What Seawater Is Really Made Of

The six ions that make the sea salty

Dissolved salts in seawater are dominated by a handful of conservative ions — species whose ratio to total salt is fixed because they are chemically inert on the timescale of ocean mixing (~1,000 years). At the reference salinity of 35‰, the roster is:

• Chloride (Cl⁻) — 19.35 g/kg, 546 mmol/kg. The single largest constituent, and the one salinity was historically pegged to.
• Sodium (Na⁺) — 10.78 g/kg, 469 mmol/kg.
• Sulfate (SO₄²⁻) — 2.71 g/kg, 28 mmol/kg.
• Magnesium (Mg²⁺) — 1.28 g/kg, 53 mmol/kg.
• Calcium (Ca²⁺) — 0.41 g/kg, 10.3 mmol/kg.
• Potassium (K⁺) — 0.40 g/kg, 10.2 mmol/kg.

Together these six supply 99.4% of all dissolved solids. Add bicarbonate (HCO₃⁻, ~0.14 g/kg), bromide (Br⁻), borate, strontium, and fluoride and you reach 99.9%. Note that the ocean is not simply concentrated NaCl: it holds far more Mg²⁺ than a table-salt solution would, and the sulfate is a legacy of pyrite weathering and volcanic SO₂. The solution is electrically neutral — total cation charge (~605 meq/kg) balances total anion charge to within measurement error, a constraint used to check analyses.

Constant proportions: the Marcet–Dittmar rule

In 1819 Alexander Marcet proposed, and in 1884 William Dittmar confirmed by analyzing 77 water samples from the Challenger expedition, that the relative proportions of the major ions are essentially the same throughout the open ocean, regardless of the absolute salinity. Evaporation in the subtropics and dilution by rain or rivers change how much salt is present, but not the recipe.

The physical reason is a mismatch of timescales. The ocean mixes end-to-end in roughly 1,000 years via the thermohaline circulation. But the residence time of a conservative ion — the total amount in the ocean divided by its annual input — is measured in millions of years:

• Na⁺: τ ≈ 55–75 million years
• Cl⁻: τ ≈ 87–100 million years
• Mg²⁺: τ ≈ 13 million years
• Ca²⁺: τ ≈ 1 million years (short, because organisms constantly pull Ca²⁺ out as CaCO₃)

Because each ion is stirred thousands of times before it is removed, local river inputs get homogenized into a globally uniform blend. The rule breaks down only where a specific process is faster than mixing — Ca²⁺ and HCO₃⁻ deviate slightly in surface waters where plankton build shells, and in isolated basins like the Black Sea. Salinity is therefore not measured ion-by-ion; measuring one conservative constituent fixes them all.

How salinity is actually measured

Nobody evaporates and weighs seawater routinely — it is slow and drives off volatile chloride. Historically, oceanographers measured chlorinity (Cl‰) by titrating a sample with silver nitrate:

• Ag⁺ + Cl⁻ → AgCl(s)↓ (Mohr / Knudsen titration, endpoint by chromate indicator)

Then applied the empirical Knudsen relation Salinity = 1.80655 × Chlorinity. Since 1978, salinity is instead defined by the Practical Salinity Scale (PSS-78): a purely electrical quantity comparing the sample's conductivity to a standard KCl solution at 15°C and 1 atm. Practical salinity has no units by definition (the old ‰ / ppt is a colloquial shorthand). Modern CTD instruments log Conductivity, Temperature, and Depth thousands of times per cast; conductivity is corrected for temperature and pressure, then converted to salinity. The current standard, TEOS-10, goes further to Absolute Salinity (g/kg) using the true dissolved mass, because subtle regional differences in composition — silica in the deep Pacific, for instance — make conductivity slightly misrepresent real mass.

