Basin and elsewhere: mixing with connate water expelled from evaporites under compaction that had undergone evaporation beforehand (‘residual brines’) or dissolution of evaporite minerals such as rock salt. Egeberg & Aagaard (Reference Egeberg and Aagaard1989) attributed the brines in oil fields on the Norwegian shelf to the former process, and Verweij (Reference Verweij2003), following Hanor (Reference Hanor1994), attributed the brines in the deeper Dutch subsurface to the latter. For deep groundwater (having Cl concentrations up to 30,000 mg l⁻¹) in Upper Cretaceous Chalk in Zealand (Denmark), Bonnesen et al. (Reference Bonnesen, Larsen, Sonnenborg, Klitten and Stemmerik2009) also deduced an origin from mixing between residual brine and meteoric water. An alternative explanation proposed by Worden et al. (Reference Worden, Manning and Bottrell2006) for the high Cl concentrations of around 120,000 mg l⁻¹ in a Triassic oil-bearing sandstone in the Wessex Basin (southern England) was dissolution of Upper Triassic evaporites. Both mechanisms thus seem operational in northwestern Europe.

The behaviour of Ba and Sr is also relevant in brines because barite (BaSO₄) may precipitate and contribute to the cementing of grains (Gluyas et al., Reference Gluyas, Jolley and Primmer1997), with Sr co-precipitating, and diagenetic anhydrite may contain Ba and Sr (Purvis, Reference Purvis1992; Sullivan et al., Reference Sullivan, Haszeldine, Boyce, Rogers and Fallick1994). Both cations have frequently been analysed in the hypersaline solutions, and concentrations range up to a few thousand mg l⁻¹ for Sr and a few tens ofmg l⁻¹ for Ba (except for one extreme of 250 mg Ba l⁻¹). Figure 10 shows the observed Ba and Sr concentrations plotted against SO₄. An inverse relationship with a 1 to 1 slope at logarithmic scale can be recognised for the data points on the upper right-hand side. This could be indicative of an equilibrium relationship with a solubility control by SO₄ minerals for these two divalent cations. Supersaturation for barite (BaSO₄) is usually found for saline or hypersaline groundwater, and for celestite, also for the hypersaline groundwater, but the saturation index of celestite varies from −3 to 1 for saline groundwater.

Fig. 10.

Strontium (with solubility line for celestite; above) and Ba (with that for barite; below) concentrations vs the SO₄ concentration for a limited series of samples from the deep, buried aquifers and the reservoirs.

Role of diagenesis in reservoirs

The oil and gas reservoirs have undergone all kinds of diagenetic processes (Purvis, Reference Purvis1992; Gaupp & Okkerman, Reference Gaupp and Okkerman2011; Fischer et al., Reference Fischer, Dunkl, Von Eynatten, Wijbrans and Gaupp2012), which are partly related to these reservoirs having higher temperatures than the younger, buried Palaeogene aquifers. Speciation calculations using PHREEQC (Parkhurst & Appelo, Reference Parkhurst and Appelo2013) indicate that only the most hypersaline solutions are supersaturated for halite (Fig. 9). This shows that the production of dissolved Na⁺ and the immobilisation of Na⁺ are spatially and temporally decoupled, because otherwise halite saturation could be maintained. The hypersaline groundwater is also saturated or supersaturated for anhydrite (Fig. 9; and also gypsum). Strong supersaturation with respect to calcite and dolomite is found for many samples, whereas the saturation state for siderite varies (Fig. 9). The calculation is, however, probably imperfect, particularly due to sampling errors for pH, shortcomings in the thermodynamic calculations and slight errors in the way temperature was estimated. Precipitation of siderite and calcite occurs under more variable geochemical conditions, including low temperatures. Precipitation of dolomite may play a role during diagenesis in the oil- and gas-bearing formations such as the Rotliegend sandstone (Purvis, Reference Purvis1992; Gluyas et al., Reference Gluyas, Jolley and Primmer1997; Gaupp & Okkerman, Reference Gaupp and Okkerman2011). This may influence the Mg concentration in addition to the Ca concentration. Dolomite precipitation is more unlikely in shallower groundwater aquifers, as dolomite precipitation is slow at low temperature whereas calcite precipitation is not, and supersaturation of both carbonates usually happens (Arvidson & Mackenzie, Reference Arvidson and Mackenzie1999).

The more disperse concentration patterns for the four major cations above 85,000 mg Cl l⁻¹ by comparison with the patterns below 70,000 mg l⁻¹ have been attributed to dissolution of halite with minor secondary reactions for Na below 70,000 mg Cl l⁻¹ and important secondary reactions above. One mechanism that has been proposed (Land & Prezbindowski, Reference Land and Prezbindowski1981; Egeberg & Aagaard, Reference Egeberg and Aagaard1989) is albitisation of Ca feldspars in the reservoirs containing NaCl brine:

$$\begin{eqnarray*} &&{\rm{N}}{{\rm{a}}^ + } + {\rm{CaS}}{{\rm{i}}_2}{\rm{A}}{{\rm{l}}_2}{{\rm{O}}_8} + {{\rm{H}}_4}{\rm{Si}}{{\rm{O}}_4} + \raise.5ex\hbox{$\scriptstyle 1$}\kern-.1em/ \kern-.15em\lower.25ex\hbox{$\scriptstyle 2$} {\rm{ H^{+}}}\ \rightarrow {\rm{NaS}}{{\rm{i}}_3}{\rm{Al}}{{\rm{O}}_8} + {\rm{C}}{{\rm{a}}^{2 + }}\\ &&\quad + 2\,{{\rm{H}}_2}{\rm{O}} + \raise.5ex\hbox{$\scriptstyle 1$}\kern-.1em/ \kern-.15em\lower.25ex\hbox{$\scriptstyle 2$} \,{\rm{A}}{{\rm{l}}_2}{{\rm{O}}_3} + \raise.5ex\hbox{$\scriptstyle 1$}\kern-.1em/ \kern-.15em\lower.25ex\hbox{$\scriptstyle 2$} \,{\rm{O}}{{\rm{H}}^ - } \end{eqnarray*}$$

where Na⁺ originates from an evaporated seawater residue or dissolution of halite:

$$\begin{equation*} {\rm{NaCl}}\ \rightarrow {\rm{N}}{{\rm{a}}^ + } + {\rm{C}}{{\rm{l}}^ - } \end{equation*}$$

Egeberg & Aagaard (Reference Egeberg and Aagaard1989) argued that transformation of K feldspar is unlikely both thermodynamically and in relation to the depletion of K below the seawater evaporation line. Illitisation of the reservoirs is commonly observed and the temperature under which illite formed varies strongly. The source of the K-bearing fluid is still uncertain (Gaupp & Okkerman, Reference Gaupp and Okkerman2011): both non-correlation between feldspar dissolution and amount of authigenic clay, and increasing feldspar dissolution in association with precipitation of illite and/or kaolinite have been observed. Zechstein rock salt has also been inferred to be the source of the K fluid.