Dissolved gases and the carbonate system

Seawater is also a solution of atmospheric gases, governed by Henry's Law (solubility ∝ partial pressure) and rising sharply with falling temperature — cold polar water holds far more gas. N₂ dominates by inertness (~10–15 mL/L), but the reactive gases run the biosphere. O₂ ranges from saturation (~5–8 mL/L) at the sunlit surface down to near-zero in oxygen minimum zones. The chemically richest system is inorganic carbon (DIC ≈ 2.0–2.3 mmol/kg), set by a chain of equilibria:

• CO₂(g) ⇌ CO₂(aq)
• CO₂(aq) + H₂O ⇌ H₂CO₃
• H₂CO₃ ⇌ H⁺ + HCO₃⁻ (pK₁* ≈ 6.0 in seawater)
• HCO₃⁻ ⇌ H⁺ + CO₃²⁻ (pK₂* ≈ 9.1 in seawater)

At the ocean's pH of ≈ 8.1, that speciation lands between the two pKₐ values, so DIC is ~90% bicarbonate (HCO₃⁻), ~9% carbonate (CO₃²⁻), and less than 1% dissolved CO₂. This mixture is a natural buffer: added H⁺ is mopped up by CO₃²⁻ → HCO₃⁻, which is why seawater pH is so stable. It is also why the ocean has absorbed roughly one-quarter to one-third of all anthropogenic CO₂ — about 170 billion tonnes of carbon since 1750.

Where the salt came from — and where it goes

The ocean is not salty because rivers keep adding salt forever; it is at a steady state where inputs match removals. The sources and sinks differ ion by ion:

• Chemical weathering of continental rock delivers Na⁺, Ca²⁺, K⁺, Mg²⁺, HCO₃⁻ and dissolved silica via rivers. Example: albite feldspar hydrolysis, 2NaAlSi₃O₈ + 2CO₂ + 11H₂O → Al₂Si₂O₅(OH)₄ + 2Na⁺ + 2HCO₃⁻ + 4H₄SiO₄.
• Volcanic and hydrothermal input. Cl⁻ is largely 'excess volatile' — degassed from the mantle as HCl rather than weathered from rock. Mid-ocean-ridge hot springs strip Mg²⁺ out of seawater and add Ca²⁺, Li, Fe, and Mn.
• Biological removal. Plankton fix Ca²⁺ + 2HCO₃⁻ → CaCO₃(s) + CO₂ + H₂O and pull Si into opal (SiO₂·nH₂O), sinking both to the seafloor. This is why Ca²⁺ has such a short residence time.
• Evaporite burial and reverse weathering remove Na⁺, Cl⁻, SO₄²⁻, and K⁺ into halite, gypsum, and clays over geologic time.

Because removal responds to concentration, the system is self-correcting: bulk seawater salinity has stayed within a factor of ~2 for hundreds of millions of years, even as CO₂, oxygen, and Mg/Ca ratios (which flip the ocean between 'aragonite' and 'calcite' seas) drifted substantially.

Trace elements, nutrients, and the parts-per-billion ocean

Below the majors lies a periodic table of trace elements, most at nanomolar to picomolar levels, whose distributions are shaped by biology and scavenging rather than by dilution. Oceanographers sort them by profile shape:

• Nutrient-type (nitrate, phosphate, silicate, and micronutrient iron and zinc): depleted at the surface where plankton consume them, enriched at depth where sinking matter remineralizes. The classic Redfield ratio of the organic matter driving this is C:N:P ≈ 106:16:1.
• Conservative trace elements (Li, Rb, Mo, U): flat profiles, long residence times, behaving like the majors.
• Scavenged-type (Al, Mn, Pb): removed onto sinking particles, so concentrations fall with depth and residence times are short (decades to centuries).

Iron is the celebrity: at ~0.05–2 nanomolar, it is so scarce that it limits phytoplankton growth across roughly a third of the surface ocean (the 'high-nutrient, low-chlorophyll' regions of the Southern Ocean and equatorial Pacific). The absolute salt content barely notices these species, but they govern how much life the ocean supports and how much carbon it buries — proof that seawater's chemistry matters far beyond the 35 grams you can weigh.

The six major ions of seawater at salinity 35‰ — and how differing residence times explain the ranking.

Ion	Concentration (g/kg)	% of salts by mass	Residence time
Chloride Cl⁻	19.35	55.0%	~87–100 Myr
Sodium Na⁺	10.78	30.6%	~55–75 Myr
Sulfate SO₄²⁻	2.71	7.7%	~9–12 Myr
Magnesium Mg²⁺	1.28	3.7%	~13 Myr
Calcium Ca²⁺	0.41	1.2%	~1 Myr
Potassium K⁺	0.40	1.1%	~6–12 Myr