Calcium concentrations above the seawater evaporation line are observed in our samples. Weathering of Ca feldspars may be a plausible explanation, although the Rotliegend Sandstone in the southern part of the North Sea Basin is usually poor in Ca feldspars but not in K feldspar (e.g. Purvis, Reference Purvis1992; Gluyas et al., Reference Gluyas, Jolley and Primmer1997; Ziegler, Reference Ziegler2006). Diagenetic overgrowths of K and Na feldspars are even found in this reservoir rock (Gaupp & Okkerman, Reference Gaupp and Okkerman2011; Fischer et al., Reference Fischer, Dunkl, Von Eynatten, Wijbrans and Gaupp2012). The Ca concentrations are not balanced by SO₄ concentrations, which suggests in the first instance that neither gypsum nor anhydrite dissolution plays a role, whereas these evaporites are commonly found as precursors when halite as evaporite is also found. However, anaerobic oxidation of methane in association with SO₄ reduction has been observed at elevated temperatures in hydrothermal systems (Holler et al., Reference Holler, Widdel, Knittel, Amann, Kellermann, Hinrichs, Teske, Boetius and Wegener2011; Biddle et al., Reference Biddle, Cardman, Mendlovitz, Albert, Lloyd, Boetius and Teske2012; Adams et al., Reference Adams, Haorfrost, Bose, Joye and Girguis2013). These observations raise the question of whether transformation of anhydrite to calcite may happen in association with methane oxidation (Holler et al., Reference Holler, Widdel, Knittel, Amann, Kellermann, Hinrichs, Teske, Boetius and Wegener2011):

$$\begin{equation*} {\rm{CaS}}{{\rm{O}}_4} + {\rm{C}}{{\rm{H}}_4}\ \rightarrow {\rm{CaC}}{{\rm{O}}_3} + {{\rm{H}}_2}{\rm{S}} + {{\rm{H}}_2}{\rm{O}} \end{equation*}$$

Preliminary modelling using PHREEQC showed that this process may give rise to water high in Ca and low in SO₄, but there is no removal of Na. It must also be remembered that anhydrite is found in the Rotliegend sandstone not only as an early-stage evaporite mineral but also as a late- or intermediate-stage diagenetic mineral (Purvis, Reference Purvis1992; Lanson et al., Reference Lanson, Beaufort, Berger, Baradat and Lacharpagne1996; Gluyas et al., Reference Gluyas, Jolley and Primmer1997; Gaupp & Okkerman, Reference Gaupp and Okkerman2011). The environmental conditions for the intermediate diagenesis are, however, not clear: the effects of uplift or mixing of groundwater close to faults have not been elucidated. It is unknown whether conversion from calcite or aragonite to dolomite happened, which would not only explain enrichment of Ca above the seawater evaporation line but also depletion of Mg below. An entirely different explanation is that the seawater composition has varied over geological time, where it is, for example, assumed that Cretaceous seawater was enriched in Ca and depleted in Na, Mg and SO₄ relative to present-day seawater (Bonnesen et al., Reference Bonnesen, Larsen, Sonnenborg, Klitten and Stemmerik2009).

Trace element concentrations

Iodine and selenium are interesting in the context of subsurface disposal of radioactive waste. All the data available on these two elements refer to the deep, buried aquifers. Our dataset contains only seven analyses for Se: five of them have a concentration below the detection limit of 1−2 µg l⁻¹; the other two also have concentrations below 1 µg l⁻¹, which is not far from the natural concentration in shallow groundwater and is below the seawater concentration (Table 4). Thus, little can be said about Se in aquifers of Oligocene and older age. Thirteen analyses are available for iodine, with most samples referring to hypersaline water and all samples plotting above the seawater I/Cl ratio (Fig. 6). The I concentration varies by two orders of magnitude and is roughly related to salinity. These groundwater samples are also enriched relative to the seawater ratio. The two groundwater samples from the Palaeogene Rupel and Cretaceous Houthem Formations at Asten show high concentrations above 10 mg l⁻¹. The other samples originate from the Palaeogene Dongen aquifer at Nieuweschans (2.5/3.7 mg l⁻¹), the Carboniferous aquifers in Limburg (0.07−1.3 mg l⁻¹) and the salt diapirs (1.3−12 mg l⁻¹). It is generally accepted that iodine in groundwater mostly originates from degradation of predominantly marine plants, as the related organic compounds contain iodine, although the potential role of carbonates has often been overlooked (Heymann, Reference Heymann1925; Krul, Reference Krul1933; Geirnaert, Reference Geirnaert1973; Claret et al., Reference Claret, Lerouge, Laurioux, Bizi, Conte, Ghestem, Wille, Sato, Gaucher, Giffaut and Tournassat2010; Li et al., Reference Li, Wang, Guo, Xie, Zhang, Liu and Kong2014). The present observations confirm that in deep groundwater I is due more to mobilisation from the sediment than to a seawater origin, where marine sediments are often present in Palaeogene and older formations.

Table 4.

Statistical data on concentrations (μg l⁻¹) of trace elements in groundwater of Oligocene and older formations in the Netherlands compared to seawater (as derived from Hem, Reference Hem1970) and the Dutch natural background concentration limits for shallow groundwater (INS, 1997).

$ Van der Weiden (Reference Van der Weiden1983) did not provide a quantitative value for his detection limits; for Zn, neglecting these samples or assuming a detection limit below this median value yields a similar median value

The number of analyses for other trace elements varies, with as many as 28 for Li. Most analyses refer to the deep, buried aquifers and for Li, Zn, Cu and B to reservoirs in South-Holland as well. Trace metals in groundwater are relevant with respect to environmental and human toxicity. Table 4 presents a statistical summary of the data. Considerable numbers of analyses lie below the detection limits, which sometimes vary considerably. The maximum concentrations observed are usually at least one order of magnitude larger than the related seawater concentrations, which can be up to 200 times higher (Se) than the Dutch background concentration limit for shallow, fresh groundwater or down to 50 times lower (Cr, Pb). Most remarkable are Pb and Zn, which exceed the seawater concentration or national background concentration limit by more than a factor of 1000. They are followed by Cr, Cu and I. The median concentrations usually lie within a factor of 5 from both the national background concentration limit and the seawater concentration. An exception is I, with a factor 30 above the seawater concentration. Lithium, boron and fluoride show high concentrations in an absolute sense, but this is common in seawater as well as in Dutch shallow saline groundwater of marine origin (Frapporti et al., Reference Frapporti, Vriend and Van Gaans1996).

A comparison with other field studies on oilfield brines indicates that the medians in our data are similar to those found by Rittenhouse et al. (Reference Rittenhouse, Fulton III, Grabowski and Bernard1969) for Co, Cr, Cu, Ni, Sn, Ti and V. Collins (Reference Collins1969) found higher maximum concentrations for Li and B concentrations, and Rittenhouse et al. (Reference Rittenhouse, Fulton III, Grabowski and Bernard1969) found higher median concentrations for Li. Our maxima for Pb, Zn and Cu (found in hypersaline reservoir samples) are extremely high, as Kharaka and Hanor (Reference Kharaka, Hanor and Drever2004) state that the concentrations of heavy metals in oilfield waters are generally below 100 µg l⁻¹ because they are limited by the low solubilities of their respective sulphide minerals. These high concentrations are assumed to originate from redbed sands, which have a terrestrial origin and little reduced sulphur (Carpenter, Reference Carpenter1989; Kharaka & Hanor, Reference Kharaka, Hanor and Drever2004). Aqueous complexing of heavy metals with Cl may further raise the solubility of their sulphides. Coetsiers et al. (Reference Coetsiers, Blaser, Martens and Walraevens2009) present natural background concentrations for brackish and saline (Na + Cl > 1000 mg l⁻¹) groundwater in Eocene marine aquifers in Flanders, which are defined as the 90-percentile value, i.e. close to the maximum concentration found. Their values lie within a factor 5 from our median values, except for B, which in their study is only 2.34 µg l⁻¹. Their 90-percentile values are thus much lower than our maximum concentrations, whereas the deposits refer to comparable geological conditions. Conti et al. (Reference Conti, Sacchi, Chiarle, Martinelli and Zuppi2000) report high maximum concentrations for a few trace elements (As, Se, V, Sb) in saline groundwater originating from evaporated seawater trapped in the Late Messinian.